Highly Siderophile and Strongly Chalcophile Elements in High-Temperature Geochemistry and Cosmochemistry 9781501502095, 9780939950973

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REVIEWS in MINERALOGY and GEOCHEMISTRY Volume 81

2016

Highly Siderophile and Strongly Chalcophile Elements in High-Temperature Geochemistry and Cosmochemistry EDITORS Jason Harvey

University of Leeds, UK

James M.D. Day

Scripps Institution of Oceanography, USA

Front-cover: Proton-induced X-ray emission maps of experimental charges of Fe–Ni–Cu–S. The experimental liquid, quenched at 1000 °C, precipitated an intergrowth of quenched dendritic monosulfide solution [(FeNi)S] and intermediate solid solution [(FeCu)S] grains. Scale 2.5 × 1.8 mm. Image courtesy of Sarah-Jane Barnes, Université du Québec à Chicoutimi, Québec, Canada. Back-cover: The analytical consideration for the analysis of highly siderophile and strongly chalcophile elements in high-temperature geochemistry and cosmochemistry. Figure courtesy of Thomas Meisel, Montanuniversität, Leoben, Austria.

Series Editor: Ian Swainson MINERALOGICAL SOCIETY of AMERICA GEOCHEMICAL SOCIETY

Reviews in Mineralogy and Geochemistry, Volume 81

Highly Siderophile and Strongly Chalcophile Elements in High-Temperature Geochemistry and Cosmochemistry ISSN 1529-6466 ISBN 978-0-939950-97-3

Copyright 2016

The MINERALOGICAL SOCIETY of AMERICA 3635 Concorde Parkway, Suite 500 Chantilly, Virginia, 20151-1125, U.S.A. www.minsocam.org www.degruyter.com The appearance of the code at the bottom of the first page of each chapter in this volume indicates the copyright owner’s consent that copies of the article can be made for personal use or internal use or for the personal use or internal use of specific clients, provided the original publication is cited. The consent is given on the condition, however, that the copier pay the stated per-copy fee through the Copyright Clearance Center, Inc. for copying beyond that permitted by Sections 107 or 108 of the U.S. Copyright Law. This consent does not extend to other types of copying for general distribution, for advertising or promotional purposes, for creating new collective works, or for resale. For permission to reprint entire articles in these cases and the like, consult the Administrator of the Mineralogical Society of America as to the royalty due to the Society.

Highly Siderophile and Strongly Chalcophile Elements in High-Temperature Geochemistry and Cosmochemistry

81

Reviews in Mineralogy and Geochemistry

81

FROM THE SERIES EDITOR It has been a pleasure working with the two volume editors and authors on this 81st volume of Reviews in Mineralogy and Geochemistry. Several chapters have associated supplemental materials that can be found at the MSA website. Any future errata will also be posted there. Ian P. Swainson, Series Editor Vienna, Austria November 2015

PREFACE INTRODUCTION TO HIGHLY SIDEROPHILE AND STRONGLY CHALCOPHILE ELEMENTS IN HIGH TEMPERATURE GEOCHEMISTRY AND COSMOCHEMISTRY In high-temperature geochemistry and cosmochemistry, highly siderophile and strongly chalophile elements can be defined as strongly preferring metal or sulfide, respectively, relative to silicate or oxide phases. The highly siderophile elements (HSE) comprise Re, Os, Ir, Ru, Pt, Rh, Pd, and Au and are defined by their extreme partitioning (> 104) into the metallic phase, but will also strongly partition into sulfide phases, in the absence of metal. The HSE are highly refractory, as indicated by their high melting and condensation temperatures and were therefore concentrated in early accreted nebular materials. Within the HSE are the platinumgroup elements (PGE), which include the six elements lying in the d-block of the periodic table (groups 8, 9, and 10, periods 5 and 6), i.e., Os, Ir, Ru, Pt, Rh and Pd. These six elements tend to exist in the metallic state, or bond with chalcogens (S, Se, Te) or pnictogens (P, As, Sb, Bi). Rhenium and Au do not necessarily behave as coherently as the PGE, due to their differing electronegativity and oxidation states. For these reasons, a clear definition between the discussion of the PGE and the HSE (PGE, Re and Au) exists in the literature, especially in economic geology, industrial, or bio-medical studies. The strongly chalcophile elements can be considered to include S, Se, and Te. These three elements are distinguished from other chalcophile elements, such as Cd or Pb, because, like the HSE, they are all in very low abundances in the bulk silicate Earth (Fig. 1). By contrast with the 1529-6466/16/0081-0000$00.00

http://dx.doi.org/10.2138/rmg.2016.81.0

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface

Figure 1. Elements plotted as a function of atomic number versus their abundance in the bulk silicate Earth normalized to the CI-chondrite composition. Highly siderophile elements (HSE), which tend to exist in the metallic state, and strongly chalcophile elements are shown as dark and light gray symbols, respectively. Omitted naturally occurring elements with atomic numbers of < 92 are H, O, and the noble gases (He, Ne, Ar, Kr, Xe). Data are from McDonough and Sun (1995).

HSE, S, Se, and Te all have far lower melting and condensation temperatures, classifying them as highly volatile elements (Table 1). Moreover, these elements are not equally distributed within chondrite meteorite groups (Fig. 2). Since their initial distribution in the Solar nebula, planetary formation and differentiation process have led to large fractionations of the HSE and strongly chalcophile elements, producing a range of absolute and relative inter-element fractionations. The chemical properties of the HSE, that set them apart from any other elements in the periodic table (Table 2), have made them geochemical tracers par excellence. As tracers of key processes, the HSE have found application in virtually all areas of the physical Earth sciences.

Figure 2. Average Te, Se, S and Os abundances in carbonaceous (CI, C2, CM2, CV3) and ordinary (H, L, LL) chondrite groups. For scaling purposes, S contents are divided by 1000. Data are from compilations in Wang and Becker (2013) and Day et al. (2016).

iv

High Volatile Chalcophile

S

Se

Ru

Rh

Pd

Te

Re

Os

Ir

Pt

Au

Sulphur

Selenium

Ruthenium

Rhodium

Palladium

Tellurium

v

Rhenium

Osmium

Iridium

Platinum

Gold

Mod Volatile Siderophile

Refractory Siderophile

Refractory Siderophile

Refractory Siderophile

Refractory Siderophile

High Volatile Chalcophile

Transitional Siderophile

Refractory Siderophile

Refractory Siderophile

High Volatile Chalcophile

Behavior

Element

79

78

77

76

75

52

46

45

44

34

16

Z

1337

2041

2719

3306

3458

723

1828

2236

2606

494

388

Melting Point (K)

1060

1408

1603

1812

1821

709

1324

1392

1551

697

664

50% Tc (K)

0.146

1.004

0.47

0.486

0.037

2.33

0.588

0.141

0.692

19.7

54100

CI-chondrite abundance

0.001

0.0071

0.0032

0.0034

0.00028

0.012

0.0039

0.0009

0.005

0.075

250

BSE abundance

> 6000 y bp

1735

1803

1803

1925

1783

1803

1803

1844

1817

Prehistory

Discovery date

Unknown

Known to native Americans - A de Ulloa (South America)

S Tennant (England)

S Tennant (England)

W Noddack, I Tacke, O Berg (Germany)

FJM von Reichstein (Romania)

WH Wollaston (England)

WH Wollaston (England)

KK Klaus (Russia)

JJ Berzelius (Sweden)

Unknown

Discoverers (location)

Table 1. General behavior and discovery of the highly siderophile and strongly chalcophile elements.

Anglo-Saxon word Gold, latin word Aurum (Au)

Spanish Platina, meaning silver

Greek goddess Iris, meaning rainbow

Greek Osme, meaning odour

Latin Rhenus, meaning Rhine

Latin Tellus, meaning Earth

After the asteroid Pallas, Greek Pallas, Goddess of Wisdom

Greek Rhodon, meaning Rose

Latin Ruthenia, meaning Russia

Greek Selene,meaning Moon

Latin Sulphur, meaning 'burning stone'

Meaning of the name

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface These elements have been used to inform on the nucleosynthetic sources and formation of the Solar System, planetary differentiation, late accretion addition of elements to planets, coreformation and possible core-mantle interaction, crust-mantle partitioning, volcanic processes and outgassing, formation of magmatic, hydrothermal and epithermal ore deposits, ocean circulation, climate-related events, weathering, and biogeochemical cycling. More recently, studies of strongly chalcophile elements are finding a similar range of applications. Their utility lies in the fact that these elements will behave as siderophile or strongly chalcophile elements under reducing conditions, but will also behave as lithophile or atmophile elements under oxidizing conditions, as experienced at the present day Earth’s surface. A key aspect of the HSE is that three long-lived, geologically useful decay systems exist with the HSE as parent (107Pd–107Ag), or parent–daughter isotopes (187Re–187Os and 190 Pt–186Os). This volume is dedicated to some of the processes that can be investigated at high-temperatures in planets using the HSE and strongly chalcophile elements.

Table 2. Chemical properties, isotopes, atomic abundances and geologically important isotope decay schemes of HSE and strongly chalcophile elements (Continued on next page). Element

Electronegativity*

Oxidation states

Density (g/cm3)

Isotopes

Mass

Atomic Abundance

S

2.58

6,4,2,−2

2.067

32

31.9721

94.99

Se

Ru

2.55

2.2

6,4,−2

8,6,4,3,2,0,−2

4.809

12.1

Rh

2.28

5,4,3,2,1,0

12.4

Pd

2.2

4,2,0

12

vi

33

32.9714

0.75

34

33.9679

4.25

36

35.9671

0.01

74

73.9225

0.89

76

75.9192

9.37

77

76.9199

7.63

78

77.9173

23.77

80

79.9165

49.61

82

81.9167

8.73

96

95.9076

5.54

98

97.9053

1.87

99

98.9059

12.76

100

99.9042

16.60

101

100.9056

17.06

102

101.9043

31.55

104

103.9054

18.62

103

102.9055

100

102

101.9056

1.02

104

103.9040

11.14

105

104.9051

22.33

106

105.9035

27.33

108

107.9039

26.46

110

109.9052

11.72

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface Table 2 (Cont'd). Chemical properties, isotopes, atomic abundances and geologically important isotope decay schemes of HSE and strongly chalcophile elements. Element

Electronegativity*

Oxidation states

Density (g/cm3)

Te

2.1

6,4,−2

6.232

Re Os

1.9 2.2

7,6,4,2,1 8,6,4,3,2,0,−2

Isotopes

20.8 22.58

Mass

Atomic Abundance

120

119.9040

0.09

122

121.9031

2.55

123

122.9043

0.89

124

123.9028

4.74

125

124.9044

7.07

126

125.9033

18.84

128

127.9045

31.74

130

129.9062

34.08

185

184.9530

37.4

187

186.9557

62.6

184

183.9525

0.02

186

185.9538

1.59

Partial decay product 190Pt

187

186.9557

1.96

Partial decay product 187Re

188

187.9558

13.24

189

188.9581

16.15

190

189.9584

26.26

192

191.9615

40.78 37.3

Ir

2.2

6,4,3,2,1,0,−1

22.42

191

190.9606

193

192.9629

62.7

Pt

2.2

4,2

21.46

190

189.9599

0.014

192

191.9610

0.782

194

193.9627

32.967

195

194.9648

33.832

196

195.9649

25.242

198

197.9679

7.163

197

196.9665

100

Au

2.4

5,4,3,2,1,−1

19.282

Comments

4.16 × 10+10 yr−1

4.5 × 10+11 yr−1

Pauling Units, where E(AB) = [E(AA)·E(BB)]1/2 + 96.48(XA − XB)2. E(AB) is expressed in kJmol−1 (1 eV = 96.48 kJmol−1) and XA − XB represents the difference in "electronegativity" between the two elements, whose individual electronegativities are given the symbols XA and XB. Comments denote geologically and cosmochemically relevant decay schemes. Data source: http://ww.rsc.org/periodic-table.

*

While this volume is not dedicated to the practical applications of the HSE and strongly chalcophile elements, it would be remiss not to briefly discuss the importance of these elements in society. All of these elements have found important societal use (Table 3), from the application of Au as a valued commodity in early societies, through to the present-day; the importance of S and Se in biological processes; the discovery and implementation of Pt, Pd, and subsequently other PGE to catalytic oxidation (Davy 1817), and the importance of the anti-cancer drug cisplatin (cis-[Pt(NH3)2Cl2]) to anti-tumour treatments (e.g., Rosenberg et al. 1969). The use of the PGE, most especially Pt, Pd and Rh, in the automotive industry to generate harmless gases has caused some potential collateral effects; the possible environmental impact vii

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface

Se

S

Element

Chip resistors and electrical contacts; electrochemical cells for Cl production; catalysts for production of ammonia and acetic acid; hardener for Pt and Pd alloys; jewelry

Additive to glass; pigment for ceramics, paints, plastics; photovoltaic and photoconductive; dandruff shampoos; essential in many biological processes

Production of H2SO4 (Sulfuric Acid) for fertilizers, polyamides; fungicide; fumigant; vulcanization of natural rubber; gunpowder; bleaching; essential in many biological processes

Major, current societal uses

12,000

2,000,000

>69,000,000,000

World yearly production (in kg)

14,000

610

500

Estimated cost pure metal (USD/kg)

Table 3. Societal uses and current (numbers current between 2008–2013) production of the HSE and strongly chalcophile elements.

Ru

240

130,000

115,000

16,000

18,600

Used in alloys of copper or stainless steel, to improve machinability; additive in lead, making it more resistant to acid, improving strength and hardness; semiconductor applications

50,000

77,000

Catalytic convertors; catalyst for nitric acid, acetic acid and hydrogenation reactions; headlight reflectors; thermocouple elements; optical mirrors; optic fibre coatings

Te

Additive to W- and Mo-based alloys for filaments; electrical contact material; additive to nickel alloys for single-crystal turbine blades; hydrogenation of fine chemicals

500

42,000

Rh

Re

Used to produce very hard alloys for fountain pen tips, instrument pivots, needles and electrical contacts; catalyst in chemical processes

10,000

130,000

58,330

Os

Most corrosion-resistant material. Special alloys; forms alloy with Os, used for pen tips and compass bearings; used in making the standard metre bar (90% Pt–10% Ir); contacts in spark plugs

200,000

55,400

250,000

Ir

Catalytic convertors; jewelry; catalyst for nitric acid, benzene, silicon production; electronic industry; chemotherapy; optical fibres; LCDs; turbine blades; spark plugs; pacemakers; dental fillings

2,770,000

Catalytic convertors; jewelry; dental fillings and crowns; ceramic capacitors; gas hydrogenation/dehydrogenation reactions

Pt

Bullion; jewelry; alloys; dentistry; electronic components; gold nanoparticles used as industrial catalysts (vinyl acetate, used to make PVA for glue/paint/resin, is made using a gold catalyst)

Pd

Au

viii

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface and human health-risks from available PGE in the environment (see Rauch and Morrison 2008, for a review). An entire volume can (and should!) equally be written on the utility of the HSE and strongly chalcophile elements during low-temperature geochemistry.

BASIC CONCEPTS AND TERMINOLOGY Of the eight HSE, only Rh and Au are monoisotopic. Important data on Au and/or Rh abundances can be generated using non-isotope dilution methodologies, but the over-whelming majority of abundance data discussed in this volume for the HSE are for Re, Os, Ir, Ru, Pt, and Pd, which can be measured using isotope-dilution methodologies. Abundances of the strongly chalcophile elements, S, Se, and Te, can also be measured by isotope dilution. Isotope dilution studies to obtain abundances tend to use isotopically enriched tracers (typically 34S, 77Se, 99Ru, 105 Pd, 125Te, 185Re, 190Os, 191Ir, 194Pt), followed by inductively coupled plasma mass spectrometry measurement. Isotopic studies of Os are described below and methods for determination of isotopic variations in other HSE and strongly chalcophile elements have been developed, but are not explicitly discussed here. The reader is referred to Meisel and Horan (2016, this volume), and associated references for details of these isotopic methodologies. In addition to their strongly siderophile tendencies, HSE exhibit contrasting behaviors during melting, with the platinum-PGE (PPGE; Pt, Pd: melting temperature < 2000 °C; Barnes et al. 1985), Re and Au typically being more incompatible during melting and crystallization, relative to the iridium–PGE (IPGE; Os, Ir, Ru: melting temperature > 2000 °C; Barnes et al. 1985). For this reason, studies of the cosmochemical behavior of the HSE will often list the HSE in order of melting temperature of the pure metal, whereas studies using the HSE to investigate mantle melting processes will order the HSE according to relative incompatibility during melting. For mantle peridotites, there is general agreement that bulk peridotite-melt partition coefficients follow the sequence (e.g., Fischer-Gödde et al. 2011; Wang and Becker 2013; König et al. 2014): Os ≥ Ru ≥ Ir > Rh ≥ Pt > Pd > Au ≥ Te > Se > S ≥ Re The ability of relative and absolute HSE abundances to record recent processes acting on rocks are complemented by the existence of the long-lived 190Pt–186Os (190Pt → 186Os + α + Q; λ = 1.48 × 10−12 a−1; Walker et al. 1997) and 187Re–187Os (187Re − 187Os + β− + n; λ = 1.6668 × 10−11 a−1; Selby et al. 2007) chronometers. Both long-lived radiogenically produced isotopes are minor constituents (186Os = 1.6%; 187Os = 1.5%; Shirey and Walker, 1998) of osmium. In the case of the 187Re–187Os system, where 187Re is a major isotope (62.6%) of rhenium, and has a half-life of 41.6 Ga, the range of natural materials spans several orders of magnitude and 187Os/188Os can reasonably range from a Solar System initial ratio of ~ 0.095 to nearly pure 187Os derived from samples essentially devoid of Os and with high concentrations of Re (e.g., molybdenite; Luck and Allègre 1982). This characteristic means that the percent-level difference of 187Os/188Os between natural samples allows routine analysis of low Os abundance samples to percent precision or better, with the most widelyused method of analysis being negative thermal ionisation mass spectrometry (N-TIMS; Creaser et al. 1991; Völkening et al. 1991). The generally accepted ‘chondritic composition’ for 187Os/188Os is 0.127 (Shirey and Walker 1998), although there are clear differences between carbonaceous chondrites (187Os/188Os = ~ 0.1262), relative to ordinary (187Os/188Os = ~ 0.1284), or enstatite chondrites (187Os/188Os = ~ 0.1280; Day et al. 2016, this volume). Chondritic evolution is established from the most primitive initial 187Os/188Os defined from early Solar System iron meteorites (initial 187 Os/188Os = 0.09531) to the average chondritic composition for the present day. For these parameters, the average 187Re/188Os of chondrites is 0.40186. To calculate the 187Os/188Os of ix

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface chondrites at any time in the past—or future—the following equation can be used: 187

Os/188Ostime = 187Os/188Osinitial +

187

Re/188Oschondrite (eλ (

4568000000 )

– eλ t )

where λ is equal to 1.6668× 10−11 a−1 (Selby et al. 2007). For ease of reference, studies will often report the percentage difference between the Os isotope composition of a samples and the average ‘chondritic’ composition for a specified time, gOs. Samples with positive gOs are often described as ‘enriched’, because it implies long-term elevated 187Re/188Os with respect to chondrites. Samples with negative gOs are often described as ‘depleted’, due to the opposite implication of long-term low 187Re/188Os, in the following way:

((

γ Os =

187

Os/188Ossample(t ) /

187

) )

Os/188Oschondrite(t ) -1 × 100

Model ages (MA) and relative rhenium depletion ages (RD) ages can all be calculated using the Re–Os isotope system (Fig. 3). Model ages (TMA) represent the timing of separation from chondritic evolution and can be estimated for low Re/Os mantle materials, as well as high Re/ Os melts or crustal materials. The assumption with this method is that the Re/Os measured in the sample is an accurate reflection of its long-term history and has not been affected by later processes:   187 Os/188Oschondrite −187 Os/188Ossample   TMA = 1/λ × ln   187 + 1  188 187 188      Re/ Oschondrite − Re/ Ossample  

By contrast, time of relative Re depletion ages (TRD), which apply to low Re/Os mantle peridotites, does not rely on the Re/Os measured in the sample, which can be affected by recent Re addition. Instead this method uses sample compositions at the time of eruption and assumes that all of the Re in the sample was removed during melt-depletion. In reality, this method provides a minimum age for samples that have experienced melt-depletion: TRD=

  187 Os/188Oschondrite −187 Os/188Ossample      + 1 1/λ × ln    187  Re/188Oschondrite   

Due to potential disturbance from terrestrial weathering, or from cosmic-ray exposure affecting Re isotopic composition, studies of meteorites and planetary rocks have used Re*, which is the concentration of Re calculated assuming chondritic 187Os/188Os at the assumed time of sample crystallization (Day et al. 2010). This notation can be calculated as:

Re* =

{([Os]

sample

/ At wt sample

)×(

187

Os/188 Ossample −187 Os/188 Oschondrite(t )

)}

0.00336(ln[ Age × λ ] − 1)

By contrast with the 187Re–187Os decay system, 190Pt is a minor isotope of Pt (0.01292%) and has a longer half-life (~ 450 Ga), so 186Os/188Os variations in the mantle are small and of the order of ~ 0.00015%, with an ‘average’ mantle value of 0.119837 ± 5 (2σ). The typically minor variations of 186Os/188Os in volcanic settings require external analytical precision of better than 30 ppm. To obtain sufficient analytical precision, large quantities of Os are needed (typically 50–75 ng of Os) to generate sufficient signals on 186Os given the ionization efficiency of Os by N-TIMS (~ 2–6%; Creaser et al. 1991). Inevitably, the analytical challenge of measuring 186 Os/188Os means that there is far less data currently available than there is for 187Os/188Os. The now-extinct 107Pd was also the parent for 107Ag in the early Solar System. This extinct radionuclide system had a half-life of 6.5 Ma (107Pd – 107Ag + β− + n; λ = 1.06638 × 10−7 a−1; Parrington et al. 1996). The 107Pd–107Ag parent–daughter isotopic decay system is a candidate for use in both constraining the timing of early planetary fractionation events, for potentially x

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface

Figure 3. Solar System chronology versus 187Os/188Os. Shown in (a) is the chondritic evolution curve, curves corresponding to gOs values of +10 and +20, and the slopes of 187Re/188Os of 0.4 (~chondritic), 0.04 (sub-chondritic) and 200 (supra-chondritic). Examples of calculation of model ages and relative Re depletion ages are shown in (b). Example a shows an example of mantle-derived basaltic lava that evolved with high 187Re/188Os, which, when corrected, indicates a model-age of ~2.9 Ga. Example b shows a peridotite that experienced melt depletion. Based solely on the 187Os/188Os of the sample, a minimum TRD age of ~ 3.6 Ga is obtained. If the Re/Os of the sample is considered to be representative and not affected by later addition of Re, then the TMA can be calculated, giving a substantially older age of ~4 Ga. Example c, shows the case of a depleted peridotite xenolith that experienced infiltration of Re from the host melt, corrected for Re-ingrowth up to the eruption age (~ 0.25 Ga). In this scenario, the TRD age can be calculated from the initial 187Os/188Os at the time of the eruption age, but the lack of knowledge of the Re/Os of the xenolith prior to entrainment in the melt prevents a TMA from being calculated.

xi

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface determining whether Earth’s core material is incorporated into mantle plumes, and for investigating the timing of volatile-element depletion in planets. The relatively short half-life renders the system sensitive to fractionation events occurring within the first 40 million years of Solar System history (i.e., Kelly and Wasserburg 1978). Because Pd is more siderophile than Ag, planetary differentiation should result in an enrichment of Pd relative to Ag in planetary cores. If this happened during the lifetime of 107Pd, a correspondingly high 107Ag core signature would develop. If Earth’s differentiation occurred within 40 million years (approximately five half-lives) of the beginning of the Solar System, an isotopic excess of 107 Ag should exist within the core. Equally, because Ag is a moderately volatile element, whereas Pd is more refractory than Ag, large ranges in Pd/Ag have been observed in volatiledepleted iron meteorites (up to 100,000), compared with a Solar Pd/Ag of ~ 3, leading to 107 Ag/109Ag ratios > 9, compared with the solar value of 1.079 (Chen and Wasserburg 1996). For the Pd–Ag isotope system, the initial 107Pd/108Pd has been determined as 5.9 ± 2.2 × 10−5 (Schönbächler et al. 2008), with 107Ag/109Ag typically reported in parts per ten thousand notation relative to the NIST SRM978a silver standard (107Ag/109Ag = 1.07976): ε107 Ag =

( 

107

)

Ag/109 Agsample /107 Ag/109 Ag NIST SRM978a  -1 × 10,000

FUNDAMENTAL PROCESSES AND OUTLINE OF THE VOLUME In this volume, a number of key areas are reviewed in the use of the HSE and strongly chalcophile elements to investigate fundamental processes in high-temperature geochemistry and cosmochemistry. It is divided into five parts. The first part of the volume concerns measurements and experiments. Chapter 1, by Brenan et al. (2016), provides an comprehensive overview of experimental constraints applied to understanding HSE partitioning under a range of conditions, including: liquid metal–solid metal; metal– silicate; silicate–melt; monosulfide solid solution (MSS)–sulfide melt; sulfide melt–silicate melt; silicate melt–aqueous fluid–vapor. Chapter 2, by Meisel and Horan (2016) provides a summary of analytical methods, issues specifically associated with measurement of the HSE, and a review of important reference materials. The second part of the volume concerns the cosmochemical importance of the HSE and strongly chalcophile elements. In their assessment of nucleosynthetic isotopic variations of siderophile and chalcophile elements in Solar System materials, Yokoyama and Walker (2016, Chapter 3) discuss some of the fundamentals of stellar nucleosynthesis, the evidence for nucleosynthetic anomalies in pre-Solar grains, bulk meteorites and individual components of chondrites, ultimately providing a synthesis on the different information afforded by nucleosynthetic anomalies of Ru, Mo, Os, and other siderophile and chalcophile elements. Chapter 4 concerns the HSE in terrestrial bodies, including the Earth, Moon, Mars and asteroidal bodies for which we have materials as meteorites. Day et al. (2016) provide a summary of HSE abundance and 187Os/188Os variations in the range of materials available and a synthesis of initial Solar System composition, evidence for late accretion, and estimates of current planetary mantle composition. The third part of the volume concerns our understanding of the Earth’s mantle from direct study of mantle materials. In Chapter 5, Aulbach et al. (2016) discuss the importance and challenges associated with understanding HSE in the cratonic mantle, providing new HSE alloy solubility modelling for melt extraction at pressures, temperatures, fO and fS pertaining to conditions of cratonic mantle lithosphere formation. Luguet and Reisberg (2016) provide similar constraints on non-cratonic mantle in Chapter 6, emphasizing the importance of combined geochemical and petrological approaches to fully understand the histories of mantle peridotites. The information derived from studies of Alpine peridotites, obducted ophiolites and oceanic abyssal peridotites are reviewed in Chapter 7 by Becker and Dale (2016). xii 2

2

Highly Siderophile and Strongly Chalcophile Elements ‒ Preface The fourth part of the volume focusses on important minerals present in the mantle and crust. Chapter 8 provides a broad overview of mantle chalcophiles. In this chapter, Lorand et al. (2016) emphasise that chalcophile and siderophile elements are important tracers that can be strongly affected by host minerals as a function of sulfur-saturation, redox conditions, pressure, temperature, fugacity of sulfur, and silicate melt compositions. Along a similar theme in Chapter 9, O’Driscoll and Gonzalez-Jimenez (2016) provide an overview of platinum-group minerals (PGM), pointing out that, where present PGM dominate the HSE budget of silicate rocks. Finally in this section, Harvey et al. (2016) examine the importance of Re–Os–Pb isotope dating methods of sulfides for improving our understanding of mantle processes (Chapter 10). The fifth and final part of the volume considers the important of the HSE for studying volcanic and magmatic processes. In Chapter 11, Gannoun et al. (2016) provide a synthesis of the most abundant forms of volcanism currently operating on Earth, including mid-ocean ridge basalts, volcanism unassociated with plate boundaries, and subduction zone magmatism. The volume is completed in Chapter 12 by Barnes and Ripley (2016), by an appraisal of the obvious importance of magmatic HSE ore formation in Earth’s crust.

ACKNOWLEDGEMENTS We are grateful to all of the authors involved in the production of this Reviews in Mineralogy and Geochemistry volume. The task of putting together the volume would not have been so easy were it not for the diligence of the reviewers, and most especially, the Series Editor, Ian Swainson.

REFERENCES Aulbach S, Mungall J, Pearson DG (2016) Distribution and processing of highly siderophile elements in cratonic mantle lithosphere. Rev Mineral Geochem 81: 239–304 Barnes S-J, Ripley EM (2016) Highly siderophile and strongly chalcophile elements in magmatic ore deposits. Rev Mineral Geochem 81:725–774 Barnes S-J, Naldrett AJ, Gorton MP (1985) The origin of the fractionation of platinum-group elements in terrestrial magmas. Chem Geol 53:302–323 Brenan JM, Bennett NR, Zajacz Z (2016) Experimental results on fractionation of the highly siderophile elements (HSE) at variable pressures and temperatures during planetary and magmatic differentiation. Rev Mineral Geochem 81:1–87 Chen JH, Wasserburg GJ (1996) Live 107Pd in the early Solar System and implications for planetary evolution. In: Earth Processes: Reading the Isotope Code. Basu A, Hart SR (eds) Am Geophys Union Monogr 95:1–20 Coursey JS, Schwab DJ, Tsai JJ, Dragoset RA (2010) Atomic weights and isotopic compositions (version 3.0). National Institute of Standards and Technology, Gaithersburg, MD Creaser RA, Papanastassiou DA, Wasserburg GJ (1991) Negative thermal ion mass spectrometry of osmium, rhenium, and iridium. Geochim Cosmochim Acta 55:397–401 Becker H, Dale CW (2016) Re–Pt–Os isotopic and highly siderophile element behavior in oceanic and continental mantle tectonite. Rev Mineral Geochem 81:369–440 Davy H (1817) Some new experiments and observations on the combustion of gaseous mixtures, with an account of a method of preserving a continued light in mixtures of inflammable gases and air without flame. Phil Trans R Soc 107:77–85 Day JMD, Walker RJ, James OB, Puchtel IS (2010) Osmium isotope and highly siderophile element systematics of the lunar crust. Earth Planet Sci Lett 289:595–605 Day JMD, Brandon AD, Walker RJ (2016) Highly siderophile elements in Earth, Mars, the Moon, and asteroids. Rev Mineral Geochem 81:161–238 Emsley J (2011) Nature’s Building Blocks: An A–Z Guide to the Elements. Oxford University Press, New York, 2nd Edition Fischer-Gödde M, Becker H, Wombacher F (2011) Rhodium, gold and other highly siderophile elements in orogenic peridotites and peridotite xenoliths. Chem Geol 280:365–383

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Highly Siderophile and Strongly Chalcophile Elements ‒ Preface Gannoun A, Burton KW, Day JMD, Harvey J, Schiano P, Parkinson I (2016) Highly siderophile element and Os isotope systematics of volcanic rocks at divergent and convergent plate boundaries and in intraplate settings. Rev Mineral Geochem 81:651–724 Harvey J, Warren JM, Shirey SB (2016) Mantle sulfides and their role in Re–Os and Pb isotope geochronology. Rev Mineral Geochem 81:579–649 Haynes WM (2015) CRC Handbook of Chemistry and Physics. Taylor and Francis, Boca Raton, FL, 95th Edition, Internet Version 2015 Lorand J-P, Luguet A (2016) Chalcophile and siderophile elements in mantle rocks: Trace elements controlled by trace minerals. Rev Mineral Geochem 81:441–488 Luck J-M, Allègre CJ (1982) The study of molybdenites through the 187Re–187Os chronometer. Earth Planet Sci Lett 61:291–296 Luguet A, Reisberg L (2016) Highly siderophile element and 187Os signatures in non-cratonic basalt-hosted peridotite xenoliths: Unravelling the origin and evolution of the post-Archean lithospheric mantle. Rev Mineral Geochem 81:305–367 König S, Lorand J-P, Luguet A, Pearson DG (2014) A non-primitive origin of near-chondritic S–Se–Te ratios in mantle peridotites; implications for the Earth's late accretionary history. Earth Planet Sci Lett 385:110–121 Kelly WE, Wasserburg GJ (1978) Evidence for existence of 107Pd in the early Solar System. Geophys Res Lett 5:1079–1082 McDonough WF, Sun SS (1995) The composition of the Earth. Chem Geol 120:223–253 Meisel T, Horan MF (2016) Analytical methods for the highly siderophile elements. Rev Mineral Geochem 81:89–106 O’Driscoll B, González-Jiménez J (2016) Petrogenesis of the platinum-group minerals. Rev Mineral Geochem 81:489–578 Parrington JR, Knox HD, Breneman SL, Baum EM, Feiner F (1996). Nuclides and isotopes. 15th ed. San Jose, CA: General Electric Nuclear Energy Rauch S, Morrison GM (2008) Environmental relevance of the platinum-group elements. Elements 4, 259–263 Rosenberg B, VanCamp L, Trosko JE, Mansour VH (1969) Platinum compounds: a new class of potent antitumour agents. Nature 222:385–386 Schönbächler M, Carlson RW, Horan MF, Mock TD, Hauri EH (2008) Silver isotope variations in chondrites: volatile depletion and the initial 107Pd abundance of the Solar System. Geochim Cosmochim Acta 72:5330–5341 Selby D, Creaser RA, Stein HJ, Markey RJ, Hannah JL (2007) Assessment of the 187Re decay constant by cross calibration of Re–Os molybdenite and U–Pb zircon chronometers in magmatic ore systems. Geochim Cosmochim Acta 71:1999–2013 Shirey SB, Walker RJ (1998) The Re–Os isotope system in cosmochemistry and high-temperature geochemistry. Ann Rev Earth Planet Sci 26:423–500 Völkening J, Walczyk T, Heumann KG (1991) Osmium isotope ratio determinations by negative thermal ionization mass spectrometry. Int J Mass Spectrom 105:147–159 Walker RJ, Morgan JW, Beary E, Smoliar MI, Czamanske GK, Horan MF (1997) Applications of the 190Pt–186Os isotope system to geochemistry and cosmochemistry. Geochim Cosmochim Acta 61:4799–4808 Wang ZC, Becker H (2013) Ratios of S, Se and Te in the silicate Earth require a volatile-rich late veneer. Nature 499:328–331 Yokoyama T, Walker RJ (2016) Nucleosynthetic isotope variations of siderophile and chalcophile elements in the Solar System. Rev Mineral Geochem 81:107–160

Jason Harvey, University of Leeds James M. D. Day, Scripps Institution of Oceanography xiv

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TABLE OF CONTENTS

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Experimental Results on Fractionation of the Highly Siderophile Elements (HSE) at Variable Pressures and Temperatures during Planetary and Magmatic Differentiation James M. Brenan, Neil R. Bennett, Zoltan Zajacz

INTRODUCTION AND SCOPE..............................................................................................1 SOLID METAL–LIQUID METAL PARTITIONING..............................................................2 EXPERIMENTAL APPROACH TO SOLID METAL–LIQUID METAL PARTITIONING (DSM/LM) ..........................................3 HSE SOLUBILITY EXPERIMENTS: IMPLICATIONS FOR METAL–SILICATE PARTITIONING ........................................13 CALCULATING THE METAL–SILICATE MELT PARTITION COEFFICIENT FROM SOLUBILITY DATA ............................................14 Controls on the metal–silicate partition coefficient .....................................................15 Metal inclusions in experiments and the analysis of contaminated phases .................16 Possible mechanisms of metal inclusion formation ....................................................17 Experimental methods to measure HSE solubility and metal–silicate partitioning ....21 Summary of experimental data ....................................................................................22 Application of results to core formation......................................................................33 SILICATE AND OXIDE CONTROL ON HSE FRACTIONATION .....................................35 Experimental approach ................................................................................................35 Spinel–melt partitioning of HSEs................................................................................36 Silicate mineral–melt partitioning of HSEs.................................................................38 Origin of the variation in partitioning .........................................................................40 Local PGM saturation during chromite growth...........................................................43 MAGMATIC SULFIDE AND ASSOCIATED PHASES .......................................................44 Experimental approach ................................................................................................45 MSS–sulfide melt partitioning.....................................................................................48 MSS–ISS–sulfide melt partitioning .............................................................................51 xv

Highly Siderophile and Strongly Chalcophile Elements ‒ Table of Contents Sulfide melt–silicate melt and MSS–silicate melt partitioning ...................................51 Role of the chalcogens (Se, Te, As, Bi, Sb) .................................................................61 SILICATE MELT–AQUEOUS LIQUID–VAPOR PARTITIONING ....................................63 Theoretical considerations ...........................................................................................64 Experimental methods .................................................................................................65 The volatile/melt partitioning of Au ............................................................................68 The volatile/melt partitioning of PPGE .......................................................................71 The volatile/melt partitioning of IPGE and Re............................................................72 CONCLUDING REMARKS ..................................................................................................73 ACKNOWLEDGMENTS.......................................................................................................73 REFERENCES .......................................................................................................................74

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Analytical Methods for the Highly Siderophile Elements Thomas Meisel, Mary F. Horan

INTRODUCTION ..................................................................................................................89 DATA QUALITY CONSIDERATIONS FOR THE HSE.......................................................90 Sample heterogeneity and reproducibility. ..................................................................90 MEASUREMENT PROCEDURES .......................................................................................92 Chemical separation of HSE .......................................................................................95 REFERENCE MATERIALS FOR HSE ANALYSIS .............................................................97 APPENDIX ...........................................................................................................................100 REFERENCES .....................................................................................................................102

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Nucleosynthetic Isotope Variations of Siderophile and Chalcophile Elements in the Solar System Tetsuya Yokoyama, Richard J. Walker

INTRODUCTION ................................................................................................................107 ORIGIN OF ELEMENTS: STELLAR NUCLEOSYNTHESIS ..........................................109 Production of elements from He to Fe via hydrogen to silicon burning ...................109 PRESOLAR GRAINS ..........................................................................................................117 Types of presolar grains and their origin ...................................................................119 ISOTOPE ANOMALIES OF SIDEROPHILE AND CHALCOPHILE ELEMENTS IN BULK METEORITES ...............................................................................................126 Isotope anomalies of siderophile elements in bulk meteorites ..................................128 Isotope anomalies of chalcophile elements in bulk meteorites .................................138 INTERNAL ISOTOPE ANOMALIES PRESENT IN CHONDRITES ...............................140 CAIs...........................................................................................................................140 Acid leachates and residues .......................................................................................142 Isotopic constraints on the s-process nucleosynthetic component. ...........................148 ORIGIN OF PLANETARY SCALE ISOTOPE ANOMALIES IN METEORITES .......................................................................150 CONCLUDING REMARKS ................................................................................................152 ACKNOWLEDGMENTS.....................................................................................................152 REFERENCES .....................................................................................................................153 xvi

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Highly Siderophile Elements in Earth, Mars, the Moon, and Asteroids James M.D. Day, Alan D. Brandon, Richard J. Walker

INTRODUCTION ................................................................................................................161 MOTIVATION FOR STUDY AND BEHAVIOR OF THE HSE IN PLANETARY MATERIALS ............................................................................163 METHODS APPLIED TO INVESTIGATING SIDEROPHILE ELEMENTS IN PLANETARY MATERIALS ...............................................................164 HSE ABUNDANCES ...........................................................................................................164 The rhenium–osmium, platinum–osmium and palladium–silver isotope systems ....166 Standardization in planetary studies ..........................................................................168 Metal-sulfide–silicate modeling in chondritic systems .............................................168 Partial melt modeling of planetary mantles ...............................................................172 “Pristinity” of crustal and mantle samples ................................................................174 Estimation of planetary mantle composition .............................................................175 What do chondritic or nearly/broadly chondritic actually mean? .............................179 PLANETARY MATERIALS ................................................................................................181 Early Solar System materials.....................................................................................181 Fragments of planetary cores and/or metal-rich melt pools ......................................187 Primitive achondrite meteorites .................................................................................191 Meteorites from differentiated asteroids....................................................................195 Mars ...........................................................................................................................198 The Moon ..................................................................................................................201 Terrestrial mantle composition ..................................................................................209 Secondary alteration effects.......................................................................................210 PLANETARY FORMATION PROCESSES ........................................................................212 Initial conditions and homogeneity of starting materials ..........................................212 Partial melting and partitioning of the HSE ..............................................................214 Core crystallization....................................................................................................216 ‘Late-accretion’ .........................................................................................................216 Alternative hypotheses for the abundances of the HSE in planetary mantles ...........218 Magmatic processes...................................................................................................220 Later impacts into planetary crusts ............................................................................221 COMPARATIVE PLANETOLOGY AND IMPLICATIONS FOR TERRESTRIAL FORMATION .............................................................................222 FUTURE DIRECTIONS ......................................................................................................226 ACKNOWLEDGEMENTS ..................................................................................................227 REFERENCES .....................................................................................................................227

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Distribution and Processing of Highly Siderophile Elements in Cratonic Mantle Lithosphere Sonja Aulbach, James E. Mungall, D. Graham Pearson

INTRODUCTION ................................................................................................................239 THE CRATONIC MANTLE SAMPLE: PECULIARITIES, OPPORTUNITIES AND PITFALLS .............................................................................................................240 xvii

Highly Siderophile and Strongly Chalcophile Elements ‒ Table of Contents DATABASE ..........................................................................................................................244 MINERALOGY AND HSE HOSTS IN CRATONIC MANTLE PERIDOTITES ....................................................................248 ANALYTICAL TECHNIQUES FOR CRATONIC MANTLE PERIDOTITES ..................250 Whole rocks and mineral separates ...........................................................................250 Single-grain techniques .............................................................................................252 UTILIZATION OF THE RE–OS ISOTOPE SYSTEM IN CRATONIC MANTLE STUDIES.............................................................................253 EFFECT OF MELT DEPLETION DURING CRATONIC LITHOSPHERE ........................... STABILIZATION ON SULFUR AND HSE SYSTEMATICS ......................................255 Sulfur and the persistence of sulfides ........................................................................255 Alloy saturation .........................................................................................................256 Chalcogens ................................................................................................................262 HSE PROCESSING DURING MANTLE METASOMATISM ...........................................262 Modification during intraplate mantle metasomatism ...............................................263 Modification during craton margin processes—subduction ......................................269 MODELLING OF PRIMARY VS. SECONDARY HSE SIGNATURES IN CRATONIC MANTLE ..............................................................................................273 HSE Concentrations of the Archean convecting mantle (ACM) ...............................273 Modelling Rationale ..................................................................................................275 Effects of partial melt extraction on HSE based on modelling .................................278 Post-core formation, sluggish downward mixing of a late veneer ............................285 SUMMARY ..........................................................................................................................286 ACKNOWLEDGMENTS.....................................................................................................288 REFERENCES .....................................................................................................................288

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Highly Siderophile Element and 187Os Signatures in Non-cratonic Basalt-hosted Peridotite Xenoliths: Unravelling the Origin and Evolution of the Post-Archean Lithospheric Mantle Ambre Luguet, Laurie Reisberg

INTRODUCTION ................................................................................................................305 CONSTRAINING THE HSE AND 187Os/188Os ISOTOPIC COMPOSITION OF THE PRIMITIVE BULK SILICATE EARTH .........................................................306 PETROLOGY AND LOCATION OF NON-CRATONIC PERIDOTITE XENOLITHS ...........................................................308 A BRIEF REVIEW OF HSE AND Os ISOTOPE ANALYTICAL METHODS AND HSE NORMALIZATION VALUES ...............................................................................309 HOST MINERALS OF HIGHLY SIDEROPHILE ELEMENTS IN NON-CRATONIC PERIDOTITE XENOLITHS ..........................................................................................311 Nature of the host minerals........................................................................................311 Petrography of the base metal sulfides ......................................................................312 Origin of the base metal sulfides and platinum group minerals ................................315 HIGHLY SIDEROPHILE ELEMENTS AND 187Os/188Os RESULTS FROM NON-CRATONIC PERIDOTITE XENOLITHS ...........................................................316 Summary of results from whole-rock studies............................................................317 xviii

Highly Siderophile and Strongly Chalcophile Elements ‒ Table of Contents Summary of results from base metal sulfides and other mineral phases ...................321 Reconciling 187Os/188Os results from whole-rock and base metal sulfide analyses ...326 THE LIFE OF A XENOLITH, AS RECORDED IN HSE- AND Os-ISOTOPE SYSTEMATICS ..............................................................................................................327 The HSE and 187Os composition of the Primitive Upper Mantle ..............................328 Whole-rock observations on samples representing melting residues ........................330 Mineralogical control of HSE fractionation during partial melting ..........................332 Ancient melt-extraction events recorded by 187Os/188Os systematics.........................335 Post-melting petrological history ..............................................................................336 Syn- to post-eruptive processes .................................................................................350 CHRONOLOGICAL INTERPRETATION OF OS-ISOTOPIC DATA AND TECTONIC IMPLICATIONS .............................................................................................................351 Obtaining age information from Os isotopes of whole rocks....................................351 Obtaining age information from Os isotopes of base metal sulfides .........................354 Tectonic interpretation of Os model ages ..................................................................355 CONCLUSIONS...................................................................................................................356 ACKNOWLEDGEMENTS ..................................................................................................358 REFERENCES .....................................................................................................................358

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Re–Pt–Os Isotopic and Highly Siderophile Element Behavior in Oceanic and Continental Mantle Tectonites Harry Becker, Christopher W. Dale

INTRODUCTION ................................................................................................................369 BREVIA OF CONCEPTS, TERMINOLOGY, AND ANALYTICAL CAVEATS ..............370 Re–Pt–Os parameters ................................................................................................370 Normalization of concentration data .........................................................................370 Precision and accuracy of concentration data and analytical issues..........................371 HIGHLY SIDEROPHILE ELEMENTS IN MANTLE TECTONITES FROM DIFFERENT GEODYNAMIC SETTINGS ...................................................................371 Summary of mantle tectonites and their geodynamic settings ..................................371 HSE IN ABYSSAL PERIDOTITES FROM SPREADING OCEANIC LITHOSPHERE ..374 HSE in mantle tectonites from continental extensional domains and continent–ocean transitions ....................................................................................380 HSE in ophiolites that formed at fast spreading ridges with little or no influence from subduction processes .....................................................................................386 High-temperature orogenic peridotites from convergent plate margin settings ........388 Highly siderophile elements in peridotites and melt-reacted lithologies of ophiolites influenced by convergent plate margin magmatism ..............................392 Highly siderophile elements in the mantle sections of ophiolites of uncertain origin ......................................................................................................................398 DISCUSSION .......................................................................................................................399 Influence of low-temperature alteration processes on the HSE in bulk rocks and minerals............................................................................................................399 The influence of melt infiltration and partial melting on HSE abundances in mantle tectonites .....................................................................................................402 Summary—Mantle melting and mantle–magma interaction—different sides of the same coin ......................................................................................................416 Os isotopic heterogeneity in the mantle ....................................................................418 xix

Highly Siderophile and Strongly Chalcophile Elements ‒ Table of Contents The role of recycled oceanic lithosphere in producing HSE and Os isotope signatures in magmas .............................................................................................423 The relationship between abyssal peridotites and MORB: an osmium isotope perspective ..............................................................................................................426 Interpretation of Re–Os model ages ..........................................................................427 ACKNOWLEDGMENTS.....................................................................................................429 REFERENCES .....................................................................................................................430

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Chalcophile and Siderophile Elements in Mantle Rocks: Trace Elements Controlled By Trace Minerals Jean-Pierre Lorand, Ambre Luguet

INTRODUCTION ................................................................................................................441 BACKGROUND ...................................................................................................................442 Sulfides in the upper mantle and mantle rocks ..........................................................443 Abundance and phase control on chalcophile and siderophile elements in the fertile upper mantle.................................................................................................446 Partial melting of the mantle: a BMS-removing and PGM producing petrogenetic process. ..............................................................................................458 Chalcophile/siderophile element systematics in pyroxenites ....................................466 Low-pressure BMS dissolution in regional-scale open system melting of the sub-continental lithospheric mantle........................................................................469 BMS precipitation associated with magma percolation/metasomatism ....................472 Platinum-group minerals and magmatic percolation of the lithospheric mantle.......475 CONCLUDING REMARKS ................................................................................................480 ACKNOWLEDGMENTS.....................................................................................................481 REFERENCES .....................................................................................................................481

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Petrogenesis of the Platinum-Group Minerals Brian O’Driscoll, José María González-Jiménez

INTRODUCTION ................................................................................................................489 PHASE RELATIONS AND ORIGIN OF THE PGM ..........................................................490 Chemical properties of the PGM ...............................................................................490 Extraterrestrial occurrences of the PGM ...................................................................492 Origin of the terrestrial PGM: Mantle melting, metasomatism, and metal transfer ..495 PGM IN LAYERED MAFIC–ULTRAMAFIC INTRUSIONS ...........................................498 Chromitite-hosted layered intrusion PGM ................................................................498 Non-chromitite-hosted PGM in layered intrusions ...................................................507 PGM IN OPHIOLITES.........................................................................................................516 PGM in ophiolite peridotites .....................................................................................516 PGM in ophiolite chromitites ....................................................................................518 PGM in sulfide-rich ophiolite lithologies ..................................................................530 PGM IN PERIDOTITES OF THE SUBCONTINENTAL LITHOSPHERIC MANTLE ......................................................531 Subcontinental lithospheric mantle peridotite massifs ..............................................531 SCLM peridotite xenoliths ........................................................................................532 xx

Highly Siderophile and Strongly Chalcophile Elements ‒ Table of Contents PGM IN CONCENTRICALLY ZONED URALIAN–ALASKAN–ALDAN-TYPE COMPLEXES ..............................................533 PGM in dunite, pyroxenite and gabbro .....................................................................534 Chromitite-hosted PGM in CUAAC .........................................................................535 PGM and sulfide mineralization in CUAAC .............................................................539 PGM IN NI-SULFIDE DEPOSITS ......................................................................................542 Komatiite-associations...............................................................................................542 Magmatic Ni–(± Cu–± PGE)–sulfide deposits in non-komatiitic rocks ....................544 EXAMPLES OF UNCONVENTIONAL PGM OCCURRENCES .....................................549 Kimberlite- and Cu-porphyry-hosted PGM...............................................................549 OUTLOOK AND FUTURE WORK ....................................................................................551 Assessing the mineralogical and textural complexity of PGM assemblages ............551 Constraints on quantifying the distribution and grain size of PGM ..........................552 Advancing our understanding of the link between PGM assemblage and PGE geochemistry ..................................................................................................553 ACKNOWLEDGEMENTS ..................................................................................................554 REFERENCES .....................................................................................................................554 APPENDIX ...........................................................................................................................570 PGM in placers associated with ophiolite complexes ...............................................570 PGM mineralization in CUAAC placer deposits.......................................................573 REFERENCES .....................................................................................................................576

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Mantle Sulfides and their Role in Re–Os and Pb Isotope Geochronology Jason Harvey, Jessica M. Warren, Steven B. Shirey

INTRODUCTION ................................................................................................................579 BACKGROUND ...................................................................................................................581 ANALYTICAL METHODS AND PRACTICAL ASPECTS OF SAMPLE PREPARATION .......................................................................................585 BASE-METAL SULFIDE OCCURRENCE, MAJOR ELEMENT GEOCHEMISTRY AND PETROLOGY........................................................................................................589 Peridotite-hosted sulfides ..........................................................................................592 Pyroxenite-hosted sulfides .........................................................................................603 Diamond-hosted sulfides ...........................................................................................606 Re–Os–Pb MASS BALANCE IN ULTRAMAFIC SAMPLES ...........................................610 Osmium mass balance ...............................................................................................610 Rhenium mass balance. .............................................................................................613 Lead mass balance .....................................................................................................614 GEOCHRONOLOGICAL METHODS, MODEL AGES, AND POTENTIAL PITFALLS. .....................................................................................615 Sulfide Re–Os isochrons, TMA, TRD, and gOs..............................................................615 Potential pitfalls with sulfide geochronology. ...........................................................620 THE UTILITY OF Re–Os AND Pb ISOTOPE GEOCHRONOLOGY ...............................626 Dating the formation of diamonds.............................................................................626 Diamond formation through time ..............................................................................628 The age of the continental lithospheric mantle and the assembly of its domains .....628 The relationship between the age of the SCLM and the overlying crust ..................630 xxi

Highly Siderophile and Strongly Chalcophile Elements ‒ Table of Contents The inherent heterogeneity within the oceanic mantle ..............................................632 CONCLUDING REMARKS AND FUTURE DIRECTIONS .............................................635 ACKNOWLEDGMENTS.....................................................................................................635 REFERENCES .....................................................................................................................635

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Highly Siderophile Element and Os Isotope Systematics of Volcanic Rocks at Divergent and Convergent Plate Boundaries and in Intraplate Settings Abdelmouhcine Gannoun, Kevin W. Burton, James M.D. Day, Jason Harvey, Pierre Schiano, Ian Parkinson

INTRODUCTION ................................................................................................................652 HIGHLY SIDEROPHILE ELEMENT DISTRIBUTION AND BEHAVIOR IN THE UPPER MANTLE .............................................................................................654 Core formation and the late accretion of impactor material ......................................654 Highly siderophile elements in mantle minerals .......................................................655 Highly siderophile element behavior accompanying fractional crystallization ........660 THE 187Re–187Os ISOTOPE SYSTEM AND THE FORMATION OF MID-OCEAN RIDGE BASALT (MORB) .....................................................................661 Introduction ...............................................................................................................661 Analytical techniques ................................................................................................662 Rhenium–Osmium elemental variations in MORB glass .........................................663 The 187Os/188Os isotope variations in MORB glass ....................................................665 Analytical issues associated with MORB..................................................................668 SULFIDES IN MID-OCEAN RIDGE BASALTS ...............................................................677 Petrology and chemistry ............................................................................................677 187 Re–187Os behavior in MORB sulfide ......................................................................680 The 187Os/188Os composition of the MORB mantle source ........................................682 LOWER OCEANIC CRUST ................................................................................................683 Assimilation of gabbroic lower crust.........................................................................686 HSE ABUNDANCES AND Re–Os ISOTOPE SYSTEMATICS OF INTRAPLATE VOLCANISM ..................................................................................686 Mantle melting processes ..........................................................................................687 Osmium isotopes as tracers of hotspot sources .........................................................688 Crustal and lithospheric mantle assimilation/contamination.....................................691 The origin of Continental Flood Basalts (CFB) and Large Igneous Provinces (LIP) ...............................................................................693 Continental intraplate alkaline volcanism .................................................................696 Processes affecting the HSE compositions of sub-aerial volcanism .........................699 HIGHLY SIDEROPHILE ELEMENT SYSTEMATICS OF ARCS ....................................700 HSE and 187Os/188Os in arc lavas ................................................................................700 HSE and 187Os/188Os in arc xenoliths .........................................................................704 Radiogenic Os from slab components or from crustal contamination ......................705 Mechanical mixing processes ....................................................................................706 CONCLUSIONS AND PERSPECTIVES ............................................................................707 ACKNOWLEDGMENTS.....................................................................................................708 REFERENCES .....................................................................................................................708

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Highly Siderophile and Strongly Chalcophile Elements in Magmatic Ore Deposits Sarah-Jane Barnes, Edward M. Ripley

INTRODUCTION ................................................................................................................725 CLASSIFICATION OF THE DEPOSITS ............................................................................731 Reef or stratiform deposits ........................................................................................731 Contact deposits ........................................................................................................731 Ni-sulfide deposits .....................................................................................................732 MINERALS HOSTING THE PLATINUM-GROUP ELEMENTS .....................................734 Base metal sulfides ....................................................................................................734 Platinum-group minerals ...........................................................................................737 Chromite ....................................................................................................................738 Mass Balance .............................................................................................................738 GEOCHEMISTRY ...............................................................................................................739 Introduction ...............................................................................................................739 Normalization to mantle or chondrite? ......................................................................740 Recalculation to 100% Sulfides or whole rock..........................................................740 Other chalcophile elements .......................................................................................742 INTERPRETATION .............................................................................................................743 Composition of the silicate melt ................................................................................743 Saturation of the magma in a sulfide liquid ...............................................................746 Upgrading of the Sulfides ..........................................................................................747 Crystallization of a sulfide liquid ..............................................................................752 Late magmatic fluids .................................................................................................754 Subsolidus events ......................................................................................................755 UTILIZATION OF THE Re–Os ISOTOPE SYSTEM IN STUDIES OF MAGMATIC Ni–Cu–PGE ORE GENESIS ..........................................................................................756 Background................................................................................................................756 The “R-factor” and its application to Re–Os isotopes ...............................................759 Examples of the application of the Re–Os isotope system to magmatic ore deposits ........................................................................................759 CONCLUSIONS...................................................................................................................765 ACKNOWLEDGMENTS.....................................................................................................766 REFERENCES .....................................................................................................................766

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Reviews in Mineralogy & Geochemistry Vol. 81 pp. 1-87, 2016 Copyright © Mineralogical Society of America

Experimental Results on Fractionation of the Highly Siderophile Elements (HSE) at Variable Pressures and Temperatures during Planetary and Magmatic Differentiation James M. Brenan1,2 1

Department of Earth Sciences Dalhousie University Halifax Nova Scotia B3H 4R2 Canada [email protected]

Neil R. Bennett Geophysical Laboratory 5251 Broad Branch Rd. NW Washington, DC 20015−1305 USA [email protected]

Zoltan Zajacz 2

Department of Earth Sciences Earth Sciences Centre 22 Russell Street Toronto Ontario M5S 3B1 Canada [email protected]

INTRODUCTION AND SCOPE The platinum-group elements (PGEs; Os, Ir, Ru, Rh, Pt, Pd), along with rhenium and gold, are grouped together as the highly siderophile elements (HSEs), defined by their extreme partitioning into the metallic, relative to the oxide phase (> 104). The HSEs are highly refractory, as gauged by their high melting and condensation temperatures, and were therefore relatively concentrated in the feedstock for the terrestrial planets, as defined by the composition of chondritic meteorites (e.g., Anders and Ebihara 1982; Horan et al. 2003; Fischer-Gödde et al. 2010). However, the planetary formation and differentiation process has since acted on this chemical group to produce a rich variety of absolute and relative inter-element fractionations. For example, analysis of iron meteorites suggests a significant decoupling of the HSE in the cores of planetesimals, and likely Earth’s core, with Os, Ir, Ru (IPGE-group) and Re concentrated in the metal phase, and Pt, Rh, Pd (PPGE-group) plus Au usually concentrated in the residual liquid (Goldstein et al. 2009). In terms of the silicate Earth, analysis of mantle rocks reveals very low levels of the HSE, but relative abundances similar to chondrites (see review by Day et al. 1529-6466/16/0081-0001$10.00

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Brenan, Bennett & Zajacz

2016, this volume), in part reflecting HSE segregation into core-forming iron (Ringwood 1966; Ganapathy et al. 1970). This is in contrast to mantle-derived melts, whose HSE abundances are highly fractionated, with relative depletions in the IPGE-group compared to PPGE-group, as well as Re and Au (Barnes et al. 1985). Resulting Re/Os and Pt/Os fractionation also influence the long-term evolution of the 187Re to 187Os and 190Pt to 186Os decay systems, and, hence, the development of distinctive Os-isotope reservoirs (Walker et al. 1997; Shirey and Walker 1998; Day 2013). The emplacement of mantle-derived magmas into Earth’s crust results in a further decoupling of the HSE suite. Crystallization of mafic and ultramafic magmas appears to leave the IPGE in magmatic cumulates, while concentrating the PPGEs, Re and Au in the more differentiated products (e.g., Brugmann et al. 1987; Puchtel and Humayan 2001, 2005). The onset of sulfide liquid saturation in these systems can produce a wholesale reduction of HSE concentrations in the silicate melt, reflecting the highly chalcophile nature of this element suite. Magmatic sulfide liquids may differentiate internally, with further concentration of the IPGEs and Re into accumulations of the early formed monosulfide solid solution (MSS), and subsequent enrichment in the liquid residue in the PPGEs and Au (Barnes and Ripley 2016, this volume). As the silicate magma differentiates further, an orthomagmatic fluid may develop, possibly disturbing the primary distribution of the HSE within the pile of accumulated solids (e.g., Boudreau et al. 1986; Boudreau and Meurer 1999). More evolved magmas emplaced at a high level in the crust, or even erupted, may have some of their remaining HSE collected into a low-density vapor phase, which may be deposited in a hydrothermal stockwork (Richards 2011), or dispersed into the atmosphere during volcanic eruptions (e.g., Naughton et al. 1976; Zoller et al. 1983; Toutain and Meyer 1989; Crocket 2000; Yudovskay et al. 2008). Hence, the HSE are not only fractionated during planetary differentiation, but during this process, these elements may exhibit four out of the five geochemical classifications originally proposed by Goldschmidt (1958; Table VI, page 25); concentration in the core, and magmatic sulfides (siderophile and chalcophile), partitioning into silicates/oxides (lithophile), and expulsion in volcanic emanations (atmophile). In order to make full use of this unique suite of elements to understand the planetary differentiation process, information on their partitioning amongst solid/liquid iron metal, sulfide, silicate, oxide phases, and vapour/fluid, as well as the stability of HSE-bearing accessory phases is required. The advent of procedures to detect very low concentrations of the HSE in these various phases has greatly expanded our empirical understanding of this behavior (Meisel and Horan 2016, this volume), but significant uncertainties still remain. Laboratory experiments offer a complimentary approach to the empirical studies, providing constraints on the nature of HSE fractionation involving specific phases, and variation with intensive parameters. This chapter provides a review of that work, with an emphasis on results pertaining to processes occurring mostly at the magmatic stage. For each of the experimental systems considered, we have provided some information on how the experiments are done, the methods of analysis and attempt to place the results in a theoretical framework.

SOLID METAL–LIQUID METAL PARTITIONING Studies of the iron meteorites have shown large variations in both the relative and absolute abundances of the HSE (e.g., McCoy et al. 2011; see review by Day et al. 2016, this volume). In part, this variation may derive from differences in the bulk HSE composition of the meteorite parent bodies, but significant differences exist within groups derived from a single parent body, reflecting the role of internal differentiation processes (e.g., Scott 1972; Scott and Wasson 1976; Goldstein et al. 2009). So-called non-magmatic iron meteorites (Types IAB and IIICD) are thought to derive by impact melting of planetesimals, with the variation in HSE concentrations due to mixing of different melt fractions (e.g., Choi et al. 1995). In contrast, HSE variation in the magmatic iron group is consistent with crystal–liquid fractionation during solidification of the parent body core; as mentioned, segregating metal concentrating the IPGE

Experimental Fractionation of the HSE

3

and Re, with enrichments in the PPGE + Au in residual melt (e.g., Scott 1972; Pernicka and Wasson 1987; Walker et al. 2008; Goldstein et al. 2009). Studies of the magmatic iron group have also emphasized the possible role of non-metal components, such as S, C, Si, and P, in affecting the solid metal–liquid metal partitioning, as accessory phases containing those elements are ubiquitous (Goldstein et al. 2009). This is also consistent with the need for a light-element component in the cores of the terrestrial planets, in order to explain their density deficit, and to satisfy cosmochemical constraints (Dreibus and Wanke 1985; McSween 1994; McDonough 2003). Laboratory experiments of solid metal–liquid metal partitioning have provided the means to verify the core crystallization model for the magmatic irons, and also inform about the effects of non-metal components (Willis and Goldstein 1982; Chabot and Drake 1999; Goldstein et al. 2009) as well as the influence of the crystalline metal (Van Orman et al. 2008; Stewart et al. 2009; Rai et al. 2013) on HSE distribution during solidification. Also, a major focus of study in recent years has been the relative fractionation of Re/Os and Pt/Os arising from inner core solidification, which bears on the development of deep planetary reservoirs with distinct 187Os/188Os and 186Os/188Os isotopic compositions (e.g., Walker et al. 1995, 1997; Brandon et al. 1998). The following sections outline the parameterizations used to describe variation in the solid-metal/liquid metal partition coefficient, DSM/LM, as a function of either liquid metal composition or solid metal structure. Several considerations for experiments at both ambient and high pressure are first described.

EXPERIMENTAL APPROACH TO SOLID METAL–LIQUID METAL PARTITIONING (DSM/LM) Past studies have focused on the separate roles of the liquid and solid phases on controlling the absolute and relative magnitudes of DSM/LM. The structure of solid Fe is expected to undergo changes due to both pressure and the incorporation of light elements. The consequences of increasing pressure are manifold, including control on the stable atomic arrangement (e.g., the FCC to HCP transition at high T; Komabayashi et al. 2009), unit cell volume, and effect on the solubility of light elements in the Fe-metal structure (e.g., Zhang and Fei 2008). Although temperature is likely to play a second-order role, most experimental data have been acquired at conditions significantly cooler than those expected to accompany core crystallization in larger planetary bodies. This makes T worthy of investigation in future work. Outlined here are several different experimental approaches that have been used to determine DSM/LM over a range of P–T–X conditions. Most of the early experimental studies of HSE partitioning in solid metal–liquid metal systems were conducted at 0.1 MPa and temperatures corresponding to the Fe (+ Ni) − X liquidus, with X being C, S, O, Si, and P, with more recent work focusing on the effect of pressure. The primary concerns for experiments in the Fe-rich systems of interest here are the attainment of equilibrium and the potential for reaction between the container material and the charge. The compositional space that can be accessed is governed by phase relations in the chosen system and must also be considered in the experimental design. For experiments done at ambient pressure, there is the additional consideration of volatile element loss from the sample. With these in mind, the experimental protocol adopted for experiments to determine siderophile element partitioning at ambient pressure in the Fe ± S, P, C systems has remained essentially unchanged since the pioneering experiments of Drake et al. (1978). The light element components are typically introduced to starting materials as Fe-sulfide, elemental phosphorus and graphite powder, for S, P, and C, respectively, which are mixed and ground with metallic Fe, Ni, and typically one HSE dopant at the wt% level (e.g., Jones and Drake 1983; Malvin et al. 1986; Chabot et al. 2006), although multiple HSE can be added at ppm levels. As shown by Fleet et al. (1999), constant partitioning is obtained over the ppm to wt% concentration range. The most commonly

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a

b

100 μm

SiO2 Ampoule

SiO2 Spacer Laser Ablation Pits Sample

S-rich

Solid Fe

S-poor

Al2O3 Crucible 0.1 MPa; 1350 oC 50 μm

Figure 1. (a) Experimental arrangement used to perform equilibrium partitioning experiments in solid metal–liquid metal systems at 0.1 MPa. (b) Back scattered electron images of the run-product texture from an experiment at 0.1 MPa and 1350 ºC. When quenched, the liquid portion of the sample unmixes to produce sulfur-rich (darker) and sulfur-poor (brighter) domains in the sample.

employed crucible material is vitreous alumina, as it is unreactive with the alloy phase and stable to the required temperatures (e.g., Drake et al. 1978; Jones and Drake 1983). Owing to the volatile nature of the light element additives, most past experiments at low pressure have been done in evacuated and sealed silica glass ampoules (Fig. 1a), although some early results (free of S and P) were obtained with open crucibles in contact with an Ar–H2 atmosphere (Willis and Goldstein 1982; Jones and Drake 1983). In order to aid sample homogenization and reduce the time required for equilibrium, some studies have employed a superliquidus step lasting for several hours, then cooling to the intended equilibration temperature of the experiment (Drake et al. 1978; Chabot et al. 2006, 2007). Typical experiment durations are between 12 and 200 h (e.g., Chabot et al. 2006, 2007), although runs as short as 5 h appear to be sufficient for equilibrium in the Fe–Ni–S–P system at 0.1 MPa and 1250 ºC (Malvin et al. 1986). Experiments are terminated by immersing the sample in cold water. Although the rate of heat loss is high, quenching an iron-rich liquid containing S, P or C by this method still results in precipitation of a heterogeneous intergrowth of Fe-rich metal with a light element-rich component (Fig. 1b). In order to obtain an average composition of the quenched melt phase, run product analysis by electron microprobe, or more recently by LA-ICPMS, employs a defocused or rastered beam. Certain experimental systems offer unique challenges. Phase relations in the Fe–Si system show that for compositions more Fe-rich than the eutectic that exists between Fe and Fe2Si, the solid and liquid alloy phases contain similar quantities of Si (Kubaschewski 1982). Isolating the effect of liquid composition on partitioning is therefore not straightforward. To circumvent this problem, Chabot et al. (2010) conducted experiments in the Fe–S–Si system and monitored trace element distribution between immiscible melt pairs; one rich in the S and the other in Si. By comparison to the well-established effects of S on partitioning, the role of Si could be evaluated independently. The effect of oxygen on DSM/LM has also been difficult to evaluate, as the solubility of oxygen in molten Fe is diminishingly small at low pressure. However, work by Naldrett (1969), and later by Fonseca et al. (2008), has shown that the

Experimental Fractionation of the HSE

5

oxygen content of molten Fe–S can reach several wt%. Taking advantage of this effect, Chabot et al. (2015) measured partitioning in the Fe–S–O system, and again, by comparison to partitioning systematics in the Fe–S system, were able to evaluate the role of O. Experiments done at elevated pressures have utilized both piston-cylinder and multi-anvil apparatus. Crushable MgO is often employed as a sample container, as it is unreactive with Fe-rich alloys over a broad spectrum of conditions, and promotes a homogenous pressure distribution (e.g., Jones and Walker 1991; Walker 2000; Lazar et al. 2004; Van Orman et al. 2008). In order to prevent leakage of the liquid alloy phase during the experiment, the crushable MgO is annealed at ~ 600−800 ºC for ~ 8−12 hrs at the target pressure (e.g., Jones and Walker 1991; Walker 2000). Hard-fired alumina crucibles have also been employed successfully to investigate partitioning in the Fe–S and Fe–C systems (Chabot et al. 2008, 2011; Rai et al. 2013), as well as graphite to study HSE partitioning in the Fe–S–O system (Fleet et al. 1991). At the low temperatures and high-sulfur content of the experiments done by Fleet et al. (1991), the solubility of carbon in the sulfur-rich liquid phase is negligible. However, partial closure of the Fe–C and Fe–S liquids miscibility gap with increasing pressure (Corgne et al. 2008), may result in significant levels of carbon in the resulting liquid, and therefore difficulty in isolating the separate effects of C and S. Graphite-saturated experiments in the Fe–S–C system may also become saturated in cementite (Fe3C), rather than Fe, as the solid phase coexisting with sulfur-rich liquids (Dasgupta et al. 2009; Buono et al. 2013). Refractory Os–Ir ± Fe alloy grains, in some cases rimmed by Pt, were observed by Fleet et al. (1991) in experiments conducted in the Fe–S-O system at 1100−1200 ºC and 4.5−11 GPa. Some experiments done at ambient pressure and temperatures of 1000−1200 ºC in the Fe–Ni–S system also display evidence for heterogeneous Os concentrations in the solid metal phase (Fleet et al. 1999). A similar problem with Ir homogeneity was noted by Jones and Drake (1983). These grains result from the use of HSE-rich starting materials that are resistant to dissolution at the chosen experimental temperatures and durations. A factor which might have exacerbated the equilibration problem in the Fleet et al. experiments was the loading of several of the PGEs into each experiment, producing significant domains of Fe-poor, IPGE-rich alloy in which diffusion of all the PGEs is intrinsically slow (Watson and Watson 2003). To prevent the formation of refractory HSE grains in experiments done at 10 GPa and 1400−1500 ºC, Walker (2000) used a novel experimental arrangement to prevent alloying between the HSEs and Fe early in the experiment. This was achieved by taking advantage of the large thermal gradient intrinsic to the multianvil assembly, and positioning powdered Pt, Os, and Re metal in what would become the hot zone of the capsule, causing complete dissolution and precipitation with newly formed Fe–Ni crystals at the cold end. Experiments that exploit a thermal gradient have also been used to determine solid metal–liquid metal partitioning in the Fe–Si system (Morard et al. 2014). As the melting loop in the Fe–Si system is only a few tens of degrees, Morard et al. (2014) imposed a vertical temperature gradient across samples so as to increase the relative proportions of metal and liquid for a given bulk composition. Effect of liquid metal composition on DSM/LM. The effect on DSM/LM resulting from changes to the liquid metal composition has been parameterized by Jones and Malvin (1990) in terms of a non-metal interaction model. Jones and Malvin (1990) consider partitioning of an element (i) between solid (SM) and liquid (LM) metal. At equilibrium, the chemical potential of i is equal in both phases: (1)

µSM = µiLM i

Expressing these chemical potentials in terms of the standard state chemical potential and the activity of element i in each phase (ai) yields: O

O

µSM + RT ln aiSM = µiLM + RT ln aiLM i

(2)

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6

Recalling ai = γ i Xi (where γ i and Xi are the activity coefficient and mole fraction of i respectively), the activity terms may be replaced to provide an expanded form of Equation (2): O

O

µSM + RT  ln γ SM + ln XiSM  = µiLiq + RT  ln γ iLM + ln XiLM  i i

(3)

Equation (3) may be rearranged to yield: O

O

LM SM  SM  µi − µi  SM  (4) − ln  γ i LM  ln  Xi LM  = Xi  γi  RT   Substituting the molar partition coefficient ( Di* SM / LM ) into the left hand side of Equation (4) O O and the Gibbs free energy of reaction ( ∆GrO ) for µiLM − µiSM yields:

*SM/LM ln Di=

−∆GrO  SM  − ln  γ i LM  γ RT i  

(5)

From Equation (5) we see that for a system at constant pressure, the molar partition coefficient will depend upon temperature and changes to the activity coefficient of element i in both the solid- and liquid-metal phase. Jones and Malvin (1990) argue that both T and changes to γ SM i have a negligible effect on Di*SM/LM relative to changes in γ iLM. The effect of T can be assessed by considering an ideal system, in which the ratio γ SM / γ iLM i *SM/LM O is unity, and variation in Di depends only on the free energy term. ∆Gr in Equation (5) corresponds to the melting reaction for pure solid and liquid phases of element i, and can be written in terms of the heat of fusion ( ∆H m ) for the element of interest. Remembering that −∆GrO = − ( ∆HrO − T ∆SrO ) and replacing the subscript r with m, to denote the melting reaction, Equation (5) becomes: ∆H m ∆Sm (6) − + ln Di*SM/LM = RT R At the melting temperature of i, −∆GrO = 0 and so ∆Sm = ∆H m / Tm . Substituting this identity into Equation (6) yields the following expression:

Di*SM/LM ln=

∆H m R

 1 1  −   Tm T 

(7)

Equation (7) can be used to assess the variation in Di*SM/LM arising from changes in temperature. Table 1 lists the melting temperatures and enthalpies of fusion for the highly siderophile elements. Selecting osmium, which has the largest enthalpy of fusion, we calculate that Di*SM/LM will only vary by a factor of ~ 2.5 if T changes by 1000 ºC . The weak dependence of Di*SM/LM on temperature suggested by this analysis is borne out by the similar correlation coefficients (R2) obtained from simple regression of Di*SM/LM vs. liquid-metal composition and multiple linear regression vs. both liquid-metal composition and 1/T (Jones and Malvin 1990). For systems at relatively low P, in which the solubility of light elements in solid Fe–Ni alloy is small, Jones and Malvin (1990) argue that changes in γ SM are negligible relative i to changes in γ iLM. They support this assertion by noting that values of Di* SM / LM in the light- element free system are typically close to unity, the solid Fe–Ni system is itself relatively ideal, and that their 0.1 MPa experimental data can be adequately modeled considering only the γ iLM term. Chabot and Jones (2003) later used a modified version of the Jones and Malvin (1990) parameterization to successfully model a much larger database of 0.1 MPa partitioning data, providing further support for an approach that considers only changes in γ iLM.

Experimental Fractionation of the HSE

7

Table 1. Heats of fusion of the highly siderophile elements. Element

Tm (K)

Hm (kJ/mol)

Re

3458

34.08

Os

3306

57.85

Ir

2739

41.12

Ru

2606

38.59

Pt

2041

22.175

Rh

2236

26.59

Pd

1828

16.74

Au

1337

12.55

From data summarized in Haynes (2014).

Accepting that changes in T and γ SM do not strongly drive changes to the solid metal– i liquid metal partition coefficient, we can return to the relationship given in Equation (5) and simplify it to yield  1  ln DiSM/LM = ln γ iLM + constant − ln  LM  + constant = γ  i 

(8)

For geo- and cosmochemical purposes, the more useful independent variable is liquid-metal composition rather than γ iLM . These are linked, however, through interaction parameters which describe the excess free energy of mixing. For a liquid-iron alloy containing solutes i through N in dilute concentration, the activity coefficient of i is often described using only the first-order interaction parameters (e): N

ln γ iLM = ln γ i0 + ∑εij X j

(9)

j =2

where γ i0 is the activity coefficient of i at infinite dilution in liquid iron and Xj is the mole fraction of the subscript component in solution. It should be noted that by ignoring interaction parameters greater than first order, Equation (9) is not thermodynamically rigorous and should only be applied to alloys in which the solutes are dilute (Ma 2001 and references therein). Despite this, the following model adequately describes the experimentally determined partitioning of several HSEs in systems containing up to ~ 30 wt% sulfur (Jones and Malvin 1990). Although Ni is also typically present in moderate concentrations, it’s relatively weak interactions with Fe, S, P, and C permit the simplification embodied by Equation (9) while still allowing most experimental data to be modeled successfully (Jones and Malvin 1990). As an example, if we consider the case of iridium partitioning in an Fe–Ni–S alloy, the activity coefficient of Ir, as given by Equation (9) is: 0 Ni S Ir ln γ LM Ir = ln γ Ir + ε Ir X Ni + ε Ir X S + ε Ir X Ir

(10)

From Equations (8) and (10) it is apparent that the solid metal–liquid metal partition coefficient of Ir can be expressed in terms of the component mole fractions weighted by their corresponding interaction parameters. Jones and Malvin (1990) propose a further simplification, whereby the alloy is described as comprising only metal and non-metal domains, which either accept or reject the HSEs respectively. In this framework, changes to the activity coefficient as a function of liquid composition can be described by a single modified interaction parameter, termed the b factor. For Ir in the Fe–Ni–S system:

Brenan, Bennett & Zajacz

8

ln γ LM Ir = β Ir ln (1 − αnX S )

(11)

where a is a constant specific to the compositional system being investigated and n is a stoichiometric coefficient related to the speciation of the non-metal component in the alloy. Equation (11) retains the linear dependence of the activity coefficient on composition inherent in Equation (10) and is subject to the same limitations discussed above. Substituting Equation (11) into (8), yields DIr*SM/LM as a function of the non-metal content of the alloy and the b factor:

ln DIr*SM/LM = βIr ln (1 − αnXS ) + C

(12)

where C is a constant. As the sulfur (or other light-element) content of the alloy tends towards zero, ln DIr*SM/LM will tend towards C, such that in the limiting case of Xs = 0, C will equal the partition coefficient in the light-element-free system ( ln Di0 ) . Figure 2 provides an example of partitioning data for Au and Re in the Fe–Ni–S system plotted in the form of Equation (12), illustrating the overall linearity of the data. For partitioning of an element i in a system containing light elements j through N, the general form of Equation (12), therefore, becomes: N   ln Di*SM/LM = ln Di0 + βi ln  1 − ∑α j n j X j  j  

(13)

The value of bi in Equation (13) for a system containing multiple light elements is related to the b factors in each of the individual light-element-bearing systems. Jones and Malvin (1990) express this relationship as a weighted average of the effects in the end-member systems:  nX βij…. N = βij  N j  nX  ∑j j

   + β k  nX k  i  NnX   ∑j j

  … + β N  nX N i   NnX   ∑j j

   

(14)

For example, the b factor for element i in the Fe–Ni–S–P system, where nS and nP are 2 and 4, respectively, is described as:

2.5

a) Au

2.0

ln DSM/LM

b) Re

8.0

1.5 1.0

10.0

6.0

slope = β

0.5 4.0

0.0 -0.5

2.0

-1.0 -1.5 -3.5

-3.0

-2.5

-2.0

-1.5

ln(1-2αXs)

-1.0

Ambient Pressure 0.7 GPa 2.7 GPa Jones & Walker 1991 7.5 GPa 8 GPa

-0.5

0.0

0.0 -3.5

-3.0

-2.5

-2.0

-1.5

ln(1-2αXs)

Ambient Pressure 10 GPa - Walker 2000 3.3 GPa Van Orman et al. 2008 18 - 22 GPa

-1.0

-0.5

0.0

5 GPa 10 GPa Hayashi et al. 15 GPa 2009 20 GPa

Figure 2. The variation in ln DSM/LM as a function of liquid metal composition, as described by the parameter ln(1 − 2aXs), for Au (a) and Re (b). Despite differences in P and T between the experiments, each element is adequately described by a single linear fit. The partition coefficient for Re, however, is significantly more sensitive to liquid composition, owing to the larger value for bRe.

Experimental Fractionation of the HSE

9

  P  4 XP  2 XS βSP βSi  i =  + βi    2 XS + 4 X P   2 XS + 4 X P 

(15)

Chabot and Jones (2003) develop further the parameterization outlined here, such that only a single beta factor need be determined in systems containing multiple light elements. In order to implement predictive models of partitioning in light-element-bearing systems, the key parameters to determine from experiments are therefore DSM/LM and the b factors. Table 2 provides a summary of these values determined for the HSE at 0.1 MPa as defined using the Jones–Malvin formalism. Values of DSM/LM and b show a ~ 5-fold difference in magnitude amongst the HSE, with the strongest melt composition effects implied for Re–Os–Ir, and the least for Pd and Au. These differences serve to further decouple the HSE during core solidification as the light-element component builds up in the residual melt. Significantly, although values of b for Re, Os, and Pt are similar, there are resolvable differences in the values of DSM/LM, in the order Os > Re > Pt. This result was noted based on empirical estimates from magmatic iron meteorites (Walker et al. 1995), and used to develop the hypothesis that mafic magmas with anomalous enrichments in 187Os/188Os and 186Os/188Os contain an outer core component (Walker et al. 1995, 1997; Brandon et al. 1998). Subsequently, experiments have been done to document the effect of pressure on the relative partitioning of these elements (3−22 GPa; Walker 2000; Van Orman et al. 2008; Hayashi et al. 2009). Some of these studies suggest that DSM/LM becomes smaller, and more similar, likely due to an increased size misfit in Fe metal (Van Orman et al. 2008), as described below. The role of the solid phase on DSM/LM. As described in the previous section, much of the available solid metal–liquid metal partitioning data can be adequately described using a parameterization which takes into account metal-solvent interactions in the liquid (e.g., Jones and Malvin 1990; Chabot and Jones 2003). This approach however, provides no theoretical explanation for the role of the crystalline Fe–Ni solid phase, as manifested by ~ 5-fold difference in values of DSM/LM between the HSE (Table 2). Several recent studies have sought to provide this theoretical framework through application of a modified form of the lattice strain model (Van Orman et al. 2008; Stewart et al. 2009; Chabot et al. 2011; Rai et al. 2013). This model, as commonly applied to silicate systems, quantifies the parabolic relationship between log DCrystal/Melt and ionic radius (Blundy and Wood 1994). This functional form arises because the variation in partition coefficients for a suite of isovalent cations originates purely in the elastic strain incurred Table 2. Summary of parameters to predict solid metal–liquid metal partitioning. Element

ln D01

b (Soret)2

Error

n1

b  (Partitioning)1

Error

n3

Ru

0.12

1.72

0.31

4

2.10

0.15

15

Rh

1.31

0.27

4

1.94

0.08

3

Pd

0.52

0.13

4

0.54

0.15

6

2.34

0.40

4

2.72

0.05

37

Re

0.47

Os

0.63

Ir

2.36

0.40

4

2.79

0.07

49

2.17

0.38

4

2.83

0.13

36

Pt

–0.18

1.80

0.33

4

2.45

0.09

34

Au

–1.24

0.71

0.19

4

1.06

0.07

23

Notes: 1 Values derived from regression of the ambient pressure data summarized by Chabot and Jones (2003) and the high-pressure results of Jones and Walker (1991), Walker (2000), Lazar et al. (2004), Van Orman et al. (2008), and Hayashi et al. (2009). 2 Beta determined by the Soret experiments of Brenan and Bennett (2014). 3 Number of experiments.

10

Brenan, Bennett & Zajacz

by the size mismatch between the substituent cation and the optimal radius for that site. An analogous approach has been taken for application to metallic systems, with DiSM/LM cast as:  1 ln = DiSM/LM ln D0SM/LM +  −4 π N A EM  r0 ( M ) ri −r0 ( M ) 2  

(

)

2

+

3  1 ri −r0 ( M )   / RT 3 

(

)

(16)

where D0SM/LM is the partition coefficient for an element with the ‘ideal’ neutral atomic radius for site M (r0(M)), NA is Avogadro’s constant, EM is the apparent Young’s modulus for site M and ri is the radius of a neutral atom of element i. Some previous studies have used the neutral atomic radii of Clementi et al. (1967) for values of ri and fit the experimental data to Equation (16) by varying D0, r0, and EM (Stewart et al. 2009; Chabot et al. 2011; Rai et al. 2013). Application of this model by Stewart et al. (2009) to previous results for HSE partitioning in the Fe–S and Fe–C systems, at 0.1 MPa to 22 GPa (Chabot and Jones 2003; Chabot et al. 2006; Van Orman et al. 2008), suggests incorporation of these elements in the lattice does not occur through simple replacement of Fe. The large value of r0 (1.83 Å) for the parabola defined by the 3rd row transition elements, Au through W, is significantly greater than the atomic radius of Fe (1.56 Å) and may instead suggest the accommodation of these elements in defects resulting from the presence of light elements in the Fe lattice (Stewart et al. 2009). However, Van Orman et al. (2008) showed that the partitioning of Re, Os, and Pt was consistent with their relative increase in metallic radius compared to that of Fe in the FCC structure. Chabot et al. (2011) considered a larger suite of data in this context, and found that the systematic partitioning trends using atomic radii as the ordinate broke down for the 1st- and 2nd-row transition elements when plotted as a function of the metallic radius. The reason for these differences are unclear, but certainly bear on our interpretation of which sites the HSE partition in the solid-metal phase, and the fundamental controls on inter-element fractionation. As a final point to this section, we note that the ~ 5-fold variation in b is similar to the differences in DSM/LM for the HSE, with the latter related to the elastic strain generated by size mismatch from the optimal site size in FCC iron. Values of b provide a measure of the affinity of a particular HSE for Fe-rich domains within the liquid structure, with larger values of b signifying enhanced sulfur avoidance (Jones and Malvin 1990). Thus, it seems reasonable that the extent to which a particular HSE will concentrate in more Fe-rich, non-metal-poor melts should be related to the size of the metal atom, provided Fe-solid and Fe-rich melt have similar structures. The structural similarity between liquid and solid metal is implied by the small ΔV for the solid-to-liquid metal transfer of components, as indicated by the relatively small effect of pressure on DSM/LM as documented in previous work (e.g., Van Orman et al. 2008; Chabot et al. 2011). As shown by Brenan and Bennett (2010), values of DSM/LM show a somewhat stronger dependence on metal, or atomic radius than b, but there is a sympathetic variation between the two, suggesting similar origins of the HSE “selectivity” for Fe-rich domains. As proposed by Van Orman et al. (2008), with increasing pressure the difference in absolute and relative values of DSM/LM for Pt, Re, and Os (and likely the other HSE) decrease due to the increased size mismatch between the substituent metal and Fe, and between individual HSE. If so, then we expect that not only will pressure decrease the absolute and relative values of DSM/LM (as per Van Orman et al. 2008), but, by analogy, pressure may also reduce the differences between individual values of b. The effects of increased pressure are not straightforward, however, as exemplified by the results of high-pressure experiments for Pt, Re, and Os done in the Fe–Ni–S system (vs. the Fe–S system studied by Van Orman et al.), which instead show a slight increase in DSM/LM with increasing P (Hayashi et al. 2009). Soret Diffusion Experiments. The effect of liquid composition on solid metal–liquid metal partitioning can also be determined from experiments that impose a thermal gradient on nonideal Fe-alloy solutions (e.g., Fe–S; Brenan and Bennett 2010). These experiments produce run-

Experimental Fractionation of the HSE

11

products with major-element compositional gradients that reflect the opposing mass fluxes of Soret and chemical diffusion. Soret diffusion arises when the system contains components that possess a different partial molar enthalpy when undergoing activated transport in the medium. A component with higher enthalpy in transport will migrate from hot to cold; transporting heat to the cold portion of the system and acting to ameliorate the imposed temperature imbalance. The opposite sense of migration is expected for components with a lower enthalpy in transit, thus also redistributing heat so as to reduce the thermal gradient. Details of the Soret process and its application to complex geologic systems are treated in detail by Lesher and Walker (1986, 1991). Chemical diffusion can limit the magnitude of segregation by Soret diffusion, due to the chemical potential gradients that arise from compositional differences along the sample length. In ideal systems, chemical diffusion may be quite effective in limiting compositional gradients due to the large change in chemical potential with composition. In strongly non-ideal systems however, which contain P–T–X regions where d mi/d X ≈ 0, and hence the driving force for chemical diffusion is negligible, experiments may exhibit large gradients in major-element composition. The magnitude of the compositional gradient is described by the Soret coefficient (σ), which for a binary system is given as (Lesher and Walker 1986): σ=

X Hot − X Cold X (1 − X ) ∆T

(17)

where XHot and XCold are the component mole fractions measured at the hot and cold ends of the sample respectively and X is the average component mole fraction in the sample. Trace elements in the system are distributed to maintain a constant activity along the sample and concentration gradients therefore depend upon changes to the activity coefficient with major element composition (Jones and Walker 1990). The following derivation relates the compositional gradient for a trace element (i), measured in a Soret diffusion experiment, to b. Values of b obtained from these experiments may then be used to predict partitioning through the use of Equation (13) (Brenan and Bennett 2010). Assuming a hypothetical solid Fe phase in equilibrium with the melt at all points along a Soret diffusion experiment, the solid metal–liquid partition coefficient for i can be defined at two points along the sample (T1 and T2) as: 0 LM ln DiSM/LM ,T1 = ln Di + ln γ i ,T1

(18)

0 LM ln DiSM/LM ,T2 = ln Di + ln γ i ,T2

(19)

Subtracting Equation (19) from Equation (18) and taking the exponent yields: γ iLM D SM/LM ,T1 ,T1 = iSM/LM LM γ i ,T2 Di ,T2

(20)

Substituting XiSM / XiLM for DiSM/LM :

( (

LM XiSM γ iLM ,T1 / X i ,T1 ,T1 = LM γ iLM XiSM ,T2 ,T2 / X i ,T2

) )

(21)

SM For adjacent positions along the sample, it can be shown that XiSM ,T1 / X i ,T2 ≈ 1, simplifying Equation (21) to give:

γ iLM XiLM ,T1 ,T1 = γ iLM XiLM ,T2 ,T2

(22)

Brenan, Bennett & Zajacz

12

Taking the natural log of Equation (22) and substituting Equation (11) for the activity coefficients relates b for trace element i to a given mole fraction of light element (Xl): LM ln XiLM ,T1 − ln X i ,T2 =−βi   ln (1 − nαX l ,T1 ) − ln (1 − nαX l ,T2 ) 

(23)

∆ ln X iLM = −βi ∆  ln (1 − nαX l ) 

(24)

Or more simply:

Values of b are therefore extracted from plots of ln XiLM vs. ln (1 − nαXl ) and are assumed to be temperature independent. Evidence for the lack of a strong T dependence for b is provided by previous equilibrium experiments (Jones and Malvin 1990; Chabot and Jones 2003) and the linearity of data plotted in the manner described above (Brenan and Bennett 2010). An example of the major and trace element data produced by Soret experiments is provided in Figure 3. Soret diffusion experiments are subject to the same design considerations as those described previously for ambient pressure and high pressure studies. The primary advantage of this approach is that a small number of experiments can provide information comparable to a large suite of isothermal solid–liquid partitioning experiments. A further application of this method is to determine the effects of light-elements that have a low solubility in Fe-rich melt. Phase relations in the Fe–O system for example, preclude large volume experiments to directly measure solid metal–liquid metal partition coefficients for liquids that span a broad range of O-contents (i.e., ≤ 2.2 wt% O at 15 GPa, Langlade et al. 2008). Soret experiments, however, can be performed at super-liquidus temperatures where light-element components more readily dissolve in the melt. This can provide access to compositional space that may be relevant to high-pressure planetary differentiation processes. A disadvantage of this approach, however, is the lack of information for the solid phase, which also exerts a strong control on inter-element fractionation, as noted above. 35 30

-6

SOR 12 SOR 21

-8

20 15

-9

-10

10

-11

5 0

slope = ß

-7

SOR 24

ln XRe

wt% S

25

-5

SOR 3

1750

1850 T (oC)

1950

a

2050

-12 -13 -1.4

-1.2

-1.0 -0.8 ln(1-2αXs)

-0.6

b

-0.4

Figure 3. The results of Soret experiments performed by Brenan and Bennett (2010) in the Fe–Ni–S system. Panel (a) displays the change in sulfur content as a function of T along the sample length, that arises due to the Soret effect. Panel (b) displays the corresponding gradients in Re concentration, that result from trace element redistribution to maintain isoactivity along the sample length. Values of bRe are determined from the slope defined by linear regression of ln XRe vs. ln(1 − 2aXs) for each experiment.

Experimental Fractionation of the HSE

13

HSE SOLUBILITY EXPERIMENTS: IMPLICATIONS FOR METAL–SILICATE PARTITIONING Accretion of the Earth from planetesimals of chondritic composition (McDonough and Sun 1995; Wood et al. 2006), with concurrent differentiation into a metal core and silicate mantle, is generally thought to have occurred over the first ~ 30 Ma of Earth’s history (e.g., Kleine et al. 2002; Yu and Jacobsen 2011). During accretion, heat generated by the decay of short-lived isotopes and the collision of large impactors is likely to have raised global temperatures sufficiently to cause widespread melting and the formation of a magma ocean, through which more dense Fe–Ni liquid could descend (e.g., Ringwood 1966; Karato and Murthy 1997; Wood et al. 2006; Rubie et al. 2007). In this scenario, the siderophile elements would be transported to the growing core, leaving a depleted silicate mantle with element ratios that depend upon the differing affinities for the metal phase (e.g., Chou 1978; Newsom and Palme 1984; Wänke et al. 1984). Past work has ascribed the behavior of the moderately siderophile elements (MSEs; Mo, W, Cr, V, Mn, etc) to a combination of metal extraction and accretion of compositionally distinct components (Wänke and Dreibus 1988; Schmitt et al. 1989; O’Neill 1991). More recently it has been shown that, if metal–silicate partitioning is appropriately parameterized, a match to mantle MSE abundances can be achieved at appropriately high pressure and temperature (e.g., Wade and Wood 2005; Righter 2011; Siebert et al. 2011; Wade et al. 2012). In contrast, metal–silicate partitioning of the HSE has provided a somewhat conflicting view on the accretion model. The majority of past work on HSE partitioning has been at 0.1 MPa and relatively low temperature (~ 1300−1400 ºC) and showed that at the relatively reduced fO2 attending core formation1 (i.e., 1.5 log units more reduced than the iron–wustite buffer; IW − 1.5) metal–silicate partitioning of all the HSEs is likely to exceed 105−108 (e.g., Borisov and Palme 1995, 1996; Ertel et al. 1999, 2001). Such results predict the quantitative removal of the HSEs from the silicate mantle, which is inconsistent with their estimated abundances in the primitive mantle (Becker et al. 2006). This apparent lack of mantle HSE depletion is a long-standing issue in geochemistry, and has come to be known as the “excess siderophile element problem”. Importantly, moderate to large differences in the relative mantle abundances of the HSEs are also predicted by the low- temperature partitioning data (compare Borisov and Palme 1996 to Ertel et al. 1999), in conflict with observed chondritic relative abundances (Becker et al. 2006) and a mantle Os-isotope time evolution requiring Pt/Os and Re/Os ratios to match those of chondritic meteorites (Meisel et al. 2001; Brandon et al. 2006). These combined observations support a model for late accretion of a small amount of dominantly chondritic material which postdated core formation; the so-called later veneer (e.g., Kimura et al. 1974). The veracity of this model came into question, however, following the proposal of Murthy (1991) that there should be a convergence of metal–silicate partitioning values at the very high temperatures likely during core formation (i.e., > 2700 ºC), resulting in a mantle HSE composition set by metal–silicate equilibrium. Although the strategy to estimate high temperature partitioning was shown to be flawed (e.g., Capobianco et al. 1993) more recent results have indeed shown that temperature is a key variable to be considered, as is the importance of measurements at reduced conditions, at which some HSE may exhibit unanticipated and unusual redox behavior, as described in the subsequent sections. 1 In this chapter two different, but related, fO2 references are used. The first, fayalite–magnetite–quartz (FMQ) denotes relatively oxidized conditions, and is near the fO2 of most terrestrial magmas. Iron–wustite (IW) is the other reference point, corresponding to highly reduced conditions, with the conversion that IW is approximately 3.5 log units more reduced than FMQ, i.e., IW ~ FMQ − 3.5.

14

Brenan, Bennett & Zajacz CALCULATING THE METAL–SILICATE MELT PARTITION COEFFICIENT FROM SOLUBILITY DATA

Experiments that equilibrate silicate melt with an Fe-rich metal liquid usually result in vanishingly low HSE concentrations in the quenched silicate melt, owing to high values for metal/silicate partition coefficients, DiMet/Sil . For example, a sample doped with 200 ppm of element i and comprising equal fractions of metal and silicate melt will result in CiSil < 0.004 ppm for values of DiMet/Sil > 105. Hence, in order to generate run-products with measurable HSE levels in the quenched silicate phase, solubility experiments are performed. In this case, pure HSE metal, or an HSE-rich alloy, is equilibrated with silicate melt at the desired P–T–fO2 conditions and values of DiMet/Sil are calculated using the formulation of Borisov et al. (1994), described as follows. The dissolution reaction for a metal in silicate melt is: n (25) M ( Met ) + O2 ( g ) = M n + O n / 2 ( Sil ) 4 where n is the oxidation state of the dissolved metal cation. Two equilibrium constants can be defined for this reaction, one for an experiment at saturation in an HSE phase (KSat) and the other pertaining to a natural system with dilute HSE concentrations (KDil): n Sat  − log fO2 = log K Sat log  aMO n /2  4 n = log  aMOn /2  − log [ aM ] − log fO2 log K Dil 4

(26) (27)

where aM and aMOn/2 are the activity of the metal and metal oxide and components. Substituting a= Xi γ i into Equations (26) and (27), where X i and γ i are the mole fraction i and activity coefficient respectively, yields: n Sat  + log  γ Sat = KSat log  X MO log MO n /2  n /2   − 4 log fO2

(28)

n = log K Dil log  X MOn /2  + log  γ MOn /2  − log [ X M ] − log [ γ M ] − log fO2 4

(29)

��

��

At fixed P and T, subtracting Equation (28) from (29) and rearranging gives:  γ MOn/2   X MOn /2  Sat log =  log  X MOn/2  + log [ γ M ] − log  Sat  X  M   γ MOn/2 

(30)

The molar partition coefficient (D*Met/Sil = XMet / XSil) may then be substituted into the left-hand side of Equation (30), yielding:  γ MOn/2   1  Sat  + log [ γ M ] − log  Sat log  = log  X MO  *Met/Sil  n/2  D    γ MOn/2 

(31)

The very low solubility of HSEs in silicate melt, even when saturated in a pure metal phase, results in negligible changes to γ MOn /2 over the possible range for X MOn/2 and the ratio γ MOn/2 / γ Sat MO n/2 may therefore be treated as unity. After taking the exponent, Equation (31) then becomes: 1 Sat = X MO γM n/2 D*Met/Sil

(32)

Experimental Fractionation of the HSE

15

Typically. it is the concentration by weight rather than mole fraction of a trace element that is reported and conversion of Equation (32) to the weight-based, rather than molar, partition coefficient is accomplished using the following conversion factors (A and B):

X M= A × CM ( Met )

(33)

X MOn/2= B × CM ( Sil )

(34)

where CM is the concentration of trace element M in the indicated phase. Rearranging, and substituting Equations (33) and (34) into Equation (32) yields:

 1 B  B   D Met/Sil =   D*Met/Sil =    Sat   A  A   B × C M ( Sil ) × γ M 

(35)

which simplifies to: D Met/Sil =

1 A×C

Sat M ( Sil )

(36)

× γM

where CMSat is the concentration of M measured in the silicate melt at the solubility limit and γ M is the activity of M at infinite dilution in Fe metal. Values of A tend toward a constant value as the concentration of M in Fe-alloy decreases.

Controls on the metal–silicate partition coefficient Temperature, pressure, oxygen fugacity and melt composition may all play a role in determining the value of DMet/Sil. The equilibrium constant for Reaction (26), describing the dissolution of a trace metal into silicate melt may be equated with the Gibbs free energy of  reaction ( ∆Gr ) as follows: n ∆Gr = ln K = ln  aMOn/2  − ln [ aM ] − ln fO2 − RT 4 Replacing the activity terms ( a=i Xi γ i ) and rearranging yields:  X ln  M  X MO n /2 

 ∆Gr n  γ MOn /2  − ln fO2 + ln   =  − T R 4  γM  

(37)

(38)

Expanding the free energy term ( ∆Gr =∆Hr − T ∆Sr + P∆Vr ) and substituting the molar partition coefficient into the left hand side of Equation (38) reveals the variables that might be expected to affect HSE partitioning: ln D*Met/Sil =

 γ MOn /2  ∆Hr ∆Sr P ∆Vr n − + − ln fO2 + ln   RT R RT 4  γM 

(39)

where ∆Hr , ∆Sr and ∆Vr are the enthalpy, entropy, and volume change of reaction respectively. Assuming these three parameters do not themselves depend strongly on P or T, Equation (39) can be simplified to yield the following relationship:  γ MOn /2 a bP n ln D*Met/Sil = + − ln fO2 + ln  T T 4  γM

 +c 

(40)

where a, b, and c are constants determined by regression of the experimental data. The partition coefficient is therefore expected to depend on temperature, pressure, oxygen fugacity

16

Brenan, Bennett & Zajacz

and the composition of both the silicate and metallic melt. Equation (40) provides the basis for the various approaches employed in past work to parameterize the results of either metal solubility measurements or direct determinations of metal–silicate partitioning. A summary of such work is provided in Table A1 in the appendix. Whereas the dependence of solubility and partitioning on fO2 is reasonably well established under oxidizing conditions (i.e., > FMQ), experiments done at reducing conditions often display strongly disparate results. The origin of this variation is thought to arise from the presence of dispersed metal inclusions in silicate run-products2. This effect has introduced some uncertainty in the accuracy of past measurements, of both metal solubility and mineral-silicate melt partitioning (see Silicate and Oxide Control on HSE Fractionation). Hence, before reviewing the partitioning results for the various HSEs, it is instructive to briefly describe the inclusion problem, and efforts to overcome it in experiments.

Metal inclusions in experiments and the analysis of contaminated phases The presence of a dispersed metallic phase contaminant in quenched silicate melts from solubility experiments has been recognized since the earliest efforts to determine values of DMet/Sil for the HSE (Kimura et al. 1974). An array of different particle sizes have been implied, ranging from ~ 0.05 mm to ~ 5 mm, also with varied spatial distributions (compare Ertel et al. 1999; Cottrell and Walker 2006; Fortenfant et al. 2006; Mann et al. 2012). Owing to their small size and dispersed nature, it has been difficult to fully document the characteristics of these inclusions, even by high resolution electron microscopy. Hence, the most commonlyused indication of their presence has been poorly reproducible solubility measurements and heterogeneous time-resolved LA-ICPMS spectra (e.g., Ertel et al. 2001, 2006). Any study of solubility and partitioning where metal inclusions are suspected must therefore assess their presence as either an exsolved, but once dissolved, component of the silicate melt, or instead a discrete metal phase at the conditions of the experiment (hereafter referred to as ‘quench’ and ‘stable’ inclusions respectively). If the metal particles have the latter origin, then their contribution to the analyses of silicate run-products whose intrinsic HSE concentration is very low will increase the apparent solubility. This issue embodies much of the ambiguity over the true solubility and partitioning of HSEs at reducing conditions. Early efforts to determine HSE solubility used bulk analytical techniques (e.g., neutron activation analysis, scintillation spectrometry), which required extra care to ensure a minimum level of metal contamination, including reversal experiments, measurements on different aliquot sizes, thorough cleaning of the exterior portion of glass shards, etc. For the cases of Pd (Borisov et al. 1994) and Au (Borisov and Palme 1996), solubility determined by these methods produced reproducible results, and systematics with fO2 consistent with thermodynamic expectations. Other metals, such as Ir, Os, and Re, produced more scattered results, which is now suspected to be a result of metal contamination (Ertel et al. 2001; Fonseca et al. 2011). The arrival of LA-ICPMS as a readily available tool for trace element analysis (e.g., Jackson et al. 1992) provided an alternative, in situ, approach to solubility measurements. The time-resolved spectra produced during sample ablation revealed metal inclusions manifested as high count-rate ‘peaks’ separated by low count-rate ‘troughs’, the latter thought to more closely represent the intrinsic signal from metal dissolved in the silicate (Fig. 4). Under this pretext, several studies have filtered their analytical results by calculating concentrations from only the “trough” portion of the spectra, (e.g., Ertel et al. 2001, 2006; Laurenz et al. 2013), yielding reasonably systematic relations between solubility and fO2 at conditions at least as reducing as the FMQ buffer. How closely the low count-rate portions of spectra represent truly inclusion-free silicate however, will depend upon the nature of both the sample (inclusion size, spatial distribution) and the analytical 2 Note that there is a similar problem encountered in sulfide-bearing silicate systems, further described in the Section Sulfide–melt/silicate melt and MSS–silicate melt partitioning.

Experimental Fractionation of the HSE

195

Pt / 43Ca (counts / sec)

a

195

Pt (counts / sec)

b

101

No Ablation

17

Ablation

100

10-1

10-2

10-3 0 105

10

20

No Ablation

30

40

50

60

70

80

Ablation

104

Minimum in signal used to determine concentrations

103

102

0

Figure 4. Time-resolved LA-ICPMS spectra for Pt in the silicate portion of inclusion-contaminated experiments by Bennett et al. (2014) (a) and Ertel et al. 2006 (b). In panel (a), counts per second Pt are normalized to the signal for Ca, an abundant lithophile element that is unaffected by the presence of metal inclusions. This procedure isolates heterogeneity arising due to the ablation of metal inclusions from instrumental signal variability. The small grey area in panel (b) shows the region used for data-reduction by Ertel et al. (2006) to determine ‘filtered’ Pt concentrations (see Metal Inclusions in Experiments and the Analysis of Contaminated Phases for details).

10 20 30 40 50 60 70 80 90 100 t (s)

conditions (spot-size, wash-out time of the ablation cell). Most importantly, solubilities determined using this method still rest upon assigning all of the heterogeneity to inclusions having a stable origin, which does not account for the possibly of some exsolved metal. The difficulty associated with analysis of inclusion-contaminated silicate run-products has driven attempts to suppress the formation of the metal particles (Borisov and Palme 1995; Ertel et al. 1999, 2001; Borisov and Walker 2000; Ertel et al. 2001; Brenan and McDonough 2009; Bennett and Brenan 2013; Bennett et al. 2014; Médard et al. 2015). Due to the success of these efforts, the database of inclusion-free solubility measurements is now sufficient to assess the oxidation state of most HSEs dissolved in silicate melt at reducing conditions, and derive accurate values of metal–silicate partitioning. The experimental and analytical approaches used to generate this database are outlined in the section Experimental methods to measure HSE solubility and metal/silicate partitioning. Prior to describing the methods to measure partitioning, and their results, a brief review is provided of some of the proposed mechanisms for inclusion formation.

Possible mechanisms of metal inclusion formation A diverse range of mechanisms have been proposed for the formation of dispersed metal inclusions in glasses synthesized at high temperature and low fO2. In the following discussion two categories of metal inclusions are considered: 1) ‘quench’ inclusions that form by exsolution from the silicate melt during the cooling step that accompanies the termination of an experiment and 2) ‘stable’ inclusions that are present as a discrete metal phase at the P and T of the experiment.

18

Brenan, Bennett & Zajacz

Quench metal inclusions. HSE solubility experiments are terminated by rapid cooling of the sample, so if the solubility of the HSE is prograde, then this step may cause oversaturation in HSE metal. In some instances, the silicate melt does not quench to a glass, so HSE metal may also form by local saturation due to a build-up in concentration in the residual melt by the crystallization of phases that exclude the HSE. The increase in solubility with temperature documented for nearly all of the HSE (See: Summary of experimental data) suggests that if quench rates are sufficiently slow, then HSE metal grains will begin to nucleate and grow. Cottrell and Walker (2006) outline several expectations for inclusions formed in this manner: an increase in diameter with decreased cooling rate, spatial variability in the distribution of inclusions due to differences in cooling rate across the sample, and a compositional difference between inclusions and the bulk HSE ± Fe source metal. Cottrell and Walker (2006) document all of these features in experiments to measure Pt solubility done at 2.2–2.3 GPa and 1940–2500 ºC, providing a strong argument that the inclusions observed in their silicate run-products were formed by exsolution when the sample was quenched. These authors also measured comparable Pt concentrations in portions of the sample both with and without visible contamination by metal inclusions; consistent with a spatial variability in the quench rate within the sample. In the metal–silicate partitioning study of Mann et al. (2012), the silicate melt phase did not quench to a glass, but instead to a fine intergrowth of quench crystals. This is a common result for melt compositions which are poor in network-forming cations, as in this case. Metal particles were found to be concentrated along the boundaries of skeletal olivine crystals that formed during quenching. These particles were interpreted to form by HSE buildup during quench crystallization, and included as part of the original high P–T melt composition, on the basis of uniform element concentrations measured across the sample and smooth correlations between DMet/Sil and fO2. It is noteworthy that even if metal-inclusions in the Mann et al. (2012) study did not form exclusively upon quench, and values of DMet/Sil were thus underestimated, Re, Ir and Ru are still found to be too siderophile to account for their PUM abundance. Although there is good evidence for the formation of a metal phase during quench in both these previous studies, it remains possible that experiments also contain stable inclusions. However, at the very high experimental temperatures, and associated high solubilities, the relative contribution of stable inclusions to the measured concentrations is likely to be small. Stable metal inclusions. Unlike inclusions that form when the sample is quenched, stable metal inclusions are expected to display a relatively uniform spatial distribution and possess the same composition as that measured for the bulk metal source. Although cooling rate should not have any effect on the nominal size of stable inclusions, it is foreseeable that they may act as sites for heterogeneous nucleation of metal precipitated by over-saturation of the melt during quench. The small size of stable inclusions can render them undetectable by scanning electron microscopy, preventing their distribution and composition from being accurately determined and their origin confirmed. Conclusive evidence that metal inclusions can form as a stable phase in experiments has now been provided by Yokoyama et al. (2009) and Médard et al. (2015). Yokoyama et al. (2009) measured metal–silicate partitioning of Os using a natural meteorite starting material. The Os isotopic compositions of inclusion-contaminated run-product glasses define a mixing array between the meteorite starting material, and inclusion-free glass. This implies that inclusions in these experiments possess the same isotopic composition as the meteorite starting material. The measured array in Os concentration and isotope composition therefore reflects different glass to inclusion ratios in the individual aliquots. By contrast, inclusions formed upon quench should resemble the isotopic composition of the silicate melt and not define mixing lines that extend to the meteoritic starting material. Médard et al. (2015) performed solubility experiments in a centrifuge piston-cylinder, and found that the Pt content of melts subject to high acceleration was lowest in the top-most portion of the sample, suggesting partial segregation of stable metal particles to the base of the sample. Combined, these results confirm the assertions of earlier studies that suggest a stable origin for inclusions on the basis of sample heterogeneity (Borisov and Palme 1997; Ertel et al. 1999, 2006).

Experimental Fractionation of the HSE

19

Perhaps the most obvious means to introduce metal particles into silicate melt is by erosion of the metal source. This mechanism was advocated by Yokoyama et al. (2009) based on the similarity in the isotopic composition of the metal contaminant with the metal phase added to experiments. In contrast, Ginther (1971) proposed that contamination of glass by platinum inclusions occurred via oxidation of the Pt container, caused by the dissociation and evaporation of alkali or alkali-earth oxide complexes in the melt. This mechanism was based on the observation that inclusions were restricted to surface layers of the unstirred melt and independent of the fO2 of the atmosphere within which the experiment was performed. Borisov and Palme (1997) noted that this mechanism would be most effective under reducing conditions, where alkali metal evaporation is enhanced, and may potentially explain the formation of Pt and Ir inclusions in their low fO2 experiments. Although this may be a viable process in open-system experiments performed using gas-mixing at ambient pressure, it cannot explain the formation of metal inclusions in experiments done at high confining pressure and is unlikely to afflict experiments performed in vacuum-sealed silica tubes. Metallic Pt inclusions have been well documented during the synthesis of large volumes (> 0.5 L) of phosphate laser glass. Campbell et al. (1989) summarized previous studies of laser glass synthesized in Pt containers and concluded that thermal gradients were responsible for the formation of metal inclusions. The experimental design used in these studies involved induction heating of the sample in a Pt crucible. Once the sample was molten, more feedstock was added to increase the level in the crucible. Campbell et al. (1989) identify relatively large thermal gradients in this arrangement, and owing to the increase in Pt solubility with temperature, such gradients will force precipitation of metal inclusions. Most petrologic experiments done at 0.1 MPa however, are performed using small sample sizes (< 2 mL), in the absence of significant thermal gradients, suggesting that this mechanism does not account for the formation of inclusions in most of the studies discussed as part of the present work. The formation of metal particles in experiments containing certain alloy-forming impurities is a mechanism proposed by Borisov and Palme (1997) that may potentially apply to both ambient- and high-pressure studies. In this scenario, dissolved oxide components in the melt react with the HSE source material to produce oxygen and an HSE alloy as the product. For example the reaction proposed by Borisov and Palme (1997) for Pd is:

Pd ( Met ) + xSiO2 ( Sil ) = PdSi x ( Met ) + xO2 ( g )

(41)

Decreasing fO2 therefore favours the formation of stable alloy grains. Borisov and Palme (1997) suggest that the presence of impurities that react more readily with HSEs than Si, such as As, Sb, Bi, Ge, Sn, and Pb, may increase the likelihood of alloy formation. In open-system gas mixing experiments, the volatile nature of potential alloying elements should lead to progressive loss at high T, resulting in metal inclusions free of these impurities by the end of an experiment (Borisov and Palme 1997). Borisov and Walker (2000) made use of these concepts and removed volatile contaminants from starting materials by fusion at controlled fO2 conditions. This resulted in inclusion-free run-product glasses produced at a fO2 lower than past work. Although it may have been the removal of contaminants from reagents that prevented inclusions from forming, pre-reduction of the sample may also have played a role (see below). Fortenfant et al. (2006) found that erosion of the sample crucible or stirring spindle was an unlikely source of the Os inclusions in their experiments, after identifying components in an inclusion that were not present in the labware used to perform the experiment. On this basis, the authors instead suggested that inclusions originate as a result of exsolution from silicate melt, driven by changes in sample fO2. Bennett et al. (2014) developed this idea and proposed a mechanism for stable metal inclusion formation based on the time-evolution of fO2 during the initial stages of an experiment. Most metal solubility experiments use starting materials that are fully or partially oxidized (i.e., Fe present as Fe2+, Fe3+ and sample capsules loaded

Brenan, Bennett & Zajacz

CHSE

20

Oversaturation of HSE in melt

Equilibrium HSE

fO2

concentration in melt

Equilibrium fO2

t Silicate Melt Growth

Stable Inclusions

Bulk HSE Source

Figure 5. Schematic for the formation of metal inclusions as described by Bennett et al. (2014). The process can be subdivided into three steps, 1) initial, high sample fO2 and rapid HSE in-diffusion; 2) reduction in sample fO2 causes oversaturation in HSE and precipitation of metal inclusions; 3) equilibrium fO2. conditions, small size of inclusions prevents gravitationally-driven segregation. Low solubility also hinders inclusion growth by Ostwald ripening (Médard et al. 2015). [Used by permission of Elsevier Limited, from Bennett, Brenan and Koga (2014) Geochimica et Cosmochimica Acta, Vol 133, Fig. 2, p. 427.]

in air), but then subject to a reducing atmosphere or encapsulated in a reducing material (e.g., graphite). Once the sample is heated, reaction between the sample and the reductant occurs, and the initially high fO2, imposed by the starting materials, begins to fall. For the case of a graphiteencapsulated experiment initially undersaturated in a fluid phase, the redox reaction is: C (s) + 3 O2 ( g) = CO32 − ( Sil ) 2

(42)

This reaction will proceed until an equilibrium CO32 − concentration is obtained.3 With reference to Equation (25), the dissolution of HSEs in silicate melt requires oxidation of the metal phase, hence reduction of oxidized starting materials by Reaction (42) causes an accompanying decrease in metal solubility. If the kinetics of HSE dissolution are suitably rapid, the silicate melt may be endowed with elevated metal concentrations initially, but as fO2 drops, saturation occurs and HSE grains may precipitate. This process is portrayed schematically in Figure 5. Médard et al. (2015) note that in all past experiments in which stable metal inclusions are observed, the equilibrium fO2 of the experiment is lower than that of the starting materials, consistent with the operation of the proposed reduction mechanism. Borisov and Walker (2000) 3 The carbonate anion is chosen in this case to represent the dissolved carbon species as it has previously been identified as an oxidized carbon species present in silicate melts (e.g., Mysen et al. 2011).

Experimental Fractionation of the HSE

21

also reported that experiments in which starting glasses were synthesized at the same conditions as the subsequent solubility determination were less susceptible to contamination by Os metal inclusions—procedures that would have suppressed an initially oversaturated state.

Experimental methods to measure HSE solubility and metal–silicate partitioning A simple method for measuring HSE solubility is by suspending a silicate melt bead within a wire loop, or encapsulated in foil, made of the pure metal, then equilibrating the sample in a gas-mixing furnace at high temperature. Samples are rapidly quenched, then analysed for the HSE by either bulk (e.g., Borisov et al. 1994; Borisov and Palme 1995, 1997) or in situ methods (Laurenz et al. 2013). Variations on this approach involve the use of alloys, whose composition either allow higher equilibration temperatures without melting (e.g., Pd–Au; Borisov and Palme 1996) or involve a combination of metals, one of which is wire-forming (e.g., Ni) with those that are not (e.g., Os; Borisov and Palme 1998; Borisov and Walker 2000). Borisov and Walker (2000) modified this technique to suppress the formation of stable Os metal nuggets by using Ni–Os alloy in the wire loop, and pre-saturating the silicate melt in Ni prior to the solubility experiment. This pre-saturation step was considered to suppress chemical erosion of the Ni–Os loop, and subsequent entrainment of metal particles in the melt. Note that for experiments employing alloy source material, HSE concentrations measured in the silicate melt must be corrected to solubilities corresponding to equilibrium with the pure HSE phase  Sat CMMeas(Sil )   CM (Sil ) = .  aM ( Met )   The mechanically assisted equilibration (MAE) method allows for the approach to metal– silicate melt equilibrium to be monitored over the course of an experiment, while the sample is subject to continuous stirring (Dingwell et al. 1994). In this technique, a relatively large mass of silicate melt is contained in a metal crucible within the hot zone of a gas-mixing furnace. The melt is subject to forced convection by a rotating metal spindle suspended axially, with material sampled from the crucible sequentially so as to obtain a solubility vs. time history. The purpose of the rotating spindle is to promote metal–silicate equilibration involving advective as well as diffusive exchange with the metal source. An extensive review of this technique as applied to HSE solubility measurements is given by Ertel et al. (2008). Despite this advance, however, it has been difficult to obtain reproducible solubility measurements at fO2 more reducing than FMQ. Brenan and McDonough (2009) and later Bennett and Brenan (2013) were able to suppress the formation of stable Ir, Os, and Re inclusions in high-pressure (2 GPa) experiments by adding these metals encapsulated in a pre-melted gold bead. This approach yields a unique experimental geometry that changes the relationship between fO2 evolution and HSE indiffusion during the early stages of an experiment. Molten Au strongly wets Re, Os, and Ir at high T to form a rind that physically separates the bulk HSE source in the experiments from the silicate melt (Fig. 6). This rind slows the diffusive transfer of the encapsulated metal into the silicate melt, thus allowing the melt to undergo reduction prior to the onset of metal dissolution. This technique has proved successful in suppressing metal inclusion formation under conditions as reducing as IW − 1.2 (Brenan and McDonough 2009). One tenet of the Au rind method is that the HSE of interest must be sparingly soluble in molten Au, so as to limit the HSE flux to the silicate melt. It is therefore unsuitable for investigating elements such as Pt and Pd, which display complete miscibility with Au at high temperature. Bennett et al. (2014) performed experiments to measure Pt solubility in molten silicate by adding elemental Si to starting materials, which serves to strongly suppress any initial oxidation of the melt, and thereby inhibit the formation of inclusions by initial oversaturation. The veracity of this approach was tested in experiments done both with and without added Si, with the timeresolved analysis of run-products as a guide to inclusion abundance. Control experiments, which

Brenan, Bennett & Zajacz

22 a

b Graphite Capsule

c

Silicate Glass

Silicate Melt

x

Au+HSE Bead Au Liquid ~750 µm

Solid HSE

Figure 6. Example of a run-product from a piston-cylinder experiment to measure the solubility of Au and Os in molten silicate at 2000 ºC and 2 GPa. (a) Reflected light image of the run-product showing the quenched silicate melt (now glass) surrounded by the graphite capsule. At the bottom of the glass portion is a bead of gold, quenched from the liquid state, encapsulating grains of osmium metal, which were solid during the experiment (Tmelting for Os = 3027 ºC). (b) Close-up of outlined area in (a) showing the efficient wetting of the osmium grains by the gold melt. (c) Schematic showing the distance, x, through which the HSE of interest must diffuse before HSE begins to dissolve into the melt.

had no silicon added, display significant heterogeneity of the time-resolved spectra for Pt, in contrast to the uniform (and relatively low intensity) signal for Pt in experiments done with 0.75–2 wt% added Si. Although this technique was successful in producing uncontaminated samples as reduced as IW − 1.6 at temperatures ≥ 1900 ºC, inclusions were identified in experiments at 1800 ºC and ~ IW, suggesting this method may be less effective at lower temperatures. The small size of stable inclusions and the low solubility of the HSE in general would seem to preclude the segregation of metal grains from silicate melt without the assistance of stirring; an effect not easily achieved in high pressure experiments. Subjecting the molten sample to high acceleration however, increases the settling velocity of dense particles from the silicate melt. With this in mind, Médard et al. (2015) used a piston-cylinder mounted in a centrifuge to perform Pt solubility experiments at accelerations of up to ~ 1500 g0. Run products from synthetic Fe-free melts achieved partial cleansing by this method. Experiments done with a natural FeO-bearing melt yielded similar solubility for static and high-g0 experiments, which was attributed to the role of dissolved FeO as an oxygen donor to facilitate. In addition to experiments which have measured the solubility of pure metals, or binary metal alloys, then calculating partition coefficients from Equation (36), other work has focused on more direct determinations using Fe-rich alloy compositions. As gold is the most soluble of the HSE in molten silicate, it is possible to measure Fe metal–silicate melt partitioning for relatively dilute Fe–Au alloys (i.e., < 4 wt% Au), for which results have been reported by Brenan and McDonough (2009) and Danielson et al. (2005). A comparison of partitioning calculated from the solubility of pure Au with this more direct method has yielded very good agreement (see Effect of oxygen fugacity). Other experiments have also involved Fe-bearing multicomponent alloys, but with the HSE more concentrated (Mann et al. 2012). In that case, the challenge is access to an accurate solution model for extrapolating apparent partition coefficients to infinite dilution.

Summary of experimental data The existing database of HSE solubility and partitioning data contains significant complexity, arising from the issues surrounding metal inclusions discussed above. To simplify the following overview of HSE behavior as a function of P, T, X, and fO2, only those data which are demonstrably free from contamination or have been shown to agree well with the data

Experimental Fractionation of the HSE

23

from such experiments (e.g., data from filtered LA-ICP-MS spectra) are presented. Unless otherwise stated, the oxidation states quoted below have been determined by linear regression of solubility vs. fO2 plots using the following relationship: log= CMSat (sil )

n log fO2 + Constant 4

(43)

Effect of oxygen fugacity–Rhenium: Several studies have determined the solubility of Re as a function of fO2 and data now exists at 0.1 MPa to 18 GPa and 1400–2500 ºC (Fig. 7). At 0.1 MPa and above ~ IW + 1, Re is thought to be dissolved as a mixture of 4+ and 6+ species, with 6+ being dominant across most of the fO2 range (Ertel et al. 2001). This result is consistent with crystal–melt partitioning experiments that also indicate 4+ and 6+ species for Re above ~ IW (Mallmann and

2.0

4.0

Au

4.0

Pd

Pt

210

2.0

2.0

1.0

2+ 2-

0.0

-1.0 -2

log CHSE (ppm)

3.0

0.0

1+

2+

0

2

4

6

8

10

2+

-2.0 -3 3.0

Ir

-1

1

3

5

-2.0 -3

7

1.0

Rh 1+

2.0

1.0

-1

1

3

5

7

9

11

4

5

Os

0.0

2+

1.0

2+

-5.0 -2

3+

0

2

4

0.0

-2.0

-1.0

-3.0

1+

1+

-3.0

-1.0

3+

-1.0

3.0

0.0

0

0

6

8

10

Ru

-2.0 -2

3.0

2.0

0

2

4

6

8

-4.0 10 -2

Re

0

1

1755 K

1573 K

1673 K 1687 K

-1.0

3+ 4+

1

3

5

7

9

-3.0 -2

Δ IW

Bennett et al. 2014 Borisov & Palme 1995

1573 K

1623 K

Borisov & Walker 2000

-1.0

3

Borisov & Palme 1996

Borisov & Palme 1997

1.0

2+

2

Bennett & Brenan 2013

0.0

-2.0 -1

-1

Borisov & Nachtweyh 1998 1755 K

1.0

3+ 4+

0

1

2

Brenan & McDonough 2009

Cottrell & Walker 2006

Ertel et al. 1999

2+

Ertel et al. 2001

Fonseca et al. 2011

4+

Fortenfant et al. 2003

Fortenfant et al. 2006

Laurenz et al. 2010

Laurenz et al. 2013

6+

-1

Borisov et al. 1994

3

Médard et al. 2015

4

5

6 GPa

18 GPa

Righter & Drake 1997

Mann et al. 2012 Righter et al. 2008

Figure 7. Solubility vs. fO2 relative to the iron–wustite buffer for all the HSEs, after correction to unit activity of HSE metal. Also indicated are slopes for the likely oxidation states of these elements in the melt. Data are from the following sources and are ambient pressure results unless otherwise noted: (Au) Borisov and Palme (1996) 1300−1480 ºC; Brenan and McDonough (2009) 2 GPa, 2000 ºC; Bennett and Brenan (2013) 2 GPa, 2000 ºC; (Pd) Borisov et al. (1994) 1350−1415 ºC; Laurenz et al. (2010) 1300 ºC; Mann et al. (2012) 6−18 GPa, 2150−2200 ºC; (Pt) Borisov and Palme (1997) 1400 ºC; Ertel et al. (1999) 1300 ºC; Cottrell and Walker (2006) 2.2−2.3 GPa, 2000 ºC; Mann et al. (2012) 6 GPa, 2150 ºC; Bennett et al. (2014) 2 GPa 2000 ºC; Médard et al. (2015) 1.2 GPa, 1900 ºC; (Ir) Borisov and Palme (1995) 1300−1480 ºC; Brenan and McDonough (2009) 2 GPa, 2000 ºC; Fonseca et al. (2011) 1500 ºC (Rh) Ertel et al. (1999) 1300 ºC; Fortenfant et al. (2003) 1300 ºC; Mann et al. (2012) 6−18 GPa, 2150−2200 ºC; (Os) Borisov and Walker (2000), 1400 ºC; Fortenfant et al. (2006) 1350 ºC; Brenan and McDonough (2009) 2GPam 2000 ºC; (Ru) Borisov and Nachtweyh (1998) 1400 ºC; Laurenz et al. (2013) 1300 ºC; Mann et al. (2012) 6 GPa, 2150 ºC; (Re) Righter and Drake (1997) 1150−1350 ºC; Ertel et al. (2001) 1400 ºC; Bennett and Brenan (2013) 2 GPa, 2000 ºC.

24

Brenan, Bennett & Zajacz

Neill 2007; see also Silicate and Oxide Control on HSE Fractionation). The lowest fO2 timeseries experiment of (Ertel et al. 2001), however, does suggest a minor contribution from a more reduced species. In isolation, high P–T data acquired between IW + 2.5 and IW − 1.5 suggest Re is dissolved in silicate melt as a 2+ species (Mann et al. 2012; Bennett and Brenan 2013). Correction of the data at 1400 ºC from Ertel et al. (2001) to 2000 ºC for comparison with the data of Bennett and Brenan (2013), however, reveals moderately good agreement between these datasets above ~ IW + 1 but evidence for a more reduced species in the higher T dataset at lower fO2 (Bennett and Brenan 2013). At present, there is insufficient data to thoroughly assess at what fO2 Re2+ becomes the dominant dissolved species, or to what extent this transition may depend upon P or T. Osmium: Experiments at 0.1 MPa and 1350−1400 ºC using Fe-free melts suggest Os is dissolved primarily as a 3+ species between ~ IW + 1 and IW + 4, although the presence of Os4+ cannot be excluded on the basis of these data (Borisov and Walker 2000; Fortenfant et al. 2006). Data from 0.1 MPa experiments performed at more reducing conditions display evidence for contamination by metal inclusions, preventing straightforward measurement of Os concentrations. Filtering of time-resolved LA-ICP-MS signals was found to be impossible due to the very low concentration of Os dissolved in the melt and the high Os background associated with available sulphide standards (Fortenfant et al. 2006). Brenan and McDonough (2009) however, were able to measure Os solubilities by LA-ICPMS in uncontaminated Fe-bearing experiments performed at 2 GPa 2000 ºC and as reduced as IW −1.6. These authors suggest solution as mixed 1+ and 2+ species based on the slope of 0.38 defined by their data. Although neither of these oxidation states is suggested by the 0.1 MPa data, it is worth noting that a mixture of 1+ and 3+ species also yields an adequate fit to the high T data. Importantly, the high T, low fO2 experiments suggest there must be some contribution from a species more reduced that Os2+ (expected slope of 0.5). Iridium: O’Neill et al. (1995) determined the solubility of Ir in CaO–MgO–Al2O3–SiO2 (CMAS) melt at 1400 ºC between IW − 1.5 and IW + 10. At conditions more oxidizing than ~ IW + 4, the change in Ir solubility is consistent with solution as a 2+ species (slope of ~ 0.5). At more reducing conditions however, O’Neill et al. (1995) observed no change in solubility with fO2 and ascribe this to either contamination by metal inclusions or solution as either Ir0 or Ir-carbonyl species. More recent solubility measurements at 1500 ºC on a similar composition over a comparable fO2 interval, reported by Fonseca et al. (2011), yielded results consistent with Ir3+ (slope of 0.75). A possible reason for the discrepancy between the two studies is that the metal contamination suggested in the experiments of O’Neill et al. (1995) are largely avoided in the work of Fonseca et al. (2011), as that study employed LA-ICPMS for sample analysis. Borisov and Palme (1995) determined the solubility of Ir-poor alloys (Ir10Pt90) over a similar range of fO2 and melt composition at 1300 ºC and 1480 ºC. Their data are consistent with solution as a 1+ species between ~ IW − 1 and IW + 8 and mixed 2+ and 3+ species at more oxidizing conditions (slope of 0.68). Results from 2 GPa experiments at 2000 ºC also suggest Ir dissolves predominantly as a 1+ species between IW + 0.5 and IW + 2.7 in basaltic melt. Mann et al. (2012) determined Ir partitioning at high P–T conditions (3.5−18 GPa, 2150−2500 ºC) and their data suggest Ir2+ as the dissolved species between ~ IW − 1.5 and IW + 0.5, although the authors acknowledge that this trend is not well defined. Spinel–melt partition coefficients measured at ~ IW + 1 to + 6 are also consistent with Ir2+ in the crystal lattice (Brenan et al. 2012). It is therefore likely that Ir2+ persists as a subordinate species at these more reducing conditions. Ruthenium: Borisov and Nachtweyh (1998) investigated Ru solubility at 0.1 MPa and 1400 ºC in anorthite-diopside melts between ~ IW + 6 and IW + 10. Their results define a slope of 0.73, corresponding to a 3+ oxidation state (Fig. 7). These results are consistent with the crystalmelt partitioning of Ru in olivine and spinel, that also suggest Ru3+ is present over much of the terrestrial fO2 range (Brenan et al. 2003, 2012). Laurenz et al. (2013) measured Ru solubility in picrite melts at 1300 ºC, 0.1 MPa and IW + 3.5 to IW + 5.5, obtaining results consistent with

Experimental Fractionation of the HSE

25

Ru dissolved as a 4+ species (Fig. 7). These authors suggest the difference in their data is due to the use of Fe-bearing melt compositions, that stabilize more oxidized species through redox exchange reactions analogous to those observed for Pd (Laurenz et al. 2010):

RuO3/ 2 (sil ) + FeO3/ 2 ( sil ) =RuO2( sil ) + FeO (sil )

(44)

In experiments done using Fe-bearing melt compositions at 6−18 GPa, 2150−2300 ºC and IW −1.5 to IW + 0.5 measured solubilities suggest Ru2+ is the dominant species (Mann et al. 2012). This is consistent with evidence from chromite-melt partitioning experiments, done at lower P-T conditions, that Ru2+ becomes dominant at conditions more reducing than ~ IW + 2.5 (Brenan et al. 2012). Rhodium: Ertel et al. (1999) determined the solubility of Rh in experiments performed at 0.1 MPa and 1300 ºC in melts with a composition close to the anorthite-diopside eutectic (Fig. 7). Between ~ IW + 6 and IW + 11, their data suggest Rh is dissolved primarily in the 2+ oxidation state, with a minor contribution from either a 3+ or 4+ species. Ertel et al. (1999) consider Rh3+ more plausible however, as the 3+ solid oxide is the stable phase at high temperature in air (Nell and O’Neill 1997). At more reducing conditions, measurements of Rh solubility at 0.1 MPa display evidence for contamination by stable metal inclusions. Regression of high P–T (6−18 GPa, 2150−2300 ºC) experiments at IW − 1.5 to IW + 0.5 yields a slope of 0.36, suggesting a mixture of Rh2+ and a more reduced species (Mann et al. 2012). Experiments to determine the olivine-melt (0.1 MPa, ~ 1330 ºC) and chromite-melt (0.1 MPa−2 GPa, 1400−1900 ºC) partitioning of Rh (see Spinel–melt partitioning of HSEs) are also consistent with the prevalence of a 2+ species between ~ IW + 5 and IW + 7 (Brenan et al. 2003, 2012). Platinum: At relatively oxidizing conditions (above ~ IW + 5), isothermal suites of experiments at 0.1 MPa define slopes of ~ 0.5 between 1300 and 1560 ºC, suggesting Pt is dissolved as a 2+ species (Borisov and Palme 1997; Ertel et al. 1999). Spectroscopic determination of Pt oxidation state in CAS glasses synthesized at 1630 ºC in air indicate the presence of 4+ species at highly oxidized conditions (Farges et al. 1999). At more reducing conditions (~ IW −0.8 to IW + 3.4), in experiments which were demonstrably free of contamination by metal inclusions, Bennett et al. (2014) observed a slight increase in Pt solubility with decreasing fO2 in basaltic glasses at 2 GPa and 2000 ºC. They interpret these results as owing to the presence of both neutral Pt and an unidentified Pt cationic complex. The solution of Pt-carbonyl species at low fO2 is discounted however, as no correlation is observed between Pt solubility and the carbon content of the melt (Bennett et al. 2014). A similar increase in solubility with decreasing fO2 was found by Médard et al. (2015) at more reducing conditions (~ IW − 2.5 to IW − 0.5) in experiments at 1.2 GPa and 1900 ºC. At lower T (1400 ºC and 1600 ºC) but similar fO2, however, Médard et al. (2015) measure constant Pt solubility with changing fO2. They interpret data with approximately constant solubility as suggesting Pt0 dissolved in the melt. Results from both Médard et al. (2015) and Bennett et al. (2014), that define a trend of increasing solubility as conditions become more reducing, are more consistent with the presence of either anionic platinum or the formation of complexes such as PtSix. Although solubilities are broadly consistent with past work at similar conditions, Mann et al. (2012) observe a slight decrease in Pt solubility with decreasing fO2 between IW + 0.5 and IW −1.5 at conditions of 3.5−18 GPa and 2150−2500 ºC. Isothermal, isobaric sets of experiments from that study define an average slope of 0.29, consistent with Pt1+ dissolved in the melt. It is unclear why these experiments yield a different result to the studies of Bennett et al. (2014) and Médard et al. (2015), as all were conducted using Fe-bearing melt compositions, and pressure appears to have little effect on Pt solubility (see Temperature and pressure). It is also unlikely that the higher temperatures employed in the Mann et al. (2012) study are

26

Brenan, Bennett & Zajacz

responsible, as measurements of Au and Pt solubility both suggest that increased T favours the reduced species (Borisov and Palme 1996; Médard et al. 2015). Cottrell and Walker (2006) also investigated Pt partitioning over a similar range of P–T conditions to Bennett et al. (2014) and Médard et al. (2015), but measured Pt solubilities that are systematically higher. Between IW − 0.75 and IW − 5.32, these authors observe no systematic dependence of solubility on fO2. These experiments may represent a continued increase in solubility with decreasing fO2, as suggested by extrapolation of the trend defined by 2000 ºC data from Bennett et al. (2014). In summary, Pt is dissolved in silicate melt predominantly as 2+ species above ~ IW + 5, with evidence for 4+ species at the most oxidizing conditions. At more reducing conditions the data is consistent with solution as Pt0, with either anionic species or non-oxide complexes becoming significant at the lowest oxygen fugacities investigated. Palladium: Borisov et al. (1994) measured Pd solubility in experiments at 0.1 MPa and 1350 ºC over a wide range of fO2 (~ IW to IW + 10) in anorthite-diopside melts (Fig. 7). These authors used a 3-species model to fit the data, comprising Pd2+, Pd1+ and Pd0. The data were found to be consistent with Pd1+ and Pd0 as the dominant species over the fO2 range investigated, with Pd2+ contributing most significantly to the dissolved Pd contents in experiments more oxidized than ~ IW + 7. Further experiments at 0.1 MPa using anorthite-diopside melts extended the database of Pd solubility measurements to lower fO2 (~ IW − 1) and temperatures of 1300, 1400, and 1480 ºC (Borisov and Palme 1996). These data are consistent with the oxidation states for Pd indicated by the earlier study of Borisov et al. (1994). Experimentallydetermined olivine-melt partition coefficients at ~ IW + 7.8 and ~ IW + 3.0 (0.1 MPa, 1335 ºC) also indicate a change from Pd2+ to Pd1+ as conditions become more reducing, in agreement with the solubility measurements (Brenan et al. 2003; Origin of the variation in partitioning). In experiments employing Fe-bearing melts however, Laurenz et al. (2010) observe lower solubilities below ~ IW + 3.5 than either Borisov et al. (1994) or Borisov and Palme (1996), and a dependence of solubility on fO2 that suggests Pd1+ not Pd0 is the predominant dissolved species. Between ~ IW + 3.5 and + 5.5 the data of Laurenz et al. (2010) again suggest Pd is dissolved as a more oxidized species than Borisov et al. 1994 (Pd2+ vs. Pd1+), and report higher solubilities for this fO2 range accordingly. These authors attribute the higher oxidation state of Pd indicated by their data between ~ IW + 3.5 and + 5.5 to redox exchange reactions with Fe, analogous to those described above for Ru. However, it is important to note that it can be difficult to assess speciation in Fe-bearing experiments which cover the high fO2 range in which both Fe2+ and Fe3+ are present, as the melt structural role of Fe3+ is different than Fe2+ (Farges et al. 2004). Hence, changes in solubility may also derive from changes in the activity coefficient for a single HSE species in response to melt structure, instead of a change in speciation (see Role of silicate melt composition on melt structure effects). Figure 7 displays the high pressure data of Righter et al. (2008) and Mann et al. (2012) collected at 1.5 GPa and 6−18 GPa respectively, corrected to unit activity of Pd. For data from the Mann et al. (2012) study, activity coefficients for Pd in the metal were taken as the values calculated in their activity model. Correction of the data from Righter et al. (2008) was made using the Margules parameters summarized by Borisov and Palme (2000) for the Pd-Fe system and ignoring the presence of Sb (< 7 wt%) in the alloy. Although only a limited number of isothermal high P data are available, those of Mann et al. (2012) (2150−2200 ºC) appear to be broadly consistent with the speciation model employed by Borisov et al. (1994). The lower T data of Righter et al. (2008) (1800 ºC) cover an insufficient range in fO2 for speciation to be reliably determined from only these datapoints. However, the solubilities recorded by their experiments are similar to those found by Mann et al. (2012) at similar fO2 but higher T. For temperatures > 1800 ºC, the experiments of Righter et al. (2008) suggest much greater Pd solubilities than measured by Mann et al. (2012). Although the reason for this discrepancy is not entirely clear, it is possible that the

Experimental Fractionation of the HSE

27

presence of Sb in the alloys used by Righter et al. (2008) affects the partitioning behavior of Pd, and subsequently, the solubilities at Pd saturation we calculate from their data. Gold: At 0.1 MPa and 1300−1480 ºC, experiments performed between ~ IW + 10 and IW using anorthite–diopside melt suggest Au is dissolved as 1+ species (average slope of ~ 0.23; Borisov and Palme 1996). Results also showed that at fO2 below ~ IW, however, the positive relationship between fO2 and solubility expected for the solution of metal-oxide species is reversed and solubility instead increases as conditions become more reducing. Borisov and Palme (1996) proposed that dissolution by way of an oxygen-producing reaction, such as the formation of silicide or carbide species, was most likely responsible for the observed increase in solubility at reducing conditions. Au solubility has also been determined in basaltic melts at 2 GPa and 2000 ºC (Brenan and McDonough 2009; Bennett and Brenan 2013). These studies reveal a weak negative dependence of solubility on fO2 between ~ IW − 1.2 and IW + 2.6, suggesting Au is dissolved in the melt as Au0 with the possibility of an additional contribution from either a silicide or other cationic species. The formation of carbide or carbonyl complexes in the studies of Brenan and McDonough (2009) and Bennett and Brenan (2013) is considered unlikely, as no difference in solubility is observed between experiments performed in graphite vs. metal-alloy capsules. From the regression of partitioning experiments done at 3−23 GPa and 1750−2500 ºC, Danielson et al. (2005) found DMet/Sil decreases as conditions become more reducing, consistent with the results of solubility experiments. The fact that Au1+ is not suggested by the high P–T data is also consistent with the experiments of Borisov and Palme (1996), which indicate the transition to reduced species occurs at a higher relative fO2 at elevated temperatures. Results of gold solubility measurements in hydrous, S- and Cl-bearing compositions at low-P and -T are discussed in Concluding Remarks. Role of silicate melt composition. Several studies have sought to determine the effect of silicate melt composition on HSE solubility. Borisov and Danyushevsky (2011) conducted a systematic investigation of Pt, Rh, and Pd solubility in air at 0.1 MPa and 1450−1550 ºC. Experiments performed in the CMAS system, with variable quantities of SiO2 added to an anorthite–diopside eutectic composition, reveal markedly different behavior for Pd compared with Pt or Rh. Pd solubility increases by ~ 55 ppm with increasing silica content between ~ 50 and 55 wt% SiO2. The addition of further silica to the system then causes a gradual decrease in solubility by ~ 75 ppm between ~ 55 and 70 wt% SiO2. Pt and Rh, however, display a monotonic decline in solubility as the SiO2 content increases from ~ 50 to 70 wt% SiO2. The change in both Pt and Rh solubility over this interval is linear and more pronounced for Rh than Pt. Wheeler et al. (2011) observed a decrease in the concentration of Pd in the silicate portion of their metal–silicate partitioning experiments with increasing bulk Pd contents. They rationalized this result, and its relationship to earlier studies, by positing the existence of a curved silicate saturation surface; which is intersected by tie-lines between co-existing liquids at progressively lower Pd concentrations as the metal alloy becomes more Pd-rich. Analogous behavior, in which a curved silicate saturation surface exists in the Pd–Si–silicate melt system, may explain the variation in Pd concentration observed by Borisov and Danyushevsky (2011) in the CMAS system. Borisov and Danyushevsky (2011) also performed experiments for Pt and Rh in the CA ± S system, and again observed a decrease in solubility for these elements with increasing silica content. Unlike the CMAS system, however, these experiments define a non-linear relationship between solubility and SiO2 content. After correcting their solubility data for CMAS melts to 1550 ºC for comparison with results from the CA ± S system, Borisov and Danyushevsky (2011) proposed the following relationships for Pt and Rh solubility as a function of melt composition: 2 log CPt = −4.59 XSiO2 + 2.09 XSiO + 1.35 X Al2 O3 + 2.92 2

(45)

2 log CRh = −3.23 XSiO 2 + 1.65 XSiO − 1.03 X Al 2 O 3 + 2.91 2

(46)

Brenan, Bennett & Zajacz

28 3.0

Pt

0.1 MPa; 1550 oC

log CHSE

2.5

3.0

Rh

2.5

2.0 2.0 1.5 1.5

1.0 0.0

0.2

0.4

XSiO

0.6

0.0

0.2

2

DAS Melts (XCaO+XMgO) XAl2O3

0.4

XSiO

0.6

2

CAS Melts 0.5 4.36 1.54 50

Figure 8. Solubility as a function of composition for CMAS melts at 0.1 MPa and 1550 ºC, as determined by Borisov and Danyushevsky (2011). Black triangles denote melts of anorthite–diopside eutectic + silica compositions, white triangles are for calcium–aluminium–silica melts (molar (CaO+MgO)/Al2O3 = 1.54 and 4.36 respectively).

The family of solubility vs. SiO2 curves produced by Equations (45) and (46) for different (CaO + MgO)/Al2O3 ratios are portrayed in Figure 8, with the accompanying experimental data. These results are qualitatively consistent with those of Dable et al. (2001), also obtained in the CAS system (0.1 MPa, 1227 ºC), who observe lower Pt solubility in melts with 70 wt% SiO2 than those with 40 wt% SiO2 for experiments done in air. In the experiments of Dable et al. (2001), however, this difference becomes less pronounced at low fO2, suggesting the compositional effect may be smaller for reduced Pt species in the melt. A negligible compositional dependence for Pt solubility at low fO2 is also supported by the agreement between high pressure experiments using basaltic to komatiitic melts (Mann et al. 2012; Bennett et al. 2014; Médard 2015). Nakamura and Sano (1997) conducted Pt solubility experiments at 0.1 MPa and 1600 ºC in air, using a variety of binary oxide melt compositions (BaO–Al2O3, BaO–SiO2, CaO–Al2O3, CaO–SiO2, Na2O–SiO2). These authors cast their solubility measurements as a function of theoretical optical basicity, a measure of the electron donation capacity of the melt components. Results show a linear increase in the logarithmic Pt concentration with optical basicity, where the slope of this trend is identical for each of the studied oxide pairs. The absolute Pt solubility however, is ~ 100 times greater for melts containing BaO as the basic oxide component. The results of Nakamura and Sano (1997) indicate a similar relationship between melt composition and solubility to those of Borisov and Palme (1997) and Borisov and Danyushevsky (2011), but also suggest that the identity of the acidic melt component (i.e., Al2O3 or SiO2) is unimportant relative to that of the basic component. Bond valence modeling of Pt in CAS melts suggests Pt is bonded to non-bridging oxygens surrounded by Ca second neighbors (Farges et al. 1999). This is also consistent with the observation that Pt solubility is enhanced in depolymerized melt compositions and tracks positively with CaO content. The study of Farges et al. (1999), however, also identifies Pt4+ not Pt2+ as the dissolved species in melts synthesized at 1630 ºC in air, contrary to that suggested by the dependence of solubility on fO2, albeit at lower T. Their model may, therefore, not apply directly to Pt in melts at low fO2. Most geologically important melt compositions contain iron, but most 0.1 MPa studies of HSE solubility have been performed using Fe-free melts. More recently, several studies have highlighted the role of Fe in both the solubility and speciation of HSEs in silicate melts (Laurenz et al. 2010, 2013). When compared with the solubilities of Pd

Experimental Fractionation of the HSE

29

log CHSE (ppm)

101

100

10-1

10

Pt Pt -2

0.0

0.2

0.4

Rh Rh

- Turchiaro 2013 - Borisov & Danyushevsky 2011

0.6 0.8 NBO / T

1.0

1.2

1.4

Figure 9. Solubility as a function of the melt structure parameter NBO/T (see the section Role of silicate melt composition for definition) for Pt and Rh at 0.1 MPa, 1400 ºC. Solid lines represent a polynomial fit to the Fe-bearing experiments of Turchairo (2013). Dashed lines represent the same fit, translated to best match the Fe-free experiments of Borisov and Danyushevsky (2011). Both studies display a similar dependence on melt polymerization, however, solubilities in the Fe-bearing composition are systematically higher.

and Ru measured in Fe-free melts, these authors observed the stabilization of more oxidized dissolved species at a given relative fO2. They suggest this difference results from redox exchange equilibria with Fe as described, for example, by Equation (44). Whether solubility in Fe-bearing melts is higher or lower than seen for Fe-free melts will therefore depend upon the fO2 conditions being investigated. Turchiaro (2013) investigated Pt and Rh solubility in basalt-rhyolite mixtures at 0.1 MPa, 1400 ºC and ~ IW + 6.8. The solubility of both elements was seen to decrease with increasing proportions of the rhyolite component. When plotted as a function of non-bridging oxygens over tetrahedrally coordinated cations (NBO/T), a measure of melt polymerization, the variation in solubility measured by Turchiaro (2013) is comparable to that measured by Borisov and Danyushevsky (2011) in CMAS melts. After correction of their data to account for differences in T and fO2 however, the absolute solubilities measured by Turchiaro (2013) are systematically higher than those of Borisov and Danyushevsky (2011) (Fig. 9). This is consistent with the idea that Fe may enhance HSE solubility through exchange equilibria of the type suggested by Laurenz et al. (2010). It should also be noted however, that the experiments of Turchiaro (2013) used natural starting materials containing several weight percent of alkali elements (Na2O, K2O) and TiO2 that may also modify HSE solubility (Borisov et al. 2004, 2006). Borisov et al. (2006) investigated the effect of sodium content on the solubility of Pd by adding various quantities of Na2O to melts with an An–Di eutectic composition. In experiments done at 0.1 MPa, 1300 ºC and in either air or a CO2 atmosphere, Pd solubility was seen to decrease by 20−30% with increasing sodium content, up to ~ 4 wt% Na2O. Further addition of Na2O however, elicits no change in Pd solubility up to ~ 11 wt% Na2O, the most sodium rich composition investigated. The inverse behavior is observed for the addition of titanium oxide to CMAS melts, where little or no increase in Pd solubility is observed up to ~ 4 wt% TiO2, after which the log Pd solubility increases in a linear fashion up to ~ 25 wt% TiO2 (Borisov et al. 2004). A similar increase in solubility with TiO2 content is observed for Ni and Fe (Borisov et al. 2004). Furthermore, the dependencies of both Ni and Co solubility on melt SiO2 content are similar to that observed for Pd, but not Pt or Rh (Borisov and Danyushevsky 2011). These features suggest a similar structural environment might be shared by Pd, Ni, Fe, and Co, which is distinct from certain other HSEs such as Pt and Rh.

30

Brenan, Bennett & Zajacz

Sulfur may also affect HSE solubility by providing an additional ligand for the formation of dissolved species. In experiments saturated in an HSE metal phase, but undersaturated with respect to an immiscible sulphide phase, solubility may be enhanced through reactions with the form (Laurenz et al. 2013): x x (47) MO x / 2 ( met ) + S2 − ( g )= MSx / 2 ( sil ) + O2 − ( g ) 2 2 Experiments for Ru at 1300 ºC, 0.1 MPa and log fS2 of −2.3 display an order of magnitude increase in solubility compared with sulfur-free experiments at otherwise identical conditions (Laurenz et al. 2013). It is possible that the addition of other complexing anions (e.g., Cl, P) may also enhance HSE solubility through reactions analogous to Reaction (47); however, Blaine et al. (2011) observed no increase in Pt solubility in Cl-bearing experiments.

Temperature and Pressure. It has been recognized for some time that the solubility of many siderophile elements in silicate melt increases with temperature (e.g., Murthy 1991; Capobianco et al. 1993; Walker et al. 1993; Hillgren et al. 1994). Recent studies have expanded the database for HSEs significantly and confirmed the presence of a T dependence for elements which previously had little or inconclusive data (e.g., Righter et al. 2008; Brenan and McDonough 2009; Mann et al. 2012; Bennett and Brenan 2013; Bennett et al. 2014; Médard et al. 2015). Measurements of HSE solubility are most often used to estimate metal–silicate partition coefficients, which can be calculated from Equation (36). This calculation however, depends not only on the HSE concentration in the melt at the solubility limit, but also on the activity coefficient for that HSE at infinite dilution in a liquid Fe solvent, gM. Values of gM may themselves vary with temperature and pressure, meaning an observed dependence of HSE solubility on these variables may not directly translate to similar behavior during metal–silicate partitioning. Where suitable thermodynamic data is available, gM may be calculated at the appropriate P–T conditions with a binary asymmetric mixing model (Thompson 1967): 2 WM + 2 X M ( WFe − WM )  RT = ln γ M X Fe

(48)

where WM and WFe are Margules interaction parameters for the HSE of interest (M) and Fe respectively, that are calculated at the required P and T from individual components relating to the excess enthalpy ( WH ) , entropy ( WS ) and volume ( WV ) of mixing:

(

)

(49)

W = WH − TWS + P − P  WV

where Pº is the reference pressure at which WH, WS, and WV were determined; typically 0.1 MPa (1 bar). The T dependence of W for most Fe–HSE binaries is relatively small, and Equation (49) provides a suitable value for W. For systems such as Fe–Pd and Fe–Pt, however, which have a greater dependence on temperature, it may be required to extrapolate values of W from the reference T at which the interaction parameters were acquired (Tº), to the conditions of the experiment (Mann et al. 2012). This can be accomplished using the following relationship (Japan Society for the Promotion of Science 1988): W=

(W

H

(



) ) TT

− T WS + P − P  WV

(50)

A summary of the available thermodynamic data for Fe–HSE and HSE–HSE systems, and more detailed discussion of modeling the activity-composition relationships in multicomponent HSE–Fe alloys can be found in Mann et al. (2012). With these considerations in mind, Figure 10 displays the variation in DMet/Sil for the HSEs as a function of inverse temperature. All data show a decrease in DMet/Sil with increasing T, in

Experimental Fractionation of the HSE 2500 oC 2000 oC 8.0

1500 oC

Au

8.0

11.0

6.0

9.0

4.0

4.0

7.0

2.0

2.0

5.0

0.0

13.0

5.5

Ir

6.5

3.5

11.0

4.5

5.5

Rh

6.5

3.5

11.0

9.0

9.0

9.0

7.0

7.0

7.0

5.0

5.0

3.0

5.0 3.5

10.0

Ru

4.5

5.5

Recalculated to ~IW

6.5

4.5

5.5

6.5

4.5

5.5

6.5

Os

3.0

3.5

10.0

Pt

3.0

0.0 4.5

11.0 log DMet/Sil

Pd

6.0

3.5

31

4.5

5.5

6.5

3.5 0.1 MPa

Re

2 GPa

Bennett & Brenan 2013: Au, Re

Bennett et al. 2014: Pt

Borisov & Palme 1996: Au, Pd

Borisov & Palme 1997: Pt, Rh

Borisov & Nactweyh 1998: Ru

8.0

8.0

Borisov & Walker 2000: Os

6.0

6.0

Ertel et al. 1999: Pt, Rh

4.0

4.0

2.0

2.0

Brenan & McDonough 2009: Os, Ir, Au ≥21 GPa

4.5

5.5

6.5 3.5 104 / T (K)

≤13 GPa

Danielson et al. 2005 Ertel et al. 2001: Re

Ertel et al. 2006: Pt

Fortenfant et al. 2003: Pt, Rh

Fortenfant et al. 2006: Os

Laurenz et al. 2010: Pd

Laurenz et al. 2013: Ru 3.5 GPa

3.5

Borisov et al. 1994: Pd

Borisov & Palme 1995: Ir

6 GPa

18 GPa

Medard et al. 2015: Pt

4.5

5.5

6.5

1.5 GPa

>1.5 GPa

Mann et al. 2012: Pd, Pt, Rh, Ir, Os, Ru O’Neill et al. 1995: Ir Righter et al. 2008: Pd Yokoyama et al. 2009: Os

Figure 10. Metal–silicate partition coefficients for the HSE as a function of inverse temperature. The solid lines in each figure represent fits to the following data: (Au) 2 GPa, IW + 1.7 to + 2.0 Brenan and McDonough (2009), Bennett and Brenan (2013); (Pd) 0.1 MPa, IW − 0.7 to + 1.0 (excepting a single, anomalously low datapoint) Borisov et al. (1994), Borisov and Palme (1996); (Pt) 2 GPa, IW to IW + 0.2 Bennett et al. (2014); (Ir) 2 GPa, IW + 1.8 to + 2.1 Brenan and McDonough (2009); (Rh) 0.1 MPa, ~ IW + 6 to IW + 8 Fortenfant et al. (2003); (Os) 2 GPa, IW + 1.7 to + 2.3 Brenan and McDonough (2009); (Ru) 0.1 MPa, IW + 7.4 to + 7.8; (Re) 2 GPa, IW + 1.7 to + 2.0 Bennett and Brenan (2013). Dashed line in the panel for Pd represents the low P fit adjusted to best match the data of Righter et al. (2008). For Rh, the dashed line represents the same fit as the solid line, adjusted to match the low-fO2 experiment of Ertel et al. (1999). The dashed triangles for Ru are high-fO2 measurements re-calculated to correspond to IW as described in the text. The dashed line in this panel is the fit to the 0.1 MPa data adjusted to match the high-P data of Mann et al. (2012). Recalculated 0.1 MPa data agree well with the fit to the high-pressure data, suggesting a negligible role for pressure. For Re and Ir, only data from experiments more oxidizing than IW are plotted for the Mann et al. (2012) study, for straightforward comparison with the other high-P data.

the order Pt > Os ≈ Ir > Pd > Ru ≈ Au > Rh > Re. The solid lines in each panel of Figure 10 represent fits to an isobaric dataset acquired at the conditions noted in the caption. For most elements the dependence of DMet/Sil on T appears independent of P and fO2. At low T however (< 1400 ºC), some data for Pt fall systematically below the trend defined by data at higher T and pressures ranging from 0.1 MPa to 18 GPa. For the data of Ertel et al. (1999) and Fortenfant et al. (2003), this is due to the higher fO2 of their experiments. It should also be noted that the weaker dependence of DMet/Sil on T indicated by the data of Fortenfant et al. (2003) is from

32

Brenan, Bennett & Zajacz

experiments performed at the same absolute fO2 but different relative fO2, unlike data used to define the solid line in Figure 10 which represent experiments at a similar relative fO2. For Pt, Re, Ru, Au, and to a lesser extent Pd, results from experiments done at 0.1 MPa to 18 GPa are generally well reproduced by a single linear regression, indicating only a weak or negligible effect of pressure on the partition coefficient. There are, however, several noteworthy exceptions. For Au, data collected at ≥ 21 GPa fall below the trend defined by results at 0.1 MPa to 13 GPa, suggesting pressure may cause a decrease in DMet/Sil at these conditions. Data for Pd also suggest little effect of P on DMet/Sil between 0.1 MPa and 3.5 GPa. Results from Mann et al. (2012) done at > 6 GPa however, fall systematically below the trend defined by 0.1 MPa data. Experiments from the study of Righter et al. (2008) done at 1.5 GPa to 15 GPa also suggest lower values of DMet/Sil for Pd than expected from the 0.1 MPa trend (Fig. 10), although this may be due to the presence of Sb ± Re ± S ± P ± C in the alloy phase used in those experiments. The effect of P on DMet/Sil for Os is difficult to assess due to the use of Os–Ni alloys as the HSE source in experiments at 0.1 MPa. Activity–composition relationships in this binary system are unknown and the measured concentrations of Os in the silicate portion of experiments performed by Borisov and Walker (2000) and Fortenfant et al. (2006) have been corrected to those expected for the pure metal by assuming ideality in the Os–Ni alloy. This leads to rather high corrected Os solubilities, accompanied by low values for DMet/Sil. The presence of a miscibility gap in the Ni–Os system (Okamoto 2009), however, indicates the assumption of ideality for these alloys is an oversimplification and calculated values of DMet/Sil should therefore be considered minimumvalues. The experiments of Yokoyama et al. (2009), done at 1−2 GPa yield values of DMet/Sil that lie between those of Brenan and McDonough (2009) done at 2 GPa, and the ambient pressure studies, but provide little extra constraint on the pressure effect. There are no data for Ru partitioning at 0.1 MPa and conditions more reducing than ~ IW + 5.5, preventing direct comparison to the available high-P results of Mann et al. (2012) obtained at more reducing conditions (~ IW −1 to IW + 0.5). Correction of the solubilities obtained from 0.1 MPa experiments at oxidizing conditions to the average relative fO2 of high-P experiments, however, reveals good agreement between data at different pressures (0.1 MPa to 18 GPa). Correction of the 0.1 MPa data was done assuming either a 3+ (Borisov and Nachtweyh 1998) or 4+ (Laurenz et al. 2013) oxidation state for Ru dissolved in the melt, as found in the respective studies. Metal–silicate partition coefficients for Au, Ir, Pd, and Rh, suggest a reduction in values with increasing P, for at least part of the investigated pressure range. For the case of Au, the effect of P is difficult to isolate, as the partitioning experiments of Danielson et al. (2005) that suggest a reduction in DMet/Sil at high P are also sulfur-bearing. Values of DMet/Sil for Au are expected to decrease with increasing S contents of the metal phase (e.g., Jones and Malvin 1990) and S itself becomes more siderophile with increasing P (Boujibar et al. 2014). Runproduct compositions are not quoted by Danielson et al. (2005), making it difficult to quantify if the low values of DMet/Sil recorded by experiments at 21 and 23 GPa are the direct result of P or the coupled effects of P and composition. Values of DMet/Sil for Rh at high P appear to lie along the trend defined by 0.1 MPa experiments (Fig. 10), however, there is a significant difference (> 4 log units) in relative fO2 between these data. Plotting the same trend through an experiment done at 0.1 MPa, 1300 ºC and more reducing conditions (dashed line in Fig. 10), reveals the discrepancy between high and low pressure experiments. Fig. 11 displays DMet/Sil for Rh as a function of pressure for the experiments of Mann et al. (2012) (~ IW − 1.5 to IW + 1.5, 2150–2500 ºC) and the 0.1 MPa experiments shown in Figure 10, corrected to 2180 ºC and IW using the T dependence of Fortenfant et al. (2003) and assuming Rh2+ in the melt. Figure 11 portrays the negative dependence of DMet/Sil on P, which likely changes magnitude somewhere between 6 and 18 GPa (Mann et al. 2012). A similar comparison of DMet/Sil vs. P can be made for Ir, after correction of data at similar fO2 (~ IW + 0.5 to IW + 0.7) to the same T, using the trend found by Brenan and McDonough (2009). Although there is a dependence of DMet/Sil on P for

Experimental Fractionation of the HSE 11.0

6.0

Ir

5.5

9.0

5.0

8.0

4.5

7.0

4.0

6.0

3.5

log DMet/Sil

10.0

5.0

0

10.0

4

8 12 P (GPa)

16

20

3.0

33

Pd

0

4

8 12 P (GPa)

16

20

Rh

log DMet/Sil

9.0 Borisov & Palme 1994 Brenan & McDonough 2009 Fortenfant et al. 2003 Borisov & Palme 1995 Ertel et al. 1999 Mann et al. 2012

8.0 7.0 6.0 5.0 4.0

0

4

8

12 P (GPa)

16

20

Figure 11. Metal-silicate partitioning for Ir, Pd and Rh as a function of pressure. All three elements display a significant decrease in DMet/Sil with increasing P. For Ir, all data are at ~ IW + 0.5 and the data of Mann et al. 2012 and Borisov and Palme (1995) have been corrected to 2000 ºC using the T-dependence found by Brenan and McDonough (2009). For Pd, we have plotted experiments done at conditions more reducing than ~ IW − 0.2 from Borisov and Palme (1996) and the high P data of Mann et al. (2012) (~ IW − 1 to IW + 0.6). Both datasets were corrected to 2000 ºC using the dependence shown in Figure 10. Data for Rh are from Ertel et al. (1999), Fortenfant et al. (2003) and Mann et al. (2012). Data from the 0.1 MPa studies was corrected to IW assuming Rh2+ in the melt. All data were corrected to 2000 ºC using the T-dependence found by Fortenfant et al. (2003).

Ir, its magnitude rests heavily on the datapoint at 6 GPa, which defines a minimum in the data. Further experiments are therefore required to better quantify the effects of P on DMet/Sil for Ir. Figure 11 shows DMet/Sil vs. P for Pd, from experiments done at conditions more reducing than ~ IW − 0.2. As observed for Rh and Ir, the magnitude of the P dependence changes above 6 GPa, although the difference is weaker for Pd than indicated for either Rh or Ir (Mann et al. 2012). In summary, all of the HSEs display a decrease in DMet/Sil with increasing T. Compilation of the literature data also suggests that DMet/Sil for several elements decreases significantly with increasing P and the magnitude of this change may vary across the investigated P range (Mann et al. 2012).

Application of results to core formation In order to apply the results of solubility and partitioning experiments to models of terrestrial accretion and core formation, values are parameterized according to Equation (40). To demonstrate this approach, we have chosen to parameterize DMet/Sil for Pt, Re, and Os as these elements comprise the long-lived Re–Os and Pt–Os isotope systems. The data summarized in earlier sections suggest that for reducing conditions, T and fO2 are the important controls on DMet/Sil. Equation 40 may thus be simplified to yield: log DiMet/Sil =

a n + ∆I W + c T 4

(51)

Values for the coefficients a, n, and c and their sources are listed in Table 3. For equilibration between metal and silicate reservoirs at a chosen set of conditions, the concentration of trace element i in the silicate phase ( CiSil ) can then be calculated using the following relationship:

Brenan, Bennett & Zajacz

34

CiSil =

(

CiTOT f + (1 − f ) DiMet/Sil

(52)

)

where CiTOT is the concentration of i in the bulk system (typically chondritic concentrations when considering core formation) and f is the fraction of silicate melt being equilibrated. Figures 12a and b display the change in CiSil for a primitive upper mantle composition (PUM) following metal–silicate equilibrium over a range of T and fO2 conditions. It can be seen from these figures that the PUM concentrations of Re and Os are not reproduced by metal–silicate equilibrium over a wide range of T–fO2 space. Conversely, PUM concentrations of platinum can be accounted for if metal–silicate equilibrium occurs at high T and low fO2 (e.g., ~ 3250 ºC at IW − 2). Figures 12c and d display the change in Re/Os and Pt/Os ratio over the same T and fO2 interval as shown in 12a and b. Approximately chondritic Pt/Os, as required to account for the Os isotopic composition of PUM, is not reproduced under any conditions. The PUM Re/Os ratio is only reproduced at relatively low T ( 4000) and Rh (~ 80−300), and uniformly low values for Pd (i.e., < 1). These results have been confirmed for Cr-bearing spinel at similar conditions in the more recent work of Righter et al. (2004), who also report a Dmineral/melt for Ir of 5 to > 10,000, and that Au and Re are incompatible (Dmineral/melt of 0.08 and 0.0012−0.21, respectively). Results obtained by Brenan et al. (2012) for Cr-spinel at 0.1 MPa and 2 GPa, 1400–1900 ºC and more reduced fO2 (FMQ – 2 to FMQ + 4) yielded generally lower partition coefficients for the IPGE than previous work, with Ru as the most compatible (Dmineral/melt of ~ 4), followed by Rh and Ir, which are moderately incompatible to compatible (Dmineral/melt range of 0.04 to ~ 1), with Pt and Pd the most incompatible (Dmineral/melt < 0.2). Mallmann and O’Neill (2007) investigated the (MgAl) spinel–melt partitioning of Re at 1275–1450 ºC and pressure of 1.5–3.2 GPa with fO2 ranging from FMQ − 2.9 to + 5.6. Their results show a remarkably systematic decrease in the Dmineral/melt for Re over this fO2 interval, ranging from ~ 0.3 to < 3 × 10−5. A summary of partition coefficients measured in these studies is provided in Figure 14, with the logical abscissa being experiment fO2. Although there is significant scatter to the data, an overall consistent result is that Ir, Ru, and Rh are more compatible in spinel then Pt, Pd, Au, and Re. In contrast to the results for Re, partition coefficients for Ir, Ru, and Rh show an overall decrease as conditions become more reducing.

Experimental Fractionation of the HSE

37

A B alloy

Pd

new olivine melt

SCO 5 mm C

TC sheath 200 m

SCO D

50 m

Au alloy

Chr

melt

Ir alloy Figure 13. Experimental configuration and run products from experiments to measure mineral–silicate melt partitioning of the HSE. (a) Example of sample configuration for an olivine–silicate melt partitioning experiment done at 0.1 MPa in a gas-mixing furnace. The olivine crucible (fabricated from single crystals of San Carlos olivine) is filled with basalt plus metal powder, then hung using a wire (in this case Pd) from a hook made from high purity fused silica, within the hot zone of a vertical tube furnace. The thermocouple, enclosed in a high purity alumina sheath, is located next to the sample during the experiment. (b) Backscattered electron image of the sectioned olivine crucible from an experiment done using the configuration shown in (a), cooled from 1350 to 1338 ºC over 5 days and log fO2 = −3.1. The sample consists of quenched Fe-bearing silicate melt (= glass), olivine phenocrysts (= new olivine) and metal source, which were Ir metal grains encapsulated in molten Au (similar to the configuration employed for the metal/silicate partitioning experiment, see also Figure 6). Circular features correspond to areas ablated for ICP-MS analysis. [Used by permission of Elsevier Limited, from Brenan, McDonough and Ash (2005), Earth and Planetary Science Letters, Vol. 237, Fig. 1a, p. 861.] (c) and (d) Reflected light photomicrographs of the sectioned run product from a chromite-silicate melt partitioning experiment done at 1850 °C, 2 GPa, for 2.8 h. The field of view in (d) corresponds to the rectangle in (c). Chromite, which was added as a finely ground powder, with a grain size less than 40 μm, has dissolved and re-precipitated to form equant crystals. The melt has quenched to a fine intergrowth of silicate and oxide crystals plus glass. Note the preferential wetting of the Ir alloy by Au. [Used by permission of Elsevier Limited, from Brenan, Finnigan, McDonough and Homolova (2012), Chemical Geology, Vol. 302−303, Fig. 2, p. 24.]

Brenan, Bennett & Zajacz

38

(a)

105 Rh: MgAl O

104

2

range from Righter et al (2004)

4

magnetite Cr-spinel Re: MgAl O

3

10

2

102

4

Ir: Cr-spinel

101 spinel/ melt 100

4+

-1

10

10-2

Re model

10-3 -4

10

6+

10-5 10-6

-4

-2

0

2 FMQ

4

6

8

(b)

5

10

4

10

3

10

102 101 spinel/ melt 100 10

-1

10-2 10

-3

10-5 10-6

Ru: MgAl O

Pd: MgAl O

Cr-spinel

magnetite

magnetite

Cr-spinel

2

10-4

-4

-2

4

2

0

4

Pt: MgAl O 2

4

Cr-spinel

2 FMQ

4

6

8

Figure 14. Summary of spinel–melt partitioning of the HSE. Dashed curve is for the calculated variation in spinel-melt partitioning of rhenium after the model of Mallmann and O’Neill (2007), using Equation (60) in the text. Values of the end-member partition coefficients for Re4+ and Re6+ are fit to the partitioning data, whereas rhenium speciation is determined from the metal solubility measurements of Ertel et al (2001). Data are from the following sources: (Rh) Righter et al. (2004), Homolova (2008), Brenan et al. (2012), Capobianco and Drake (1990), Capobianco et al. (1994); (Pd) Homolova (2008), Brenan et al. (2012); (Re) Mallman and O’Neill (2007); (Ru) Homolova (2008), Brenan et al. (2012); (Ir) Righter et al. (2004), Homolova (2008), Brenan et al. (2012); (Pt) Homolova (2008), Brenan et al. (2012).

Silicate mineral–melt partitioning of HSEs Aside from spinel, most of the previous mineral-silicate melt partitioning measurements for the HSE have been measured for olivine, with a generally similar sense of fractionation: Ir, Ru, and Rh moderately compatible, and Pt, Pd, Au, and Re incompatible. Results are summarized as a function of fO2 in Figure 15. Olivine melt partition coefficients measured at 0.1 MPa and 1330−1343 ºC for Ru and Ir increase from values of ~ 0.5 at FMQ > + 4 to ~ 2 at lower fO2. Values for Rh partitioning are ~ 2 over the entire fO2 range considered thus far (FMQ + 2 to + 7; Brenan et al. 2003, 2005; unpublished data). Partition coefficients for Au and

Experimental Fractionation of the HSE

39 (a)

101 10

Rh

0

Ir 4+

-1

10

olivine/ melt 10-2 10-3 10-4

Re

-5

10

-4

-2

0

2

4

6+

6

FMQ

8

(b)

1

10

100 Ru 10

-1

Pd 2+

olivine/ melt 10-2

Pt 1+

10-3 10

-4

10

-5

-4

-2

0

2

4

6

8

FMQ

Figure 15. Summary of olivine–melt partitioning of the HSE. Dashed curves are for the calculated variation in olivine–melt partitioning of Re and Pd, after the model of Mallmann and O’Neill (2007), using Equation (60) in the text. Values of the end-member partition coefficients for Re4+ and Re6+ and rhenium speciation are determined as per the caption to Figure 14. The palladium partitioning model assumes speciation determined from the metal solubility measurements of Borisov et al. (1994), with end-member partition coefficients for Pd1+ and Pd2+ fit to the data. Data are from the following sources: (Ru) Brenan et al. (2003), black inverted triangles; Righter et al. (2004), gray inverted triangles; (Pd) Brenan et al. (2003) and unpublished, black circles; Righter et al. (2004), gray circles; (Re) Mallman and O’Neill (2007), black diamonds; Brenan et al. (2003), gray diamonds; Righter et al. (2004), open diamonds; (Rh, Pt) Brenan et al. (2003); (Ir) Brenan et al. (2005).

Pt measured at similar conditions are uniformly low (0.12 or less), with no apparent change with fO2 (Righter et al. 2004 ; Brenan et al. 2005). Values for Pd measured by Brenan et al. (2003; unpublished data) decrease systematically with fO2: from 0.05 at FMQ + 7 to ~ 0.006 at FMQ − 0.5. The single value of ~ 0.1 reported by Righter et al. (2004) under oxidizing conditions is somewhat higher, but broadly consistent with this trend. In contrast to Pd, olivine-melt partition coefficients for Re measured by Mallmann and O’Neill (2007) increase with decreasing fO2, from 1.5 × 10−5 at ~ FMQ + 6 to ~ 0.5 at FMQ − 3, showing a similar trend

40

Brenan, Bennett & Zajacz

to the results for spinel. Values for Re partitioning measured by Righter et al. (2004) are systematically larger than those measured by Mallmann and O’Neill (2007) at similar relative fO2, possibly reflecting some metal contamination in the run-product olivines. Re partition coefficients measured by Brenan et al. (2003) are significantly lower than the Mallmann and O’Neill (2007) determinations, which the latter authors attribute to a low abundance of charge balancing hydrogen-related point defects, present in their high pressure experiments, but absent in the 0.1 MPa, dry experiments of Brenan et al. (2003). Watson et al. (1987) did a reconnaissance study on clinopyroxene at 1275 ºC and 0.1 MPa, obtaining Dmineral/melt of ~ 0.04 for Re, and ~ 0.08 for Os, but fO2 was not controlled. Results of other silicate mineral–melt partitioning experiments are summarized in Table A2. Aside from the extensive data for Re, of note is the Dmineral/melt for Pt of 1.5 determined for clinopyroxene by Righter et al. (2004). This value is somewhat unexpected, however, given that Pt2+ is the likely species at the fO2 of those experiments (Ertel et al. 1999), with an ionic radius of 0.8 in VI-fold coordination (Shannon 1976), and therefore an expected partition coefficient of ~ 0.1 based on the Blundy–Wood partitioning model (see next section). Results for Re partitioning include the measurements of Mallmann and O’Neill (2007) for clinopyroxene, orthopyroxene, and garnet (in addition to olivine and spinel mentioned above) determined over a significant range in fO2 (FMQ: 2.9 to + 5.6) at 1.5 to 3.2 GPa. Re was found to be highly incompatible in these phases at the most oxidized conditions, but values increase with decreasing fO2. Values of Dmineral/melt for Re involving clinopyroxene and garnet approach or exceed unity at ~ FMQ − 2, whereas Re is incompatible in orthopyroxene at all conditions studied. Righter et al. (2004) and Righter and Hauri (1998) also found Re to be incompatible in clinopyroxene and orthopyroxene at similar fO2, and Righter and Hauri (1998) measured DRe > 1 for garnet at FMQ − 2.9 to − 4.8.

Origin of the variation in partitioning As described previously for solid metal–liquid metal partitioning (see The role of the solid phase on DSM/LM ), Equation (16) relates the mineral–melt partition coefficient to the degree of size misfit between the substituent cation, and the optimal value for the crystallographic site (Blundy and Wood 1994). This model accounts well for HSE partitioning into olivine and spinel, (Mallman and O’Neill 2007; Brenan et al. 2003, 2005), with the additional importance of crystallographic site occupancy, and its variation with composition, to fully account for the spinel partitioning data (Brenan et al. 2012). A further complication to the partitioning of the HSE is that the valence state for some changes over the fO2 at which partitioning has been investigated, thus influencing cation size misfit, and possible substitution mechanisms as well. A change in valence state with fO2 has been demonstrated by Mallmann and O’Neill (2007) to control Re partitioning into pyroxene, olivine and spinel, and suggested by Brenan et al. (2003, 2005) for Ru, Ir, and Pd substitution into olivine. In the latter case, the increase in D for Ru and Ir was interpreted to result in a shift from Ir3+ and Ru3+ in the melt to divalent species, as the ionic radius of Ru2+ and Ir2+ is estimated to be close to Mg2+, hence nearly optimal for substitution into the olivine structure, as predicted by the model of Blundy and Wood (2001). Estimates of valence state from Ir and Ru solubility experiments are somewhat conflicting, however, as it is possible to fit both 2+ and 3+ species to some data (Ru, Borisov and Nachtweyh 1998; Ir, Brenan et al. 2005), whereas other results suggest higher valence states (Ru4+, Laurenz et al. 2013; Ir3+, Fonseca et al. 2011). Partitioning experiments over a broader range of fO2 would be useful to better understand this behavior. Changes in element partitioning involving a shift in valence state can be understood by considering the case for Re partitioning, in which there is very good agreement between partitioning systematics and valence state changes implied by the solubility data. Following the approach of Mallman and O’Neill (2007), the solution of Re metal into silicate melt can be described by the reaction:

Experimental Fractionation of the HSE Remetal + x/4 O2 = Rex+Ox/2

41 (53)

which has a solubility product of the form: Qx+ = [Rex+ Ox/2] / fO2x/4

(54)

As demonstrated by Ertel et al. (2001), the variation in the solubility of Re with fO2 in the basalt-analogue they investigated can be modeled by contributions from both Re4+ and Re6+ species. Hence, the total solubility of Re in the melt at a given fO2 can be expressed as: [Σ ReOx/2] = [Re4+ O2] + [Re6+ O3]

(55)

with square brackets denoting the melt phase. Substituting the individual solubility products yields: [Σ ReOx/2]= Q4+ + Q6+ fO23/2

(56)

x+

Values of Q are calculated by fitting the experimental solubility data. An equation equivalent to Equation (55) can be written for the crystalline phase, yielding the following relation for the partition coefficient including all melt species: DRe = ({Re4+ O2} + {Re6+ O3})/ ([Re4+ O2] + [Re6+ O3])

(57)

in which curly brackets denote the crystalline phase. Partition coefficients for the individual melt species can be written as: DRex+ = {Rex+ Ox/2} / [Rex+ Ox/2]

(58)

Equation (58) can be used to eliminate the crystalline phase terms from Equation (57), yielding the relation: DRe = ([Re4+O2]D4+) + ([Re6+O3]D6+) / ([Re4+O2] + [Re6+O3])

(59)

x+

and expressing [Re Ox/2] in terms of the solubility product yields: DRe = (Q4+ fO2 D4+) + (Q6+ fO23/2 D6+) / ((Q4+ fO2) + (Q6+ fO23/2))

(60)

Hence, a full description of the partitioning in systems involving multiple valence states can be obtained with knowledge of the partition coefficients for the pure valence state, as well as the solubility product for the metal dissolution reactions. Mallman and O’Neill (2007) applied this analysis to their Re partitioning data, using values of Qx+ obtained from the Re solubility experiments of Ertel et al. (2001), and estimates for Dx+ that provided a best fit to their data. Results of fitting the Re partitioning data for olivine and spinel to Equation (60) are shown in Figures 14 and 15, with similar behavior for the other minerals studied. Re6+ is highly incompatible in olivine and spinel, owing to charge and size mismatch, with Re4+ much less so, behaving similar to Ti4+ (Mallmann and O’Neill 2007), accounting for the significant variation in partitioning with fO2. Following a similar approach, the olivine-melt partitioning for Pd (Brenan et al. 2003; unpublished data) can also be well fit to a form of Equation (60), but assuming Pd2+ and Pd1+ (Fig. 15), consistent with the metal solubility data of Borisov et al. (1994) and Laurenz et al. (2010). By this analysis, the partition coefficients are expected to change markedly for the fO2 interval within which both species are abundant, then level off to constant values at more reduced, or oxidized conditions, in which a single species dominates. Given the evidence for changes in valence state of the HSEs described in Solid Metal–Liquid Metal Partitioning, it seems clear that characterizing partition coefficients over a range of fO2 is essential to capture the behavior likely for natural magmas. In addition to the role of ionic radius and charge, changes in the crystallographic site occupancy as a function of composition is an additional factor that may influence the HSE

Brenan, Bennett & Zajacz

42

incorporation into spinels. The structural formula for spinel can be written as (A1-xBx)[AxB2-x]O4 in which A and B are di- and trivalent cations (so-called 2,3 spinel) or di- and tetravalent cations (so-called 2,4 spinel) in IV-fold (curved brackets) or VI-fold (square brackets) coordination, and x is the inversion parameter. The inversion parameter ranges from 0, corresponding to fully “normal” spinel, to 1, or fully “inverted” spinel, in which half of the octahedral sites are occupied by divalent cations. Hence, the availability of sites for di- and trivalent cations stabilized by VIfold coordination will vary according to the degree of inversion. As demonstrated by Brenan et al. (2012), this effect may become evident when considering the partitioning involving spinels in which the chromite component is replaced by magnetite with increasing fO2, which is an important exchange in spinels from mafic and ultramafic magmas. Chromite (FeCr2O4) is a normal spinel, with all Cr in VI-fold coordination owing to the very high octahedral site preference energy (OSPE) of Cr3+ (d 3 valence electron configuration; McClure 1957; Dunitz and Orgel 1957). In contrast, magnetite (Fe3O4) is an inverse spinel at room temperature, but can show a decrease in the amount of divalent octahedral substitution at high temperature (e.g., Wißmann et al. 1998). A general formula for magnetite–chromite solid solutions takes the form: (Fe2+1-x Fe3+x)[Fe2+x Fe3+2-2z-x Cr2z]O4 where z = Cr3+/(Fe3+ + Cr3+). Site occupancies across the chromite–magnetite join, calculated after the method of Kurepin (2005), are displayed in Figure 16. For end-member chromite, the octahedral site is completely filled by trivalent cations (Cr3+), thereby restricting the uptake of divalent cations which prefer VI-fold coordination. As the magnetite component increases, however, there is a rise in the divalent cation occupancy of the octahedral site. In terms of HSE partitioning, Rh and Ir are likely to be dissolved as divalent species in oxide solutions at the fO2 of terrestrial magmas (Borisov and Palme 1995; O’Neill et al. 1995; Ertel et al. 1999; Brenan et al. 2005), with a d 7 valence electron configuration and in the low spin state (e.g., Zhang et al. 2010). There is, however, evidence for Ir3+ in the CMAS system at 1500 ºC (Fonseca et 08 0.8 0.7

o

site occupancies at 1300 C "inverse"

"normal"

0.6 0.5 site fraction 0.4

3+

Fe e (oc (oct))

3+

Cr (oct)

3+

[also Ru ]

3+

[also Ru ]

Fe2+(oct) 2+ 2+ [also Rh ,Ir ]

0.3

2+

Fe (tet)

0.2 0.1 00 0.0 0 Fe O 3

4

0.2

0.4 0.6 Cr/(Fe3+ + Cr)

0.8

1 FeCr O 2

4

Figure 16. Variation in the tetrahedral and octahedral site occupancies for di- and trivalent cations across the magnetite-chromite join. Curves are labeled according to the identity of the cation and the site occupied, e.g., Fe3+(oct) refers to the fraction of trivalent iron in VI-fold coordination. Note that the curve for Fe2+(oct) also corresponds to the amount of Fe3+ in tetrahedral coordination. All of the Cr3+ is assumed to be in VI-fold coordination. Values of the inversion parameter, x (see text) are calculated at 1300 ºC after the method of Kurepin (2005). The species Ru3+, Rh2+, and Ir2+ are likely to have strong affinity for octahedral coordination, so their substitution into spinel will be controlled by the proportion of di- and trivalent octahedral sites. Consequently, end-member chromite can accomodate Ru3+, but only with increased magnetite component are divalent octahedral sites available for Rh2+ and Ir2+. [Used by permission of Elsevier Limited, from Brenan, Finnigan, McDonough and Homolova (2012), Chemical Geology, Vol. 302−303, Fig. 1, p. 24.]

Experimental Fractionation of the HSE

43

al. 2011). Based on trends in cation size with charge in VI-fold coordination, Rh2+ and Ir2+ are estimated to have ionic radii of ~ 72 and 74 pm (1 pm = 10−12 m), respectively, similar to Fe2+ and Mg2+ (78 and 72 pm, respectively; Shannon 1976). Both HSE cations are therefore expected to have a high affinity for octahedral sites in the chromite structure, and owing to their divalent charge, incorporation of Rh and Ir into chromite is predicted to be sensitive to the degree of inversion, as influenced principally by the magnetite component. As described above, results of metal solubility measurements suggest that Ru is dissolved as a 3+ or 4+ cation in oxide solutions at moderate to high fO2 (Borisov and Nachtweyh 1998; Laurenz et al. 2013). Results of mineral–melt partitioning experiments seem to be most consistent with a 3+ oxidation state, which has a d 5 electron configuration, in low spin state (Geschwind and Remeika 1962), suggesting a strong affinity for octahedral coordination. Based on trends in cation size with charge in VI-fold coordination, Ru3+ is estimated to have an ionic radius of ~ 68 pm, similar to Fe3+ and Cr3+ (64.5 and 61.5 pm, respectively; Shannon 1976). These considerations are consistent with the overall compatibility of the IPGEs in chromium-rich spinels, and with larger partition coefficients as the magnetite component increases, owing to a higher abundance of divalent octahedral sites (Righter et al. 2004; Brenan et al. 2012). Brenan et al. (2012) developed a model for partitioning of di- and trivalent cations into chromiumspinel taking into account the variation in site occupancy with magnetite component, which provides reasonable agreement to the existing experimental data, as well as empirical estimates from natural samples (i.e., Puchtel and Humayun 2001; Locmelis et al. 2011; Pagé et al. 2012). In terms of the HSEs that are highly incompatible in olivine and spinel (Au, Pd, and Pt) either ionic radius or steric effects account for this behavior. At the conditions of the partitioning experiments, Au and Pd solubility systematics indicate the predominance of Au1+ and Pd1+ (Borisov et al. 1994; Borisov and Palme 1996), which in VI-fold coordination have ionic radii of 150 pm and 100 pm, respectively (estimated from Shannon 1976). These are considerably larger than either Mg2+ (72 pm) or Fe2+ (78 pm), so the low D values are consistent with a large mismatch in ionic radius compared to the dominant substituent cations in these phases. Metal solubility experiments have shown that Pt2+ is the likely oxidation state at the fO2 of past partitioning experiments (Ertel et al. 1999), although Pt4+ has been inferred spectroscopically, but only under highly oxidizing conditions (Farges et al. 1999). Pt2+ has a d 8 electronic configuration, and an ionic radius of 80 pm in VI-fold coordination (Shannon 1976), similar to Mg2+ and Fe2+, predicting a strong preference for octahedral sites in the olivine and spinel structures. The relatively low values for DPt would therefore seem anomalous. Documented occurrences of VI-fold complexes containing Pt2+ are rare, however, with the square planar coordination being most common, stabilized by the enhanced bond strength overwhelming the pairing energy required for this configuration (Cotton and Wilkinson 1988). Although square planar sites are unavailable in chromite and olivine, it may be that Pt2+ forms such complexes in the coexisting silicate melt, accounting for the low value for DPt.

Local PGM saturation during chromite growth Both empirical observations and recent experiments suggest that the HSE may not always be fractionated by forming a homogeneous solution in a mineral or melt phase. For example, the occurrence of PGM inclusions in chromite is well-documented, with the most common association involving minerals of the laurite–erlichmanite series (RuS2–OsS2), as well as Pt–Fe and Os–Ir–Ru-bearing alloy (Legendre and Augé 1986; Talkington and Lipin 1986; Garuti et al. 1999; Merkle 1992; Cabri et al. 1996; Gervilla and Kojonen 2002; Zaccarini et al. 2002). Phase equilibrium experiments confirm these PGMs are stable at chromian spinel liquidus temperatures over some range of fS2, provided the system is undersaturated with sulfide liquid, indicating that such inclusions can be interpreted as a primary magmatic texture (Brenan and Andrews 2001; Andrews and Brenan 2002; Bockrath et al. 2004). The origin of these inclusions has been somewhat enigmatic, but may in part account for the association

44

Brenan, Bennett & Zajacz

between chromite and enrichment of certain HSE. In the course of experiments designed to measure chromite–silicate melt partitioning of the HSE, which involved re-equilibration or growth of chromite in molten silicate, Finnigan et al. (2008) documented the occurrence of PGMs (including metal alloys and laurite) at the crystal–liquid interface. The mechanism of formation proposed by these authors involves the development of a redox gradient, owing to local reduction within the mineral–melt interfacial region, occurring as a consequence of the selective uptake of trivalent Cr and Fe from the melt, relative to the divalent species. Recalling that at conditions more oxidizing than the IW buffer, the solubility of the HSE increases with fO2, hence local reduction provides a driving force for precipitation of the PGMs in magmas that are not too far from metal saturation. Finnigan et al. (2008) modeled the processes of growth, as well as crystal–melt re-equilibration by Cr–Al exchange, to show that sufficient reduction occurs such that metal solubilities will decrease by several percent in the silicate melt at the melt–crystal interface. Once a sufficient degree of oversaturation occurs, the PGMs nucleate and continue to grow until either the redox gradient dissipates, or they become entrapped within the adjacent chromite crystal. Inclusion of such PGMs, then subsequent accumulation of chromite, constitutes a mechanism to fractionate the HSEs via mechanical means, rather than as a dissolved component in a major crystallizing phase. González-Jimenez et al. (2009) argue that the occurrence of zoned laurite–erlichmanite grains entrapped in chromite from different ophiolite localities arise from the redox gradients induced by chromite growth, providing support for the Finnigan et al. (2008) model.

MAGMATIC SULFIDE AND ASSOCIATED PHASES During melting or solidification, sulfur-bearing silicate magmas can reach saturation in a sulfide phase, typically rich in Fe, with lesser amounts of Ni and Cu. Phase equilibrium experiments on typical magmatic sulfide compositions predict an immiscible sulfide liquid to form crystalline Fe-rich monosulfide solid solution (MSS; [Fe,Ni]1-xS), which at 0.1 MPa occurs at Tmax of 1190 ºC, corresponding to Fe0.917S (Jensen 1942). The exact liquidus will depend on pressure, Ni and Cu content and fS2/fO2 (Naldrett 1969; Fleet and Pan 1994; Ebel and Naldrett 1996; Bockrath et al. 2004). With cooling, MSS is followed at T of < 900 ºC by a Cu-rich Intermediate Solid Solution ([Cu,Fe]1-xS; ISS; Dutrizac 1976), and magnetite, with the MSS–ISS assemblage undergoing subsolidus crystallization to mostly pyrrhotite (Fe1-xS), pentlandite ((Fe,Ni)9S8) and chalcopyrite (CuFeS2). In the crustal environment, there is evidence for efficient magmatic sulfide differentiation associated with relatively large igneous bodies, as documented in the world-class Ni–Cu–PGE deposits of the Sudbury (Canada) and Norilsk-Talnakh (Russia) Districts, for example, with separation of ores rich in Fe and Ni, interpreted as MSS cumulates, from those which are Cu-rich, representing mixtures of evolved sulfide liquid and cumulate ISS (e.g., Naldrett et al. 1992, 1996; Li and Naldrett 1992; Zientek et al. 1994; Ballhaus et al. 2001; see also Barnes and Ripley 2016, this volume). This process has resulted in a significant separation of the HSE, with the IPGE and Re concentrated in the MSS cumulates, and the PPGE and Au following the evolved liquid. Past studies of sulfide in upper mantle peridotites and diamonds have also identified both trapped sulfide liquid and residual MSS, albeit on a much smaller scale (Szabó and Bodnar 1995; Guo et al. 1999; Alard et al. 2000; Lorand and Alard 2001; Luguet et al. 2001, 2004). In that context, a distinction is made between so-called Type 1 and Type 2 sulfides, using criteria and nomenclature from Luguet et al. (2001). Type 1 sulfides are characterized by high Ni relative to Cu abundances, and primitive upper mantle (PUM)normalized depletions in Rh and Pd relative to Ir (as well as Ru and Os), and interpreted to be residual MSS. Type 2 sulfides have variable Ni/Cu, and similar PUM-normalized abundances of Ir (Ru, Os), Rh, and Pd, considered to be consistent with trapped immiscible sulfide liquid. Although sulfur has long been implicated as an important ligand in the

Experimental Fractionation of the HSE

45

concentration of the HSE in magmatic sulfide systems, field evidence suggests that As, and the other chalcogens Se, Te, and Bi could be important in some cases. For example, Gervilla et al. (1996, 1998), Hanley (2007) and Godel et al. (2012) have reported a close textural relationship between relatively PGE–(Au)-depleted base-metal sulfide and coexisting PGE–(Au)-rich arsenide phases (NiAs, nickeline; Ni11As8, maucherite; NiAsS, gersdorffite) in the magmatic sulfide segregations within the Ronda and Beni Besoura peridotite bodies, the Kylmakoski (Finland) Ni–Cu deposit and komatiite-hosted base-metal sulfide mineralization (Dundonald Beach South, Ontario; Rosie Ni Prospect, Western Australia), implying the presence of an immiscible arsenide melt at the magmatic stage. Recent work on samples from Creighton Mine, Sudbury (Dare et al. 2010) have shown that the base-metal sulfides are not the dominant hosts for some PGE, and that Ir, Rh, Pt occur as chalcogenrich discrete platinum-group minerals (PGMs; i.e., irarsite–hollingsworthite, IrAsS–RhAsS; sperrylite, PtAs2), possibly crystallizing before or with MSS. Chalcogen-bearing phases are also associated with late-stage low sulfur precious metal haloes around massive sulfide bodies, as has been documented at various locations around Sudbury, Ontario (Farrow and Watkinson 1997). There is also evidence for remobilized chalcogen-rich melts associated with high grade metamorphism of base-metal deposits (Frost et al. 2002; Tomkins et al. 2007). Thus the chalgogens may affect the distribution of HSE within a magmatic sulfide system in several ways, including early sequestration as immiscible semi-metal-rich liquids or discrete PGMs at the magmatic stage, ligands to maintain the PGEs in solution during ore solidification, and agents of remobilization during subsequent metamorphism. In this section, we focus primarily on the results of experiments to measure the partitioning between base-metal sulfide phases (MSS, sulfide melt) and silicate melt, but also include the limited (but likely to grow) body of results available for the chalcogens. Past experimental studies to measure partitioning amongst MSS–sulfide melt–arsenide melt–silicate melt are listed in Tables A3 and A4 in the Appendix.

Experimental approach As is the case for solid metal–liquid metal partitioning experiments, loss of volatile sulfur, and some of the HSEs, is a concern, so experiments to measure sulfide–silicate partitioning are done using gas-tight containers, or under a S-bearing vapour phase. At 0.1 MPa, this involves either encapsulation in vacuum-sealed silica ampoules, or the use of sulfur-bearing gas mixtures (i.e., SO2–CO2–CO), whereas at high pressure, experiments are done in containers made from high purity graphite or natural mineral capsules (i.e., olivine), which prevents chemical interaction between the sample and the outer noble metal capsule which is typically used to ensure a gastight seal. In order to assist in the efficient separation of sulfide and silicate liquids during the experiment, Brenan (2008) and Mungall and Brenan (2014) subject samples to high acceleration at high temperature using a specially-designed centrifuge furnace (Roeder and Dixon 1977). Partitioning experiments done at 0.1 MPa with evacuated silica ampoules have employed several methods for either buffering, or monitoring of fO2 and fS2. For example, Mungall et al. (2005) used synthetic solid buffers to fix both fO2 and fS2 in their MSS–sulfide melt partitioning experiments (DMSS/SulLiq) done at 950–1050 ºC, using the combined equilibria: 3Fe2SiO4 + O2 = 2Fe3O4 + 3SiO2 (FMQ)

(61)

Pt + 1/2S2 = PtS

(62)

and:

This was accomplished by loading the buffer powders, along with the sample, in silica cups placed within an outer silica ampoule, with the experiment physically separated below the

46

Brenan, Bennett & Zajacz

buffers, but in communication via the gas phase. Liu and Brenan (2015) employed a similar approach in their MSS–ISS–sulfide melt partitioning experiments done at 860−926 ºC (Fig. 17), but without the Pt/PtS mixture, as it was found to readily absorb the chalcogens, As, Se, and Te for which partition coefficients were also measured, in addition to the HSE. Instead the fS2 was adjusted by the metal/sulfur ratio in the sample, and monitored using the FeS content of pyrrhotite added to the FMQ assemblage, using the calibration of Toulmin and Barton (1964). A similar approach was employed by Fleet and coworkers (Stone et al. 1990; Fleet et al. 1996, 1999) in their experiments to measure sulfide melt/silicate melt partitioning (DSulLiq/SilLiq) of PGEs and Au, with fO2 buffered at relatively reducing conditions using materials representing the following equilibria: 2Fe + SiO2 + O2 = Fe2SiO4 (IQF) (63) FeO + O2 = Fe3O4 (MW)

(64)

Fe + ½ O2 = FeO (IW)

(65)

or C-O gas equilibrium involving graphite, used either as a sample holder or added as a solid rod, also known as the CCO buffer, defined by the reactions: Cgraphite + ½ O2 = CO

(66)

CO + ½ O2 = CO2

(67)

Brenan (2008) and Mungall and Brenan (2014) achieved somewhat more oxidizing conditions in their sulfide melt-silicate melt partitioning in experiments done at 0.1 MPa and 1200 ºC, in which samples were encapsulated in crucibles made from natural olivine or chromite (see Figure 2 of Mungall and Brenan 2014). Sulfur fugacity was fixed using either Pt–PtS, Ru–RuS2 or Ir2S3–IrS2 buffers, the latter two involving the sulfidation reactions: Ru + S2 = RuS2

(68)

Ir2S3 + ½ S2 = 2IrS2

(69)

and

Once sulfur fugacity is fixed, Brenan (2008) calculated fO2 from the heterogeneous equilibrium: FeOsilicate melt + ½ S2 = FeSsulfide melt + ½ O2

(70)

whereas Mungall and Brenan (2014) used the Cr content of the silicate melt, held in a chromite crucible, to estimate fO2, as the solubility of chromite varies in response to changes in the speciation of chromium (Cr2+ and Cr3+; Berry and O’Neill 2004), hence fO2, as demonstrated by Roeder and Reynolds (1991). Partitioning experiments done at high pressure have employed graphite caspules (in some cases in a sealed Pt outer capsule; Peach et al. 1994; Sattari et al. 2002), which fix fO2 near the CCO buffer, or techniques in which the external fH2 is controlled, allowing a range of fO2 to be investigated. One method of external fH2 control employs a double capsule configuration, in which the sample + H2O is loaded into a hydrogen permeable noble metal inner capsule, then sealed, and placed into an outer capsule containing an assemblage of H2O plus solid metal + oxides (i.e., Ni–NiO) or mineral mixtures (i.e., assemblages for Reactions (61), (63), and (64); Li and Audétat 2012). The external buffer fixes both fH2 and fO2, which is transmitted to the inner sample by H2 diffusion through the noble metal capsule. For experiments done with pressurized gas vessels, the fH2 can be buffered using water as the pressure medium, and the intrinsic fO2 of the vessel (Jugo et al. 1999; Simon et al. 2008), or by adding known amounts of H2 gas to the Ar pressure medium

Experimental Fractionation of the HSE (a)

47

(b)

Figure 17 (a) Configuration for experiments to measure MSS–sulfide melt partitioning of PGEs and the chalcogens employed by Liu and Brenan (2015) in which fO2 is controlled by the FMQ buffer (Eqn. 61 in the text), and fS is monitored using the composition of pyrrhotite added with the FMQ mixture (method of 2 Toulmin and Barton 1964). (b) Backscattered electron image showing homogeneous MSS crystals coexisting with quenched sulfide liquid from experiment done at 915 ºC at 0.1 MPa for three days. Dark globular structures are bubbles. [Used by permission of Elsevier Limited, from Liu and Brenan (2015), Geochimica et Cosmochimica Acta, Vol. 159, Fig. 1, p. 140 and Fig. 5a, p. 149.]

(Bezmen et al. 1994; Botcharnikov et al. 2011, 2013). The fS2 in the experiments in which fH2 is buffered is estimated using the FeS content of an added pyrrhotite sensor. Control of fS2 and fO2 in these experiments is important for several reasons. First, the oxidation state of the HSE can change with fO2, as has been previously described, as well as with fS2, as documented by Fonseca et al. (2007, 2009, 2011, 2012). Also, the degree of metal deficiency (metal/sulfur) in MSS varies with fS2, with metal-deficient MSS dissolving more of the HSE (Ballhaus and Ulmer 1995). The stability of sulfide liquid in molten silicate depends on the FeO content of the silicate, as well as the fO2/fS2 ratio (e.g., O’Neill and Mavrogenes 2002) through the heterogeneous equilibrium described by Equation (70), so sulfide may be absent if inappropriate fO2 or fS2 are imposed on the system. As the partitioning of the HSEs between sulfide and silicate melt can be expressed as an exchange reaction similar to 70 (see below), then the magnitude of partition coefficients will in turn depend on the relative fO2/ fS2 ratio. Some of the past experiments to measure partitioning amongst MSS-sulfide melt– silicate melt were done unbuffered, however, with the metal/sulfur of the MSS and sulfide melt varying with the bulk composition of the sample. It then becomes important, therefore, to relate the metal/sulfur to the fO2/fS2 in order to accurately apply the data to modeling natural systems. For experiments in which fO2 and fS2 can be measured or estimated, extrapolation of results to natural systems is less uncertain. As for the case of previous solubility and partitioning experiments, analysis of run products requires careful avoidance of inclusions (metal or sulfide) which contain a significantly higher HSE concentration than the phase of interest. This problem is not so acute in experiments involving sulfide mineral-sulfide melt equilibrium, although high spatial resolution is still important to obtain single phase analyses. The majority of past experiments to measure MSS–ISS–sulfide melt partitioning were doped at the 10−100 ppm level, which is close to the natural concentration range, then analyzed by either secondary ion mass spectrometry (SIMS; Fleet et al. 1993) or LA-ICPMS (Ballhaus et al. 2001; Mungall et al. 2005; Li and Audétat 2012; Liu and Brenan 2015). As sulfides dissolve relatively high concentrations of most HSE, some MSS-sulfide melt partitioning experiments were doped up to wt% levels, with runproducts measured by electron microprobe (Li et al. 1996; Brenan 2002), with the results of Li et al. (1996) reproduced using the more sensitive proton microprobe method (Barnes et al. 2001). In that case, it must be shown that Henry’s Law is valid over the concentration range investigated; this is demonstrated for Os at the ~ 2–5000 ppm levels and for Rh and Pd

Brenan, Bennett & Zajacz

48

at the ~ 10–20,000 ppm levels, by comparison of concentrations in MSS of similar metal/ sulfur ratio in the studies of Fleet et al. (1993), Li et al. (1996) and Brenan (2002). Earlier experiments to measure DSulLiq/SilLiq for the PGEs and Au were doped at the ~ 100−1000 ppm level, and analyzed by bulk methods (neutron activation; see summary in Table A3), whereas more recent determinations have been done at the > 1000 ppm doping level, with analyses by LA-ICPMS. As discussed below, extreme partitioning of the PGEs into the sulfide melt, as revealed by the in situ LA-ICPMS measurements, renders bulk measurements of quenched silicate highly susceptible to overestimation, (hence, underestimation of sulfide melt/silicate melt D-values) even if only miniscule amounts of sulfide contamination are present.

MSS–sulfide melt partitioning A summary of MSS–sulfide melt partitioning of the HSEs is provided in Figure 18. With the exception of some data for Re, Os and Au obtained at 1−3 GPa (Brenan 2008; Li and Audétat 2012, 2013), all other past measurements are from experiments done at 0.1 MPa, using evacuated silica ampoules (buffered or not). Values of DMSS/SulLiq for Ru, Os, Ir, Re, and Rh are > 1, with Ru as the most compatible (DMSS/SulLiq for Ru as high as ~ 20; Liu and Brenan 2015), and DOsMSS/SulLiq > DReMSS/SulLiq, giving rise to significant differences in the Os isotopic evolution within magmatic ore bodies (e.g., Lambert et al. 1998). In contrast, values of DMSS/SulLiq for Pt, Pd, and Au are all < 1. Fleet et al. (1993) noted that the partition coefficients between MSS and sulfide liquid change progressively between the three chemical subgroups of PGE, being higher for the iron triad (Ru, Os) and the lower for the nickel triad (Pd, Pt). Subsequent experiments have confirmed this observation, and extended it to the copper triad (Ag, Au), representing the most incompatible transition elements in MSS. In addition to these inter-element fractionations, there are some systematic differences in partitioning behavior amongst past studies, as noted below. These observations can be rationalized in the context of ligand field theory, as well as the effect of MSS and sulfide melt composition. Monosulfide solid solution has a NiAs-type structure, with triangular Fe clusters surrounded by distorted S octahedra, incorporating vacancies on the Fe sites and Fe3+ holes to satisfy the charge imbalance in metal deficient MSS (see review by Wang and Salveson 2005). Ballhaus and Ulmer (1995) showed that Pt and Pd (and by extension, the other PGEs, Re and Au) substitute for Fe in MSS on a one-for-one basis. Insight into the possible mineral structure control on HSE incorporation into MSS can therefore be gained by considering the relative solubilities of the PGEs in MSS compared to a fixed standard state (pure metal or 2

10

1

D(MSS/SulLiq)

10

0

10

-1

10

-2

10

-3

10

Liu and Brenan (2015) Fleet et al. (1993) Li et al. (1996) Ballhaus et al. (2001) Mungall et al. (2005) Li and Audetat (2013) Brenan (2002) Li and Audetat (2012)

Ru

Os

Ir

Re

Rh

Pd

Pt

Au

Figure 18. Summary of MSS/sulfide melt partitioning of the HSEs. [Used by permission of Elsevier Limited, from Liu and Brenan (2015), Geochimica et Cosmochimica Acta, Vol. 159, Fig. 7, p. 149.]

Experimental Fractionation of the HSE

49

pure metal sulfide). Past work has shown that the solubilities of Os, Rh, Pt, and Pd in MSS increase with decreasing metal/sulfur, M/S (Makovicky and Karup-Møller 1993, 2002; KarupMøller and Makovicky 2002; Majzan et al. 2002; Makovicky et al. 2002), indicating a decrease in metal activity coefficients, and that PGE substitution is enhanced by the presence of Fe vacancies (Ballhaus and Ulmer 1995). With this in mind, values of DMSS/SulLiq for the PGEs and Au are portrayed as a function of the M/S in the MSS in Figure 19. Where there are data for a significant range in MSS composition (Rh, Ir, Pd, Pt), partition coefficients show a weak increase with decreasing M/S—a trend consistent with the metal solubility results. Although the overall variation in partitioning with MSS composition seems consistent with expectations, the sense of HSE fractionation by MSS-melt partitioning is not reflected in the metal solubility data. Specifically, 1) Rh is found to be more soluble than Os, 2) the solubility of Pd is significantly higher than Pt, and 3) Os and Pt have similar solubilities. These differences are inconsistent with the overall incompatibility of Pt and Pd relative to Os and Rh, and the generally similar DMSS/SulLiq for the pairs Pt–Pd and Os–Rh (Fig. 18). As proposed by Liu and Brenan (2015), the inconsistencies in this comparison imply that MSS-melt partitioning of the PGEs must also be controlled by coordination complexes formed in the sulfide melt phase. Whereas Ru, Rh, Ir, and Os and Re are in octahedral coordination in their known sulfides, both Pd and Pt are in IV-fold coordination (summarized in Raybaud et al. 1997). Notably, at the conditions of past experiments, the likely oxidation state for Pt and Pd in molten sulfide is 2+ (Fonseca et al. 2009), which has a d 8 electronic configuration, and hence stabilized by square planar coordination (cf., Cotton and Wilkinson, 1988). In the absence of such sites in the NiAs-type structure, it therefore seems reasonable that both Pt and Pd are stabilized in IV-fold coordination by the more “permissive” sulfide-liquid structure. A similar argument may also hold for Au, which, assuming a 1+ oxidation state, is stabilized in low coordination number (II-fold to IV-fold) complexes (Carvajal et al. 2004). Hence, although the PGEs may be soluble in the MSS structure, it seems that their relative preference for the melt or solid phase depends on which coordination environment is most energetically favored. Both Mungall et al. (2005) and Liu and Brenan (2015) report partition coefficients for the compatible HSE (Ru, Re, Os, Ir, and Rh) that are systematically higher than other studies for a given MSS composition (Fig. 19). Whereas all previous work to measure partitioning of these elements was done unbuffered, and nominally oxygen free, as noted above, both Mungall et al. (2005) and Liu and Brenan (2015) employed techniques to fix fO2 at the FMQ buffer, resulting in sulfide melt with oxygen contents at the 1−2 wt% level. Results of previous experiments have documented a sharp decrease in the solubilities of Re, Os, Ir, Ru (and Pt) in molten sulfide at an fO2 of ~ FMQ − 2 to − 3 (depending on the metal, and the fS2; Fonseca et al. 2007, 2009, 2011; Andrews and Brenan 2002), corresponding to a sharp rise in the oxygen content of the sulfide liquid from nil to ~ 1−5 wt%. The solubility decrease over this interval is ~ 10-fold, and implies a complementary increase in the activity coefficient for these metals in the melt. The effect of an increase in the activity coefficient for a metal in the melt phase would be to increase DMSS/SulLiq, which is the sense of the offset noted above. In this context, it is also worth mentioning that the addition of Cu and Ni to an Fe–S melt composition has been shown to change the solubility of Ru, Ir and Os (Fonseca et al. 2007, 2009, 2011; Andrews and Brenan 2002; Brenan 2008) with these additives acting in opposite ways. Whereas Ni increases the solubility of these metals (e.g., 0−23 wt% Ni results in ~ 2-fold increase in Os solubility), Cu results in a decrease (e.g., 0−26 wt% Cu results in a 3-fold drop in Os solubility; Fonseca et al. 2011), implying sympathetic changes in the activity coefficients for these PGEs in the melt phase. Hence, the relatively high copper content (~ 30 wt%) of the melts produced in the study of Liu and Brenan (2015) compared to previous work (~ 4 to 13 wt%) would also result in a modest increase in partition coefficients. Although the variation in D with M/S for the compatible PGEs seems consistent with known activity-composition relations in the sulfide melt, the significant differences in partitioning seen for Au, and to a lesser extent Pt and Pd, is less clear. For MSS with a similar range in M/S, values of DMSS/SulLiq for Au are found to vary by ~ 10-fold, with results from Li and Audétat (2013) and

Brenan, Bennett & Zajacz

50

(a) 28

Liu and Brenan (2015)

40

Fleet et al. (1993) Brenan (2002)

16 12 8

20

10

0

24

0.87 0.93 M/S in MSS

0.99

0 0.75

1.05

(c)

0.81

0.87 0.93 M/S in MSS

15

20

D(MSS/SulLiq) Rh

D(MSS/SulLiq) Ir

0.81

Liu and Brenan (2015) Li et al. (1996) Mungall et al. (2005) Fleet et al. (1993) Ballhaus et al. (2001)

28

Fleet et al. (1993)

30

4

0.75

Mungall et al. (2005) Ballhaus et al. (2001)

20

D (MSS/SulLiq) Ru

D(MSS/SulLiq) Os

24

(b)

Liu and Brenan (2015)

16 12 8

0.99

1.05

(d) Liu and Brenan (2015) Li et al. (1996) Mungall et al. (2005) Fleet et al. (1993) Ballhaus et al. (2001)

10

5

4 0.8

0.85

0.9 M/S in MSS

0.95

1

(e)

0.5

Liu and Brenan (2015) Li et al. (1996) Mungall et al. (2005) Fleet et al. (1993) Ballhaus et al. (2001)

0.4 D(MSS/SulLiq) Pd

0 0.75

1.05

0.2

0.1

12

0.85

0.9 M/S in MSS

0.95

1

1.05

(f) Liu and Brenan (2015) Li et al. (1996) Mungall et al., (2005) Fleet et al. (1993) Ballhaus et al. (2001)

0.25

0.3

0 0.75

0.8

0.3

D (MSS/SulLiq) Pt

0 0.75

0.2 0.15 0.1 0.05

0.81

0.87 0.93 M/S in MSS

0.99

0 0.75

1.05

(g)

0.81

0.87 0.93 M/S in MSS

0.99

(h)

100

Liu and Brenan (2015) Brenan (2002)

1.05

10 D (MSS/SulLiq) Au

D(MSS/SulLiq) Re

10-1

8 6 4

10-2

10-3

2 0 0.75

-4

0.8

0.85

0.9 M/S in MSS

0.95

1

1.05

10

0.75

Liu and Brenan (2015) Mungall et al. (2005) Fleet et al. (1993) Li and Audetat (2013) Li and Audetat (2012)

0.8

0.85

0.9 M/S in MSS

0.95

1

1.05

Figure 19. Summary of MSS/sulfide melt partition coefficients for the HSE plotted as a function of the metal/sulfur (M/S) ratio of the coexisting MSS. With the exception of the studies by Mungall et al. (2005), Liu and Brenan (2015) and Li and Audétat (2012, 2013), all other experiments were done unbuffered, with the sulfide liquid nominally oxygen-free. [Used by permission of Elsevier Limited, from Liu and Brenan (2015), Geochimica et Cosmochimica Acta, Vol, 159, Fig. 13, p. 154−155.]

Experimental Fractionation of the HSE

51

Fleet et al. (1993) recording higher values than past determinations. Unlike past experiments, in which the HSE were added at ppm to low wt% levels, experiments done by Li and Audétat (2013) were at saturation in pure Au, corresponding to 9 to ~ 15 wt% Au in the sulfide liquid. The effect of such high metal loading on partitioning is unknown, but could very well be outside the concentration limits of Henryian behavior, and are certainly beyond natural abundance levels. Hence, such anomalously high values of DMSS/SulLiq for Au could be reasonably excluded as applicable to modeling natural processes. Fleet et al. (1993) also measured elevated values of DMSS/SulLiq for Au, as well as Pt and Pd. HSE dopant levels were low, so adherence to Henry’s Law is likely not an issue, and the composition of MSS and sulfide melt are similar to previous work. The only difference in method was the use of SIMS for sample analysis, with partition coefficients for Au determined using the ratio of sulfur-normalized count rates in the MSS and sulfide melt. As documented by Fleet et al. (1993), this is a robust technique for measuring Au in sulfides. However, it is possible that the rather small spot employed (20−30 mm), and small number of analyses acquired (2) might not have fully captured the true variation in the Au content of the texturally inhomogeneous quenched sulfide melt (see Fig. 17b).

MSS–ISS–sulfide melt partitioning Experiments to measure the partitioning of the HSE between ISS and MSS have been done by Jugo et al. (1999) at 100 MPa, 850 ºC (Au) and Liu and Brenan (2015) at 0.1 MPa and 850–875 ºC (PGEs, Re, Au). Results are summarized in Figure 20, which shows that Ru, Os, Ir, Rh, and Re are more compatible in MSS than in ISS, whereas Pd, Pt, and Au partition preferentially into ISS. Results for Au partitioning are consistent between the two studies. ISS–sulfide melt partition coefficients (DISS/SulLiq) were estimated by Liu and Brenan (2015) by combining their average values for MSS–sulfide melt, and MSS–ISS partitioning (Fig. 20). The calculated partition coefficients indicate that all the HSE should behave similarly to each other when partitioning between ISS and melt, with each weakly preferring melt relative to ISS.

Sulfide melt–silicate melt and MSS–silicate melt partitioning Experiments to measure the sulfide melt-silicate melt partitioning of the HSEs have been of two types. Most experiments involve equilibration of both sulfide and silicate in the same sample, with the added HSE usually below the solubility limit (i.e., activity of the HSE, aHSE, < 1). Another, much smaller subset of experiments are of the “indirect” type, in which the concentration of a particular HSE in either sulfide or silicate melt is measured at (or corrected to) saturation

2

10

1

partition coefficient

10

0

10

-1

10

-2

10

-3

10

Average MSS/SulLiq (Liu and Brenan, 2015) Average MSS/ISS (Liu and Brenan, 2015) MSS/ISS (Jugo et al., 1998) Calculated ISS/melt (Liu and Brenan, 2015)

-4

10

Ru

Os

Ir

Re

Rh

Pd

Pt

Au

Figure 20. Measured values of MSS/ISS partitioning (Liu and Brenan 2015; Jugo et al. 1999), and calculated values of ISS–sulfide melt partitioning based on the average MSS–sulfide melt partition coefficients of Liu and Brenan (2015).

52

Brenan, Bennett & Zajacz

(aHSE = 1), but in separate experiments involving a single sulfide or silicate melt phase (Andrews and Brenan 2002; Fonseca et al. 2007, 2009, 2011). In this approach, partition coefficients are calculated by the ratio of concentrations in the sulfide melt–silicate melt measured in the separate experiments, for which the metal activity is the same. The rationale for this approach is that it allows for the fO2–fS2 solubility systematics for each phase to be measured independently, and without the complication of sulfide contamination of the silicate phase. Rhenium. Experiments to measure values of DSulLiq/SilLiq for rhenium have employed both of the aforementioned methods. In terms of the “indirect” approach, Fonseca et al. (2007) measured the solubility of rhenium metal in molten sulfide at 0.1 MPa, 1200−1400 ºC and fO2 of ~ FMQ − 6 to − 2 (and variable fS2, using a gas-mixing furnace) which, combined with the solubility measurements of Ertel et al. (2001) for rhenium in diopside–anorthite eutectic melt, yielded estimates of DSulLiq/SilLiq. Fonseca et al. (2007) showed that, up to an fO2 equivalent to FMQ − 2, the solubility of Re in molten sulfide is independent of fO2, but then exhibits a sharp decrease (data at higher fO2 could not be obtained). At fixed fO2 of FMQ − 4.4, Re solubility in molten sulfide shows a progressive increase with fS2, consistent with a change in speciation from Re0 (low fS2) to Re4+ (high fS2). As mentioned previously in the context of the mineral–melt partitioning of Re, Ertel et al. (2001) modeled their silicate melt solubility data in terms of contributions from both Re4+ and Re6+, with the species equivalence point at ~ FMQ−3, so the presence of Re4+ is consistent with expectations. Fonseca et al. (2007) developed a partitioning model which takes into account the combined fO2 and fS2 dependences on solubility, and showed that at the fO2–fS2 conditions of mid-ocean ridge basalt (MORB) genesis (high fS2, and reduced; FMQ − 1 to − 2), DSulLiq/SilLiq for Re is ~ 1–100, thereby exhibiting chalcophile behavior. However, at the conditions of island arc basalt (IAB) genesis (low fS2; oxidized; FMQ + 2), partition coefficients are ~ 1 × 10−4, hence Re would become strongly lithophile. Brenan (2008) measured the partitioning of Re between coexisting sulfide melt and silicate melt at 1200 ºC, 0.1 MPa and ~ FMQ – 2 to FMQ + 1, with fS2 buffered using Equilibria (62), (68), and (69). Most experiments were done by first equilibrating the two melts at static conditions, then subjecting samples to high acceleration to enhance phase separation. Similar to the results predicted from the work of Fonseca et al. (2007), DSulLiq/SilLiq for Re was found to vary over a wide range, from > 20,000 to ~ 20, depending on the fO2–fS2 conditions imposed on an experiment. Following the approach of Gaetani and Grove (1997), the origin of this variation was modeled by Brenan (2008) according to the exchange of rhenium between molten sulfide and silicate as expressed by the reaction: RexOy,silicate + z/2 S2 = RexSz,sulfide + y/2 O2

(71)

The extent to which the partition coefficient is sensitive to fO2and fS2 depends on the value of the stoichiometric coefficients, x, y and z4. Normalized to one cation, this reaction becomes: (72) ReOy/x,silicate + z/2x S2 = ReSz/x,sulfide + y/2x O2 which has an equilibrium constant of the form: K5−12 = [ReSz/x,sulfide] fO2 y/2x / [ReOy/x,silicate] fS2 z/2x (square brackets denoting activities) and can be rearranged to yield: log {[ReSz/x,sulfide]/ [ReOy/x,silicate]} = log fS2z/2x – log fO2 y/2x + log K5−12

(73)

(74)

Assuming that the ratio of activity coefficients in the sulfide and silicate melts is constant 4

The reader is also referred to Kiseeva and Wood (2013) who propose an alternate approach, that makes use of an exchange reaction involving the element of interest and Fe in the sulfide or silicate, and does not have an explicit dependence on fO2 or fS2; this requires knowledge of the silicate melt FeO content, however, making it difficult to compare results with the model of Fonseca (2007).

Experimental Fractionation of the HSE

53

over the fO2–fS2 range of experiments, then this value can be combined with K5−12 and the factor to convert moles to wt%, to yield a single constant, KK5−14. Then, by assuming that z = y, Equation (74) becomes: (75)

log DSulLiq/SilLiq = y/x {½ log fS2 – ½ log fO2} + KK5−14

If the above conditions are satisfied, then a plot of log DSulLiq/SilLiq vs. ½ log fS2 – ½ log fO2 should yield a linear relationship, with the slope equal to the anion to cation ratio for the rhenium species. The variation in DSulLiq/SilLiq modeled in this way yielded two linear, but offset, data trends, defined by the fS2 of the experiments (Fig. 21). Treated separately, the low fS2 data define a slope of ~ 3, consistent with predominantly Re6+ in both silicate and sulfide, whereas the high fS2 data are defined by a shallower slope (2.4). A possible reason for the shallower slope in the high fS2 data set is the presence of Re–S species in the silicate melt, which has been shown to occur for Ru (Laurenz et al. 2013), Pt (Mungall and Brenan 2014), Ni (Peach and Mathez 1993; Li et al. 2003), and Cu (Ripley et al. 2002). For example, Re could be dissolving in molten silicate as a mixed Re–O–S species, by the model reaction (assuming Re6+ as the dissolved species in molten sulfide at the fO2 of experiments): ReOqS3-q, silicate + q/2 S2 = ReS3, sulfide + q/2 O2

(76)

which has an equilibrium constant of the form: K5−16 = [ReS3,sulfide] fO2q/2 / [ReOqS3-q, silicate] fS2q/2

(77)

and can be cast in a similar fashion as Equation (75) to yield: log DSulLiq/SilLiq = q {½ log fS2 – ½ log fO2} + KK5−17

(78)

in which KK5−17 is a combination of K5−16 and a mole to wt% conversion factor. For the slope of 2.4 exhibited by the high fS2 experiments, the proportion of sulfur-bonded Re is estimated Ru-RuS2

6.0

Sattari et al 2002

5.0

IrS2-Ir2S3

Pt-PtS Pt-PtS hi Ni 1.5 GPa Loihi

4.0 log D

MORB

sulfide/silicate

3.0 2.0

-1.74 -0.39

1.0

model curves from Fonseca et al (2007) 0.0

3

3.5

4 1/2 logfS - 1/2 logfO 2

4.5

5

2

Figure 21. Sulfide melt–silicate melt partition coefficients for Re as a function of ½ log fS2 – ½ log fO2, for experiments involving direct measurements on coexisting sulfide and silicate (Sattari et al. 2002; Brenan 2008), as well as empirical estimates from coexisting phases in oceanic basalts (Roy-Barman et al. 1998; Sun et al. 2003). Curves are calculated from the model of Fonseca et al. (2007) based on the solubility of Re metal in sulfide and silicate melts, corrected to 1200 ºC.

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as (3 − 2.4)/3 × 100 = 20%. Also included in Figure 21 are the partition coefficients estimated empirically from sulfide and glass in oceanic basalts from Loihi and FAMOUS using data from Roy-Barman et al. (1998) and Sun et al. (2003), as well as curves calculated using the model of Fonseca et al. (2007) corrected to 1200 ºC. The model curves capture the low fS2 measurements remarkably well, although the slope of the model curve is more shallow, as it assumes Re4+ as the main sulfide species. Calculated partition coefficients are systematically higher than the high fS2 measurements, however. If the lower values of DSulLiq/SilLiq determined by Brenan (2008) result from sulfur complexing in the silicate melt, then this would not be captured in the Fonseca et al. (2007) model, as it relies on Re solubility measured for molten silicate under sulfur-free conditions. This is one aspect of the “indirect” method that may compromise the accuracy of calculated values for DSulLiq/SilLiq, given the evidence for metal-sulfur complexing. In light of the strong dependence of DSulLiq/SilLiq for Re on ½ log fS2 – ½ log fO2, it is worth considering whether large variations in sulfide–silicate partitioning are likely in different mantle environments, as predicted by the model of Fonseca et al. (2007). Although Fonseca et al. (2007) correctly point out that differences in fO2 and fS2 likely exist between MORB and IAB sources, an important aspect not considered is the requirement that, for coexisting sulfide and silicate melts, the value of ½ log fS2 – ½ log fO2 will be fixed by the heterogeneous equilibrium between FeO in the silicate melt and FeS in the sulfide melt (described by Reaction 70), and therefore if the [FeO] and [FeS] in the melts does not vary much, neither will DSulLiq/SilLiq (see also Kiseeva and Wood 2013). Although it is not easy to predict the variation in [FeS] in different basalt sources, the range in [FeO] during melting is reasonably well constrained by the iron content of primitive magmas evolving by olivine control (Francis 1985, 1995). In this context, the range in the iron content of primary Phanerozoic magmas is rather limited, and bounded by picritic lavas from Iceland (~ 8 wt% FeO; Jakobsson et al. 1978) and Hawaii (~ 11 wt% FeO; Humayun et al. 2004). For a fixed [FeS] of ~ 0.7, which corresponds to XFeS in a sulfide liquid with 15 wt% Ni, meant to be in equilibrium with mantle olivine with 3000 ppm Ni (Bockrath et al. 2004), this range corresponds to values of ½ log fS2 – ½ log fO2 of 3.99 (11 wt% FeO) to 4.13 (8 wt% FeO) or Dsulfide/silicate of ~ 380 to ~ 820 (using the high fS2 partitioning trend, for example). It is important to note that although DSulLiq/SilLiq is not likely to vary more than ~ 2-fold for this range of FeO, the bulk partition coefficient for Re could change markedly, due to the effect of fO2 on the silicate and oxide-melt partition coefficients (Mallmann and O’Neill 2007), or by differences in modal sulfide content. For example, low Re/Os and Re abundances in lunar basalts suggest that Re becomes more compatible in the residual assemblage at the reduced fO2 of the lunar mantle (Birck and Allegre 1994; Day et al. 2007). If the lunar basalt source is sulfide saturated, Dsulfide/silicate for Re may be somewhat higher than for terrestrial MORB genesis, owing to the much lower fS2 required for sulfide saturation, making the low fS2 partitioning results most applicable. Compounding this effect is the overall increase in bulk solid–melt partition coefficients for Re, as a consequence of the higher compatibility of Re4+ in the peridotite phase assemblage (Mallmann and O’Neill 2007). For the case of highly oxidized arc environments, sulfide is likely to be destabilized in the mantle source (Mungall 2002), combined with a higher abundance of the more incompatible Re6+, resulting in very low bulk partition coefficients. Gold. Values of DSulLiq/SilLiq and DMSS/SilLiq for gold are summarized in Figure 22. With reference to the exchange reaction described by Equation (72), past solubility experiments have shown that Au1+ is the likely dissolved species in molten silicate at conditions more oxidizing than FMQ − 4 (Borisov and Palme 1996), so partitioning data plotted in the form of Equation (75) should have a slope of one half. Consideration of the full datasets for either MSS—or sulfide liquid–silicate liquid partitioning shows considerable scatter, with no welldefined linear relation in the manner predicted. Selecting just experiments done below Au saturation (for reasons described in MSS–ISS–sulfide melt partitioning), and measured by

Experimental Fractionation of the HSE Bezmen et al 1994 Crockett et al 1997 Stone et al 1990 Fleet et al 1999 Li and Audetat 2012 Mungall and Brenan 2014 Botcharnikov et al 2013

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Li and Audetat 2013 Botcharnikov et al 2013 (MSS) Li and Audetat 2012 (MSS) Jugo et al 1999 (MSS) Li and Audetat 2013 (MSS) Simon et al 2008 (MSS) Zajacz et al 2013 (MSS)

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Figure 22. Sulfide–silicate partitioning of Au as a function of ½ log fS2 – ½ log fO2. Solid symbols: MSS as the stable sulfide; open symbols: sulfide liquid as stable sulfide. Experiments labeled R and A were done with rhyolite and andesite silicate melt compositions, whereas all other experiments were done with basaltic silicate melts. Small font symbols are experiments measured by bulk analytical techniques (Fleet et al. 1990, 1999; Stone et al. 1990; Bezmen et al. 1994; Crocket et al. 1997; Jugo et al. 1999).

LA-ICPMS, presents a somewhat more coherent behavior for sulfide liquid–silicate liquid partitioning, as the data show a weak, albeit scattered correlation with the abscissa parameter, defining a slope of ~ 0.4 (single low value from Li and Audétat 2012, excluded), and an inferred oxidation state of + 0.8. The significant scatter, and generally low values of DSulLiq/SilLiq obtained by bulk analytical methods likely results from the presence of small amounts of sulfide contamination in the glass separate, which is avoided by the LA-ICP-MS method. However, the much larger values of DSulLiq/SilLiq reported by Bezmen et al. (1994) suggest more extreme partitioning of Au into sulfide liquid than suggested by any other studies. Bezmen et al. (1994) noted this discrepancy and attributed such large partition coefficients to the presence of hydrogen in the sulfide melt, enhancing the uptake of Au (and PGEs) in some undefined way. This interpretation is in conflict with the results for DSulLiq/SilLiq measured by Li and Audétat (2012), which were also done under high pressure, hydrous conditions, but overlap with values measured by Mungall and Brenan (2014) done under dry conditions at 0.1 MPa. Silicate melt composition may also affect HSE partitioning in sulfide-bearing systems, as Zajacz et al. (2013) has documented ~ 10-fold decrease in DMSS/SilLiq for melt compositions varying from rhyolite to basalt. The enhanced levels of Au (and therefore lower DMSS/SulLiq) in melts with lower silica content reflect a similar decrease in the activity coefficient for the dissolved HSE species as implied by the metal solubility data (see Role of silicate melt composition). The much lower values of DMSS/SulLiq measured for a rhyolite composition by Jugo et al. (1999) could reflect small amounts of trapped sulfide in the glass, as samples were analyzed by bulk methods. This does not seem surprising, since the more viscous rhyolite could trap small emulsified sulfide droplets easily. Simon et al. (2008) also report low values of DMSS/SilLiq involving a rhyolite composition, but with glasses measured by LA-ICPMS. Reported glass compositions are not homogeneous, however, with Au concentrations varying from < 0.2–6 ppm, suggesting either chemical equilibrium was not obtained, or the higher glass values are contaminated by

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ablated sulfide. Glasses produced in the study of Zajacz et al. (2013) are reported to contain Au metal contamination, which was minimized by accepting only the lowest portion of the time-resolved signal for quantitation. Accepting the lower glass values reported by Simon et al. (2008) would result in high values of DMSS/SilLiq, in line with the measurements of Zajacz et al. (2013). Hence, the Au-scavenging potential by early MSS crystallization in felsic systems may have been significantly underestimated by Simon et al. (2008). Perhaps the most notable aspect of the Au partitioning dataset is the systematically smaller values of DMSS/SilLiq than DSulLiq/SilLiq, which is especially well-defined when considering only those studies in which run-products were measured in situ by LA-ICPMS (e.g., Jugo et al. 1999; Li and Audétat 2012, 2013; Botcharnikov et al. 2011, 2013; Mungall and Brenan 2014) and involving similar melt compositions. The average value of DSulLiq/SilLiq measured from the in situ analytical studies on basaltic melts is ~ 4400. This contrasts with DMSS/SilLiq of ~ 170 measured in previous studies on the same composition (Li and Audétat 2012; Botcharnikov et al. 2011, 2013; Zajacz et al. 2013), yielding DSulLiq/SilLiq/DMSS/SilLiq of ~ 30, similar to the value of ~ 20 measured by Li and Audétat (2012), in which MSS, sulfide and silicate liquids coexist in the same experiment. Clearly, the identity of the residual sulfide phase will have a significant impact on the efficiency of gold partitioning into the silicate melt during mantle melting, and therefore the crust-tomantle transfer of this element. As shown by Bockrath et al. (2004), residual MSS can coexist with silicate melt, in the absence of sulfide liquid, along the low temperature, hydrous peridotite solidus. One may therefore expect that relatively low temperature mantle-derived melts, such as hydrous, alkalic compositions, produced in the presence of MSS, would contain a higher inventory of gold than dryer, high temperature magmas which leave behind residual sulfide liquid. This behavior was modeled in detail by Botcharnikov et al. (2013) and Li and Audétat (2013), who also emphasized the strong control of fO2 on the efficiency of mantle-to-crust transfer of gold, favouring environments in which residual sulfide is either eliminated by oxidation (Mungall 2002) or rapidly dissolved into the silicate melt under oxidizing conditions (Jugo 2009). Platinum-Group Elements (PGEs). A summary of sulfide–silicate melt partition coefficients for PGEs is shown in Figure 23, which includes the results from both laboratory measurements and values from natural samples. Most past measurements of DSulLiq/SilLiq have used bulk analytical methods to measure the metal content of glass and sulfide. The exceptions to this are the studies of Brenan (2008), and Mungall and Brenan (2014) in which analyses were done by LA-ICPMS. Also included in this comparison are the “indirect” estimates of sulfide-silicate partitioning done by Andrews and Brenan (2002) and Fonseca et al. (2009, 2012), which are based on the solubility of the metal in sulfide and silicate, measured in separate experiments involving only one melt phase. As is clear from the figure, nearly all of the previous measurements involving bulk analysis of the quenched silicate melt yield partition coefficients which are significantly lower than values determined either by in situ analysis of coexisting phases, or by the “indirect” method. Mungall and Brenan (2014) provide evidence to suggest the presence of micro-inclusions of sulfide melt in the silicate glass produced in their experiments, which would likely result in an overestimation of the intrinsic PGE content of the silicate melt, if measured by bulk methods. First, the time-resolved spectrum for the PGEs in run-product glasses was found to be inhomogeneous, with intensity peaks and troughs (similar to Fig. 4), with count-rates on all the PGEs added to a given experiment oscillating in unison. This is in contrast to the uniform signal observed for lithophile elements, like Ca, monitored at the same time. Second, both static and centrifuge partitioning experiments were done on Pt-doped experiments, and it was found that the Pt content of run-product glasses was always lower (by 24−70%) in the sample subject to high acceleration (~ 500 g). This suggests the high acceleration step had a cleansing effect on the silicate melt, removing some of the sulfide contamination owing to enhanced settling. In light of these observations, Mungall and Brenan (2014) concluded that past work, in which glasses were analysed by bulk methods, provided

Experimental Fractionation of the HSE Fleet et al., 1996

Mungall & Brenan (2014)

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Fonseca et al., 2009; 2011

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Roy-Barman et al., 1998 9

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Ir

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Figure 23. Summary of sulfide/silicate partition coefficients determined by experiment (Fleet et al. 1996; Crockett et al. 1997; Brenan 2008; Mungall and Brenan 2014) and from natural samples (Peach et al. 1990). Calculated values (Andrews and Brenan 2002; Fonseca et al. 2009, 2011) are based on the ratio of pure metal solubility in sulfide and silicate melt, calculated at FMQ − 1, and fS2 sufficient for sulfide saturation of a silicate melt with 15 mol% FeO. The S-corrected value for Ru (Andrews and Brenan 2002) corresponds to the ~ 30-fold increase in Ru solubility in sulfur-bearing silicate melt (Laurenz 2012), whereas the correction for Pt comes from the experiments of Mungall and Brenan (2014) in which solubilities in molten silicate are increased by ~ 100-fold relative to sulfur-free experiments.

minimum values of DSulLiq/SilLiq. They also suggested that glass concentrations measured by LA-ICPMS could also be susceptible to trace sulfide contamination, even if low count-rate domains of the time-resolved spectra are selected, and so-derived values of DSulLiq/SilLiq from their study should also be regarded as minimum. Accepting this interpretation therefore indicates that minimum partition coefficients for the PGEs are > 105, with some values for Ir and Pt documented by Mungall and Brenan (2014) exceeding 106. Calculated results of the “indirect” method to measure DSulLiq/SilLiq yield values for Os, Ir, Ru, and Pt of ~ 105, ~ 107, ~ 108 and ~ 109, respectively (Andrews and Brenan 2002; Fonseca et al. 2009, 2012), determined at FMQ − 1 and log fS2 of −1.44, which is sufficient for sulfide saturation of a silicate melt with 15 mol% FeO. It is currently unknown if the “true” partition coefficients could be as high as some of the values estimated from metal solubility experiments. As mentioned in the section Gold, there is now experimental evidence to suggest that metal–sulfur complexing could enhance the solubility of certain transition metals in molten silicate, so partition coefficients calculated using metal solubility in sulfur-free experiments could be overestimated in some cases. In the current context, Laurenz et al., (2012) have shown that the solubility of Ru in molten silicate is enhanced ~ 30-fold in sulfur-bearing (but not FeS-saturated) experiments, relative to sulfur-free compositions at the same fO2. This effect would reduce the Andrews and Brenan (2002) estimate for sulfide-silicate partitioning of Ru from ~ 108 to ~ 106, in closer agreement with the minimum values estimated by Mungall and Brenan (2014). In terms of the effect on Pt, Mungall and Brenan (2014) report enhanced solubility of this metal by ~ 100-fold in molten silicate in their sulfide-saturated partitioning experiments. This would translate to a ~ 100-fold reduction in the calculated value of DSulLiq/SilLiq for Pt to ~ 1 × 107, which is still higher, but in closer agreement to the direct measurements. The DSulLiq/SilLiq of ~ 105 estimated for Os by Fonseca et al. (2011) seems unusually low in the context of results for the other PGEs

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determined by the same method. One possible source of inaccuracy in that determination is the data for Os metal solubility in molten silicate. The source of solubility data is from Fortenfant et al. (2006), whose measurements were done by bulk methods, so there was no control on glass contamination by undissolved metal particles. Brenan et al. (2005) provide a maximum bound on Os solubility of 10 ppb, in demonstrably particle-free glasses produced at FMQ + 0.6, as measured by LA-ICPMS. At the same oxygen fugacity, Fortenfant et al. (2006) estimate the Os solubility to be ~ 1100 ppb. Selecting the lower solubility of Brenan et al. (2005) would yield DSulLiq/SilLiq for Os of ~ 106, which is more in line with the other PGEs estimated by the indirect method, and also for the minimum values for DSulLiq/SilLiq determined by Brenan (2008) for coexisting sulfide and silicate melts (Table A3, Fig. 23). The large values of DSulLiq/SilLiq implied by both the in situ measurements and indirect method have two important implications, in terms of alloy saturation during mantle melting, and the generation of extraordinarily PGE-rich sulfide deposits. Each is discussed in detail by Fonseca et al. (2009, 2012) and Mungall and Brenan (2014), and explored briefly below. Formation of PGE alloys during mantle melting Os–Ir–Ru-rich alloy grains constitute an important accessory phase in upper mantle peridotites, accounting for a significant fraction of the PGE in some samples (e.g., Luguet et al. 2001; Kogiso et al 2008). Such phases also appear to be remarkably refractory, and are likely responsible for the preservation of ancient (> 1 Ga) melt depletion events in much younger peridotite host rocks (e.g., Parkinson et al. 1998; Brandon et al. 2000). The presence of stable alloy grains is also an important consideration to models of PGE extraction from the mantle to the crust (Rehkämper et al. 1999; Mungall and Brenan 2014). Luguet et al. (2001) have proposed that the PGEs are initially concentrated into base-metal sulfide, and that alloys precipitate following the loss of sulfur during partial melting. In that context, Fonseca et al. (2012) have shown that saturation of a sulfide liquid in Os–Ir–Ru alloy is limited by the Os content of the liquid, as that is the least soluble component. Their results predict alloy saturation when the sulfide liquid concentration reaches ~ 40 ppm Os, assuming the availability of sufficient Ru and Ir. For mantle rocks with between 100−250 ppm sulfur, mass balance dictates that the base-metal sulfide which forms will be undersaturated in Os–Ir–Ru alloy for a typical undepleted upper mantle Os concentration of ~ 4 ppb (Becker et al. 2006). The key to alloy precipitation, therefore, is sufficient retention of Os (as well as Ir and Ru) in the sulfide phase, such that concentrations increase rapidly enough (due to the decrease in sulfide mass) to reach saturation before sulfide removal. Simple batch melting calculations which assume quantitative retention of the Os in sulfide result in alloy precipitation after 6−20% melting, depending on the total sulfur content of the source (Fig. 24; Fonseca et al. 2012). The occurence of alloy within this melting interval was also regarded by Fonseca et al. (2011) as necessary to reproduce the observed range in Re/Os in MORB, as the difference in DSulLiq/SilLiq for Re and Os was found to be too small. Mungall and Brenan (2014) showed that Pt metal saturation would occur during the sulfide-saturated melting interval, close to the point of sulfide exhaustion, for the same reasons described for Os. Importantly, they showed that use of the sulfur-free Pt solubility data to model the composition of mantle-derived magmas resulted in large, negative, Pt anomalies in model liquids, which are not observed in natural magmas. Instead, models in which solubilities are enhanced by Pt–S complexing provide the best fit to the data. Whereas Os was not included in their model, Mungall and Brenan (2014) did show that Ir and Ru alloy would also form in the restite just after complete exhaustion of sulfide liquid, in response to lowering fS2 and diminished metal–sulfide complexation in the silicate melt. Such results support a scenario in which alloys form as a result of the partial melting process, which is consistent with their occurence in sulfur-poor refractory mantle rocks, such as harzburgites of the Lherz massif (Luguet et al. 2007), depleted abyssal peridotites (Luguet et al. 2001), or in placer deposits likely derived from ophiolites (Nakagawa and Franco 1997; Meibom and Frei 2002; Pearson et al. 2007; Coggon et al. 2011), most of which are the residues of melting within a back-arc basin environment.

Experimental Fractionation of the HSE 100000

total sulfur = 100 ppm

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Figure 24. Calculated variation in the osmium content of residual sulfide melt as a function of the degree of batch melting for different sulfur contents of the source mantle ranging from 100 to 250 ppm (tick marks are at 10 ppm intervals, then as marked from 10 ppm and below). This encompasses the sulfur concentrations estimated for the depleted mantle (MORB source) of 119 (Salters and Stracke 2004) to 150 ppm (O’Neill 1991) to more elevated concentrations of 200 (O’Neill 1991) to 250 ppm (McDonough and Sun 1995) estimated for the primitive mantle. The solubility of sulfur in the silicate melt is assumed to be 1000 ppm, with source mantle Os of 3.9 ppb, corresponding to the primitive mantle composition of Becker et al. (2006). During melting of a sulfide-bearing source, quantitative sequestration of the HSE, including Os, into the residual sulfide phase, will cause concentrations in the residual sulfide to increase until saturation is reached. For reference is the estimated solubility of alloy having Os, Ru, and Ir in the same relative proportions as the primitive upper mantle (Ru62Os19Ir19; at%) and assuming that Os is the least soluble component (modified from Fonseca et al. 2012).

In the context of a discussion of the saturation of PGE-bearing alloy at the magmatic stage, it is also important to note the evidence for saturation following the ascent and emplacement of mantle-derived magmas. In all cases, the alloy composition is inferred to be either Pt- or Ir-( + Os + Ru)-rich (Peck et al. 1992; Barnes and Fiorentini 2008; Song et al. 2009; Ireland et al. 2009; Pitcher et al. 2009; Park et al. 2013). Evidence for saturation comes from either direct observation of alloy “phenocrysts”, in some cases with chromian spinel, olivine, and pyroxene included within metal grains (Heazlewood River Complex; Peck et al. 1992), or by the covariation of the PGE with other indices of differentiation, (e.g., Barnes and Fiorentini 2008; Ireland et al. 2009; Pitcher et al. 2009; Song et al. 2009; Park et al. 2013). As described in HSE Solubility Experiments: Implications for Metal–Silicate Partitioning, the solubility of the HSE under conditions more oxidizing than FMQ − 4 shows a general decrease with decreasing fO2, with the magnitude proportional to the formal oxidation state of the metal in molten silicate. There is a secondary, but less strong decrease in HSE solubility with decreasing temperature. As pointed out by Mungall and Brenan (2014), although the intrinsic temperature dependence on solubility is not strong, the absolute fO2 of a magma which is cooling along a buffer decreases markedly; at FMQ, the log fO2 changes from −5.4 at 1500 ºC to – 8.4 at 1200 ºC (O’Neill 1987). This magnitude of fO2 variation corresponds to a change in Pt solubility from 22 ppb to 0.7 ppb (Ertel et al. 1999); a compounding effect for FeO-bearing magmas is that the concentration level in the melt required for saturation will be below that for the pure metal owing to alloying with Fe (Borisov and Palme 2000), with higher Fe concentrations in the alloy at lower fO2. Hence, magmas that approach or reach alloy saturation during the melting process will tend to achieve and then continuously remain at alloy saturation as they cool during their ascent to the surface. Given their likely small size, the Stokes settling velocity of such alloy grains will be

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low, although attachment to larger grains, such as chromite or olivine, would likely enhance the segregation process (Hiemstra 1979; Ballhaus et al. 2006; Finnigan et al. 2008). Concentration of PGEs by R-factor processes. Previous work to understand the concentration mechanisms for the PGEs has emphasized the scavenging role of immiscible sulfide liquid at the magmatic stage, and the importance of the silicate to sulfide mass ratio, or R-factor (e.g., Campbell and Naldrett 1979). The R-factor, derived from mass balance considerations, is defined by the relation: (79)

R = [DSulLiq/SilLiq (Di – 1)] / [DSulLiq/SilLiq – Di]

In which Di is the ratio of the element concentration in the sulfide (Csulf) to the initial concentration in the silicate melt (Ci,silicate). Figure 25 depicts the relation between R and DSulLiq/SilLiq for different values of Di. The exceptionally-rich concentrations of PGE in sulfide horizons from the Stillwater and Bushveld complexes yield values of Di of > 105 (Campbell and Barnes 1984). Examination of Figure 25 indicates that it is not possible to achieve such high values of Di by an R-factor-like process if DSulLiq/SilLiq is in the range of 103−104, as determined by most previous partitioning studies involving bulk analytical techniques. This is because Di reaches a limiting minimum value equal to DSulLiq/SilLiq when R > 10 × Di; in other words, DSulLiq/SilLiq must be at least as large as Di to produce the necessary R-factor enrichments. Campbell and Barnes (1984) argued that DSulLiq/SilLiq must be at least 105 for the PGEs, a value which is borne out by minimum partition coefficients documented by Brenan (2008) and Mungall and Brenan (2014), and by combining PGE solubility determinations in sulfide and silicate melts (Andrews and Brenan 2002; Fonseca et al. 2009, 2011). In this context, Mungall and Brenan (2014) present a first order model to explain the high Pt concentrations in the sulfides of the Merensky Reef by mixing between primitive and fractionated B1 magmas, in which the fractionated composition is Pt and sulfur-free due to earlier sulfide removal. The Pt content of the primitive B1 magma is taken to be the same as the Bushveld marginal rocks of

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Figure 25. Values of R-factor (see Eqn. 79) as a function of DSulLiq/SilLiq for different values of the apparent partition coefficient, Di (= Csulfide/Ci,silicate), modified from Fonseca et al. (2009). If DSulLiq/SilLiq < R, then Di depends on DSulLiq/SilLiq. If DSulLiq/SilLiq > R, Di depends on R and is independent of DSulLiq/SilLiq. Very large values of DSulLiq/SilLiq for the HSE reported by Mungall and Brenan (2014), Brenan (2008) and Fonseca et al. (2009, 2011, 2012) suggest that extreme enrichments of the HSE in sulfide melt are possible, provided sufficient amounts of silicate melt have equilibrated with the sulfide.

Experimental Fractionation of the HSE

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the Lower and lower Critical zones, which is 15−25 ppb, and only slightly higher than values in most primitive mafic magmas (Barnes et al. 2010). The observed ore concentrations are reproduced with a 50:50 mix of primitive and evolved B1 compositions, and a resulting R-factor of ~ 28,500, which is certainly permissible in the context of current minimum estimates of DSulLiq/SilLiq for Pt. Thus, very high PGE tenors can be achieved at the magmatic stage, obviating the necessity of either enrichment by orthomagmatic fluids (e.g., Meurer and Boudreau 1998) or the involvement of magmas with unusually high PGE concentrations (Naldrett et al. 2009).

Role of the chalcogens (Se, Te, As, Bi, Sb) The relevant phase equilibria for melting and solidification of chalcogen-bearing compositions is reviewed in detail by Frost et al. (2002), Makovicky (2002) and Tomkins et al. (2007), and not repeated here. In terms of the role of chalcogens in affecting the HSE budget of magmatic sulfide systems, key parameters are the identity of the crystallizing phases, their timing of formation and the relative partitioning of elements between chalcogen and base-metal sulfide phases. The existence of S-bearing but arsenic-rich liquids is described by Skinner et al. (1976) who measured an extensive two-liquid field in the system Pd–As–S at 1000 ºC and 0.1 MPa. Helmy et al. (2013b) showed that similar phase relations extend into systems with Fe and Ni, with Pd-As-rich liquids stable to 770 ºC (and possibly below) at 0.1 MPa, with a strong preference of Ni over Fe relative to coexisting sulfide melt. Experiments on the Pt–As–S system (Skinner et al. 1976; Mackovicky et al. 1990, 1992) document extensive solid solution between As–S melts, and that sperrylite is a possible early-formed phase, although the minimum As content of the As–S liquid coexisting with sperrylite at 1000 ºC is quite high (several wt%). Helmy et al. (2013b) showed that the addition of Fe (and Ni) to this system significantly reduces the solubility of sperrylite in the sulfide melt to values ranging from 9400 ppm at 1150 ºC to 6200 ppm at 770 ºC. Similarly, Helmy et al. (2013b) determined that the arsenic content of sulfide melt coexisting with Pd–Ni-rich arsenide melts varies from 37800 ppm at 1150ºC to ~ 400 ppm at 770 ºC. Wood (2003) provided preliminary measurements of the solubility of molten Fe-As-S in basalt melt at 1200ºC and 1 GPa (fO2 < than FMQ buffer), reporting values of 1100 and 3300 ppm, for silicate melts with ~ 19 and ~ 13 wt% FeO, respectively. These values are likely to change with fO2, however, as arsenic is dissolved as a cation in molten silicate (As3+ and As5+; Chen and Jahaanshi 2010; Borisova et al. 2010). In any case, such high As solubility in molten sulfide and silicate would suggest that sperrylite or Pd–Ni–As-bearing melt is not likely to form early in the magmatic sulfide crystallization sequence unless the system has acquired unusually high As levels, likely due to assimilation of sulfidic sediments enriched in organic matter, as proposed by Godel et al. (2012). Alternatively, arsenic levels can become elevated in residual sulfide melts by extensive crystallization of MSS, in which As is sparingly soluble (DMSS/SulfLiq for As varies from 0.4 to 0.01; Liu and Brenan 2015 and Helmy et al. 2010). This, combined with the decrease in arsenic solubility with falling temperature, would promote the formation of Pd–Ni–As melts at a late stage. Gervilla et al. (1994) measured phase equilibria in the system As–Pd–Ni at 790 and 450 ºC, relevant to the Ni-rich assemblages associated with what are interpreted to be natural immiscible arsenide melts. High solubilities of Pd in both nickeline and maucherite are demonstrated in that work, which at 450 ºC is ~ 5 and 9 wt%, respectively, consistent with very high concentrations of Pd (and Pt) reported for these minerals from natural occurrences (Cabri and Laflamme 1976; Cabri 1992; Watkinson and Ohnenstetter 1992). Such high solubilities would indicate these minerals can retain at least a portion of the elevated PGE content of any precursor arsenide liquid. Insight into the solubility and phase relations for the Ir- and Rh-bearing sulfarsenides of the irarsite–hollingsworthite series can be gained by considering the results of dynamic crystallization experiments performed by Sinyakova and Kosyakov (2012). In that work, crystallization of a representative Noril’sk massive sulfide ore composition, doped with PGEs, Au, Ag, and As at the ~ 1000 ppm level of each, was simulated at 0.1 MPa and 1061–821 ºC.

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The sample was encapsulated in an evacuated silica tube, then completely melted, followed by cooling by moving the bottom of the sample from the hot to the cold zone of the furnace over several hours. This time–temperature history resulted in compositional and mineralogical zonation across the sample, from a MSS-rich portion, crystallized initially, to an ISS-rich portion, formed from the residual melt. Various arsenic-bearing phases were also produced over the length of the sample, including euhedral irarsite–hollingsworthite solid solution present as inclusions in early-formed MSS, as well as sperrylite, formed along the MSS–ISS cotectic, then Pt-, Pd- and As-rich segregations with a droplet-like morphology, interpreted to represent immiscible arsenide melts. The relative timing of PGM formation is remarkably similar to that proposed by Dare et al. (2010), in which crystals of irarsite–hollingsworthite, and sparse sperrylite, occur in what was early-formed MSS in contact ore at the Creighton Mine, Sudbury. The euhedral irarsite–hollingworthite inclusions produced in the crystallization experiments were suggested to form by local enrichment of components near the growing MSS crystals. The large MSS–sulfide melt partition coefficients which have been measured for Rh and Ir (see MSS–ISS-sulfide melt partitioning) would cause a depletion of these elements near growing MSS, although DMSS/SulfLiq < 1 for As would suggest that As enrichment is possible. The relatively high levels of As added to experiments, combined with the likely enrichment of this element in the melt during MSS crystallization, suggest that unusually elevated levels of this element are required for irarsite–hollingsworthite saturation in natural sulfide liquids. There is a rather limited database for measurements of the relative partitioning of the HSE between sulfide and arsenide melts, with results from both experiments (Wood 2003; Helmy et al. 2013a) and empirical observations (Hanley 2007; Pina et al. 2013) that indicate preferential uptake by the arsenide phase. Partition coefficients, expressed as the ratio of concentration in the arsenide/sulfide liquids (DAsLiq/SulfLiq) are summarized in Table A3. Pina et al. (2013) provide the most comprehensive dataset, which includes all the HSEs, as measured by reconstructing sulfide and arsenide liquid compositions in samples from the Beni Bousera magmatic Cr–Ni mineralization (Morocco). Values of DAsLiq/SulfLiq estimated in this way are ~ 100 for the PGEs and Au, and 6 for Re. Hanley (2007) estimated relative arsenide–sulfide partitioning by comparing concentrations (on a 100% sulfide basis) between an As-rich high grade lens, with lower grade As-poor segregations, occurring in a series of mineralized komatiite flows. The arsenic-rich high grade lens was found to be enriched by ~ 7−60 × for Pt and Pd, and 2–30 × for Au. The magnitude of DAsLiq/SulfLiq measured in laboratory experiments is similar, with a value for Pd > 34 reported by Wood (2003), and 41 for Pt at 1230 ºC (Helmy et al. 2013b). The partitioning of Pd measured by Helmy was found to be strongly dependent on temperature, increasing from 18 at 1150 ºC to ~ 3500 at 770 ºC. Helmy et al. (2013b) measured the effect of As on the partitioning of Pt between MSS and sulfide melt at the fS2 imposed by the Fe–FeS buffer at 950 ºC and 0.1 MPa. Their results showed a ~ 10-fold decrease in DMSS/SulLiq as the As content of the melt increased from 0 to 40 ppm. The decrease in DMSS/SulLiq was attributed to the formation of Pt–As nanoclusters in the sulfide liquid, as evidenced by the presence of various combinations of nm-sized crystalline PtAs2, and amorphous Pt–As phases, present as inclusions trapped in run-product MSS crystals. Importantly, experiments in which these phases were found were done at arsenic concentrations well below macroscopic saturation in immiscible Pt-arsenide melt or sperrylite. Helmy et al. (2013b) use these results to argue that the behavior of the HSEs in the presence of such ligands as As (and possibly other chalcogens) could be controlled by the surface properties of pre-crystalline nanoclusters. For comparison, Fleet et al. (1993) measured MSS–sulfide melt partitioning in experiments doped with up to 3500 ppm As, 4600 ppm Bi and 3300 ppm Te, compared to 40−50 ppm of the PGEs, in which DMSS/SulLiq was found to be identical to equivalent chalcogen-free experiments. A similar result was found by Liu and Brenan (2015) who measured DMSS/SulLiq involving sulfide melt with < 0.1 ppm to 150 ppm As (and approximately similar

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contents of other chalcogens) with no apparent effect on Pt or other HSE partitioning. A notable difference between studies is sulfur fugacity; whereas Fleet et al. (1993) used sulfur-excess starting materials, as did Liu and Brenan (2015), who performed experiments at log fS2 in the range of −2 to −2.6 (i.e., near values for natural silicate magmas), the fS2 of experiments done by Helmy et al. (2013b) was buffered at much lower values near Fe–FeS (approximate log fS2 of −7). Assuming both As- and S-related Pt species in the sulfide liquid, whose relative abundances are expressed as a homogeneous exchange reaction of the form: PtAsx + y/2S2 = PtSy + x/2As2

(80)

indicates that an increase in the sulfur fugacity will shift the equilibrium to the right, promoting the formation of the Pt–S species. Therefore, it seems plausible that the general lack of any effect of arsenic on Pt (or other HSE) partitioning observed at high fS2 indicates that the Pt–As species existing at low fS2 has been consumed by Reaction (80). Results would therefore suggest that arsenic is not an important complexing agent at the much higher fS2 required to stabilize sulfide liquid in an FeO-bearing silicate magma, due to the effect of fS2 on As speciation. In terms of the effect of other chalcogens on HSE behavior, Helmy et al. (2007) measured the high temperature phase equilibria and partitioning in the Fe–Cu–Ni–Pd–Pt–Te–S system. These experiments tracked phase compositions from 1150 to 370 ºC at 0.1 MPa, in both metal rich and Te-rich compositions. Experiments show that a Te-rich immiscible melt forms at 920–700 ºC, which highly concentrates Pd and Pt; moncheite (PtTe2) is the hightemperature PGM in this system, initially crystallizing at 920 ºC. These results are relevant to the behavior of Te in orogenic and ophiolitic peridotites, in which Te resides not only in base-metal sulfides, but also in trace amounts of an accessory Pt–Pd–Te–Bi phase (moncheite– merenskyite, PtTe2–PdTe2; e.g., Lorand et al. 2010; König et al. 2014). Notably, Te levels in natural peridotite-hosted sulfides are well below the concentration levels of ~ 3000 ppm required for PtTe2 saturation, as reported by Helmy et al. (2007). Hence, it seems that the observed PtTe2 is likely a subsolidus phase, exsolved from the Te-bearing base-metal sulfide during the protracted cooling history of the peridotite body (Lorand et al. 2010).

SILICATE MELT–AQUEOUS LIQUID–VAPOR PARTITIONING Silicate melts generally reach saturation in a volatile phase due to decompression and crystallization during their ascent and storage in crustal reservoirs. This process is particularly significant at convergent plate margins where volatile elements such as H, C, Cl, and S may be recycled from the subducting slab into the mantle wedge, thus the generated magmas are enriched in volatile components. The exsolving magmatic volatile phase (MVP) is known to extract significant amounts of metals and may contribute to the formation of magmatic–hydrothermal ore deposits (Hedenquist and Lowenstern 1994; Williams-Jones and Heinrich 2005; Richards 2011). The most notable HSE mobilized by the MVP is gold. Gold commonly occurs in porphyry-type ore deposits along with Cu, as well as in both high- and low-sulfidation epithermal environments (Muller and Groves 1993; Richards and Kerrich 1993; Sillitoe 2002, 2008; Heinrich et al. 2004). The gold budget of most of these deposits is thought to be dominantly derived from underlying upper crustal magma reservoirs. Palladium and Pt are also enriched in some porphyry-type deposits (Eliopoulos and Economoueliopoulos 1991; Tarkian and Stribrny 1999; Berzina and Korobeinikov 2007). Their concentration shows a positive correlation with that of Au with several notable exceptions. The presence of these metals in porphyry ores, however, indicates their partitioning into the MVP at certain conditions, which may be similar to those preferential for gold extraction. In addition, there is evidence from some occurrences that Pd and Pt have been re-mobilized by late magmatic brines in compacting cumulates or by hydrothermal systems associated with magmatic-sulfide ore

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deposits (Boudreau et al. 1986; Boudreau and McCallum 1992; Molnar et al. 1997; Hanley et al. 2005a, 2008). Porphyry molybdenum deposits are the primary resource of rhenium, which commonly occurs as a substituting cation in molybdenite (Berzina et al. 2005; Berzina and Korobeinikov 2007; Grabezhev 2013). It is likely that Re is extracted along with molybdenum by the MVP from underlying magma reservoirs at certain physical–chemical conditions and the two metals are later co-precipitated as a sulfide phase. Gold, Pd, Pt, Ir, Os, and Re have also been found in significant concentrations in high-temperature magmatic gases sampled from vents on the surface, further signifying the mobility of these metals in the MVP (Naughton et al. 1976; Zoller et al. 1983; Toutain and Meyer 1989; Crocket 2000; Williams-Jones and Heinrich 2005; Yudovskaya et al. 2008). To our knowledge, elements within the IPGE group do not show notable enrichment in magmatic-hydrothermal ore deposits.

Theoretical considerations Metals are commonly dissolved in the MVP associated with various ligands, most notably Cl−, HS−, OH−, H2So, HClo and potentially SO42− (Candela and Piccoli 1995; Barnes 1997; Williams-Jones and Heinrich 2005). Therefore, the concentration of these ligands in the MVP significantly impacts their volatile/melt partition coefficients. The affinities of cation–ligand pairs are highly variable. A simple way to predict relative complex stabilities is the application of the Hard Soft Acid Base (HSAB) theory (Pearson 1968). The HSAB theory distinguishes Lewis acids (electron acceptors, usually metal cations) and Lewis bases (electron donors, “ligands”). Hard acids and bases have high electronegativity, low polarizability and generally high valence state, whereas soft acids and soft bases have low electronegativity, high polarizability and low valence state. Therefore, cations with small ionic radii and high positive charge behave as hard acids, whereas cations with large ionic radii and low positive charge are typically soft acids. Similar statements stand for bases except the charge is negative. Out of those occuring in significant concentrations in the MVP: OH−, Cl− and SO42− are hard, whereas HS− and H2So are soft bases. Considering HSE, Au and PPGE metals form ions that behave as soft acids (i.e., Au+, Pt2+ and Pd2+) whereas even the lowest oxidation states of IPGE ions (i.e., Ru2+, Os2+, Rh2+, Ir2+) are harder and classified as borderline acids (Pearson 1968; Parr and Pearson 1983). As hard acids prefer to bind to hard bases and soft acids prefer to bind to soft bases, a general tendency in the volatile/melt partitioning of HSE can be foreseen simply by applying the HSAB principle. Accordingly, Au, Pt, and Pd are most likely to strongly partition into reduced sulfur-bearing volatiles, whereas other HSE may prefer chloride-rich brines. As O2− can be considered as a hard base and metals in the silicate melt are generally coordinated by oxygen, one may speculate that IPGE and Re will also show less tendency to partition into the volatile phase in general. Although Cl and S may partitioning strongly into magmatic vapor, the concentrations of these elements in natural silicate melts may still remain relatively high, reaching levels of up to a few thousands of ppm (Wallace 2005). The concentrations of reduced S and Cl has been shown to significantly affect the solubility of Au and PGE in silicate melts (Botcharnikov et al. 2010; Jego and Pichavant 2012; Zajacz et al. 2012b; Mungall and Brenan 2014), and must therefore be incorporated in parameterizations attempting to model their partitioning. The volatile/melt partition coefficients of Cl and S depend on P, T and melt composition,further adding to the complexity of predicting volatile/melt partition coefficients for HSE in P–T–X space (Webster and De Vivo 2002; Lesne et al. 2011; Zajacz et al. 2013). An additional important variable that may affect the volatile/melt partitioning of HSE is the density of the MVP. Metal complexes are stabilized in solution by the formation of hydration shells around them, and the extent of hydration shell formation and associated drop in the Gibbs free energy of the system is primarily controlled by the dielectric constant of water (Helgeson et al. 1981; Barnes 1997), which in turn positively correlates with the density of the MVP (Pitzer 1983). Therefore, the volatile/melt partition coefficients of most metals can be expected to

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positively correlate with the density of the MVP provided that the activity of various ligands is kept constant. This is manifested in the much higher metal concentrations determined in high-T fluid inclusions (FI) in magmatic–hydrothermal quartz veins formed under confining pressure at upper crustal depths (Ulrich et al. 1999; Klemm et al. 2007; Audétat et al. 2008; Zajacz et al. 2008; Landtwing et al. 2010), than those typically measured in high-T volcanic gases (Taran et al. 1995; Williams-Jones and Heinrich 2005; Wardell et al. 2008; Yudovskaya et al. 2008). As the above discussion illustrates, the construction of a thermodynamic model to predict the volatile/silicate melt partition coefficients of HSE requires complicated thermodynamic modeling relying on a wealth of experimental data, which is still missing for most HSE. In the following sections we discuss the most important experimental data produced in the past on Divolatile/SilLiq for these elements.

Experimental methods Most volatile/melt partitioning experiments on HSE were conducted in externally heated cold-seal pressure vessel apparatus (CSPV), which is capable to reproduce the P–T conditions of magma reservoirs in the upper crust. Such studies commonly used Ni-based superalloy vessels (e.g,. Rene 41) which limit the maximum experimental temperatures to ~ 800 ºC (Frank et al. 2002; Hanley et al. 2005b; Simon et al. 2007). Some of the later work has employed Titanium Zirconium Molybdenum (TZM) or Molybdenum–Hafnium Carbide (MHC) alloy vessels, allowing experiment temperatures up to 1000 ºC (Zajacz et al. 2010, 2012b). A typical starting phase assemblage for volatile/melt partitioning experiments consists of a silicate glass, an aqueous solution containing various ligands and sometimes a mineral assemblage to buffer important intensive variables such as pH, fO2, and fS2. Additional salts or sulfur may be added as solids. The starting phase assemblage is sealed into a noble-metal capsule that is inert with respect to the capsule load and ductile enough to transmit pressure. Most typically, the capsule itself is the source of the HSE in the experiment to avoid problems with alloying between the capsule and the metal under study. Similar to methods described in the section Experimental approach, the control of fO2 is usually achieved by external buffering, as the often used Au, Pt, and AuPd alloys are permeable to hydrogen at magmatic temperatures (Chou 1987), and the experimental charge is generally at high water activity (aH2O) due to the presence of an aqueous volatile phase. Oxygen fugacity is therefore controlled through the dissociation of water in the capsule. Hydrogen fugacities are either established using a double capsule technique with a redox buffer assemblage in the external capsule (e.g., Ni–NiO, Re–ReO2) or by establishing a constant hydrogen fugacity in the pressure medium itself. This is conveniently done in Ni-based superalloy vessels by using water as pressure medium, as hydrogen fugacity is buffered by the reaction: Ni(s) + H2O(g) = NiO(s) + H2(g)

(81)

The activity of Ni in these alloys is somewhat below one, therefore fH2 is slightly lower than that corresponding to the equilibrium constant of Equation (81). Due to the ease of fO2 control and the widespread availability of superalloy CSPV facilities, many of the previous volatile/melt partitioning experiments were conducted in such apparatus. Tattitch et al. (2014) determined that Rene 41 vessels impose fH2 corresponding to fO2 of about NNO + 0.1 (FMQ + 0.8) at aH2O = 1 and T = 800 ºC by using the CoPd alloy–CoO redox sensor (Taylor et al. 1992). Others imposed fH2 by mixing H2 into the Ar pressure medium in Mo-based alloy vessels (TZM and MHC), and determined that fH2 decreases only 0.3−0.4 log units in the vessel in 24 h at 1000 ºC due to diffusive loss of H2 through the vessel walls (Zajacz et al. 2010, 2012b, 2013). Note that this value is very much dependent on the pressurized gas volume/heated vessel surface area, therefore, the relative change in fH2 will be much faster when filler rods are used (Shea and Hammer 2013).

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The HCl/total chloride ratio is another important variable that significantly affects the solubilities of metals in the volatile phase, and therefore their volatile/melt partitioning. The concentration of HCl has been determined by mass balance (Frank et al. 2002), pH measurement in the quench fluids (Simon et al. 2005), or by using the formula of Williams et al. (1997) to calculate it based on the aluminium saturation index (ASI) of the silicate melt in equilibrium with the volatile phase (Simon et al. 2005; Simon and Pettke 2009). Hanley et al. (2005b) used an albite–andalusite–quartz assemblage to buffer fHCl, via the reaction: NaAlSi3O8(s) + HCl0(aq) = 0.5 Al2SiO5(s) + 2.5 SiO2(s) + NaCl0(aq) + H2O(l)

(82)

Similar to a method employed to determine DSulLiq/SilLiq (see the section Sulfide-melt/silicate melt and MSS–silicate melt partitioning), Zajacz et al. (2010a, 2012a,b) conducted metal solubility experiments in the volatile phase and the silicate melt separately at T = 1000 ºC to be able to model volatile/melt partition coefficients for Au and Cu in mafic–intermediate systems. In the absence of a silicate melt phase, the HCl concentrations in the volatile phase could be directly imposed by the starting fluid composition because no cation exchange reactions could occur during the run that could consume or produce HCl. The exception are FeCl2 bearing experiments, where Fe diffusion into the Au capsule produces HCl; however, as all metal concentrations were measured by LA-ICPMS in the synthetic fluid inclusions, the equilibrium HCl concentrations could be calculated even in those experiments (Zajacz et al. 2011). In S-bearing experiments, it is of critical importance to understand the concentration, oxidation state and speciation of S in the volatile phase and the silicate melt, as various sulfur species can have contrasting properties in terms of metal complexation (e.g., HS− is a soft base, whereas SO42− is a hard base). Simon et al. (2007) added S in the form of arsenopyrite and determined the concentration of S in the volatile phase using the measured S concentrations in the silicate glass and the solubility model of Clemente et al. (2004). Zajacz et al. (2010, 2011 2012a,b) loaded elemental S into the capsules and used mass-balance constraints to determine S concentrations in the volatile phase in multiphase experiments. Most commonly H2S and SO2 are assumed to be the most dominant S-species in magmatic volatiles, due to the reduced tendency of neutral molecular species to dissociate because of the low dielectric constant of water at these P-T conditions. The H2S/SO2 ratios are often predicted using gas phase standard state thermochemical data to calculate the equilibrium constant of the following reaction: H2S(g) + 1.5O2(g) = SO2(g) + H2O(g)

(83)

However, Binder and Keppler (2011) and Ni and Keppler (2012) have shown that H2SO4 species may also be stabilized in relatively high-density magmatic fluids at fO2 ≥ Re–ReO2 (FMQ + 2.6). Furthermore, Jacquemet et al. (2014) suggested that trisulfur ions (S3−) may be present in moderately S-rich (~ 3 wt% S) aqueous fluids at magmatic temperatures based on extrapolation of in situ Raman spectroscopic data obtained at 25−500 ºC. The oxidation state of S in the silicate melt may be inferred from the value of fO2 using previous calibrations (Jugo et al. 2010; Klimm et al. 2012a), or directly determined using the S-Ka line wavelength shift measured by EPMA or by monitoring changes in S K pre-edge spectral features by XANES (Wallace and Carmichael 1994; Klimm et al. 2012b). One of the main challenges in volatile/melt partitioning experiments is posed by the often unquenchable nature of the volatile phase. Early studies on volatile/melt partitioning used mostly Cl-bearing, S-free fluids and focused on metals which remained in solution upon quenching. The quenched fluids were extracted from the capsule and analyzed by various bulk methods (e.g., mass spectrometry, atomic absorption spectrometry; Holland 1972; Candela and Holland 1984). The applicability of this technique is, however, questionable for the HSE, as experiments involving these elements are generally conducted in the presence of the pure metal phase, which generally is

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the capsule itself. As the solubility of these metals at high P–T is typically higher than at ambient conditions, one may assume that a part of the dissolved metal could precipitate on the already present metallic phase upon quenching, and therefore cannot be representatively recovered for solution-based analysis. Consequently, the fluid phase has to be sampled at the experimental P–T and analyzed later, or it has to be directly analyzed at run conditions by in situ spectroscopic methodologies. Technological developments made in recent years allow in situ determination of trace element concentrations in high P–T fluids by X-ray Absorption Spectroscopy (Testemale et al. 2005), and abundant data were obtained on Au (Pokrovski et al. 2009a,b). Nevertheless, experimental difficulties arising at magmatic temperatures have so far limited the applicability of these techniques to lower-T, hydrothermal conditions. The development of the synthetic fluid inclusion technique paralleled by the rapid evolution of more and more sensitive, high spatial resolution, in situ but non-surface analytical methodologies such as LA-ICP-MS and protoninduced X-ray emission (PIXE) opened new pathways for the investigation HSE solubilities in magmatic volatiles. The volatile phase can be trapped at the experimental run P–T in the form of synthetic fluid inclusions (SFI) in quartz (Bodnar and Sterner 1987), and the bulk composition of the SFI can subsequently be analyzed by the above analytical techniques. The first such study used PIXE to determine the concentration of Au and Pt amongst other elements (Ballhaus et al. 1994). This was followed by several others after a new methodology had been established to analyze FI in quartz by the more widely available LA-ICP-MS (Gunther et al. 1998; Heinrich et al. 2003). This technique is capable to detect HSE concentrations down to sub-ppm levels in FI as small as 20−40 µm in diameter. Frank et al. (2002) used an alternative methodology to study the partitioning of Au between a rhyolite melt and a high-salinity aqueous liquid phase (brine) taking advantage of the fact that microscopic sized (~ 0–30 µm) bubbles containing the brine phase do not physically separate from the rhyolite melt due to its high viscosity on the experimental time scale. The bulk composition of glasses containing varying mass fraction of brine inclusions was obtained by instrumental neutron activation analysis (INAA), whereas the clean glass was analyzed by in situ methods (EPMA, SIMS). The composition of the brine was then derived by mass balance. Others working on felsic systems also analyzed FI in the silicate glass by laser ablation ICP-MS in addition to, or instead of SFI in quartz (Hanley et al. 2005b; Simon et al. 2007; Simon and Pettke 2009). Another challenging aspect of volatile/melt partitioning experiments on HSE is the likely occurrence of HSE metal nuggets in the silicate melt and in the volatile phase. Nuggets of Au and PGE commonly occur in the silicate melt and have been discussed in detail in Possible mechanisms of metal inclusion formation. Gold nuggets in the silicate glass were observed in many volatile/melt partitioning studies (Frank et al. 2002; Hanley et al. 2005b; Simon et al. 2005, 2007; Zajacz et al. 2012b). Some studies interpreted these to be quench products and integrated them into the glass composition (Frank et al. 2002; Simon et al. 2005, 2007), whereas others considered them as stable phase at run conditions and removed them from the signal (Hanley et al. 2005b; Zajacz et al. 2012b). Zajacz et al. (2012b) repeated some of the most nugget-rich experiments after adding a small amount of oxalic acid to the starting solution which secured low fO2 at the beginning of the run before equilibrium with the apparent fH2 in the pressure medium has been reached. This effectively eliminated the occurrence of Au nuggets, and the such-determined Au concentrations in the glass were equal to those obtained from the identical, but oxalic acid-free experiment by eliminating Au nuggets from the LA-signal. This observation strongly suggests that Au nuggets are indeed stable phases in hydrous silicate melts at run conditions, and should be excluded from the glass composition. This result is consistent with the model described in Possible mechanisms of metal inclusion formation, that there is a time evolution in the fO2 of metal solubility experiments—generally from high to low—so the elimination of the high fO2 portion prevents initial HSE dissolution, then subsequent precipitation of metal particles as the fO2 decreases to the equilibrium value.

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Though Au nuggets in the volatile phase are likely non-existent or insignificant based on the low standard deviation of Au concentrations in SFI, the situation is less clear in the case of Pt. All previous studies found highly scattering concentrations of Pt in synthetic FI from individual experiments (Ballhaus et al. 1994; Hanley et al. 2005b; Simon and Pettke 2009). Ballhaus et al. (1994) argued that this may relate to the heterogeneous entrapment of polynuclear complexes; however, these would need to be extremely large to induce the observed variation in Pt concentrations. Simon and Pettke (2009) speculated that the erratic concentrations may relate to the re-equilibration between the fluid inclusions and the host silicate glass upon quenching. Hanley at al. (2005b) noted a positive correlation between the formation time of the FI and the measured Pt concentrations, and suggested that the range in Pt concentrations is due to premature fracture healing before equilibrium between the brine and the Pt metal has been reached. However, even for a specific entrapment time interval, about an order of magnitude scatter was observed in Pt concentrations. This may suggest the heterogeneous entrapment of nuggets. In addition, Hanley et al. (2005b) proposed that the time dependent concentrations may relate to the re-precipitation of Pt in the unhealed fractures under a temperature gradient.

The volatile/melt partitioning of Au Due to its pronounced presence in magmatic-hydrothermal ore deposits, Au has received by far the most attention out of all HSE with respect to volatile/melt partitioning behavior. A summary of the data is provided in Table A5 and displayed in Figure 26. Most studies investigated felsic melts in equilibrium with chloride-bearing magmatic volatiles in Rene-41 vessels at fO2 = FMQ + 0.7. Only Simon et al. (2007) also added sulfur to the system. As all these studies were conducted in Au capsules, the measured concentrations in the silicate glass and the volatile phase represent the solubility of Au and are shown separately. Zajacz et al. (2010) and Zajacz et al. (2011) studied the solubility of Au in S- and Cl-bearing low-density vapors at 1000 ºC, and later combined these data with Au solubilities obtained in andesite melts and S, Cl partitioning, volatile/SilLiq and Cl speciation data to develop a model to calculate DAu (Zajacz et al. 2012a). In S-free systems, a common conclusion of Au solubility studies is that the concentration of Cl and HCl/total Cl ratios significantly affect the Au content of the volatile phase and therefore. Simon et al. (2005) noted a strong preferential partitioning of Au into the brine phase relative to the vapor, and Zajacz et al. (2010) also found a positive correlation between the concentration of chloride species in the vapor phase and the solubility of Au. On the other hand, Hanley et al. (2005b) observed a strong negative correlation between Au solubilities and total chloride concentration in brines and partially attributed this to the salting-out effect, wherein stronger dipole NaCl ion pairs successfully compete against weaker-dipole neutral Au chloride complexes for the limited number of water molecules to form hydration shells around them. Considering that Hanley et al. (2005b) covered the 4.3–39.9 m total-chloride concentration range, whereas vapors in Simon et al. (2005) and Zajacz et al. (2010) were characterized by mCltotal < 1.5, it may be possible that there is a maximum in Au solubility at intermediate chloride concentrations. However, this requires further experimental confirmation. It can also be noted that the Au solubility values of Hanley et al. (2005b) are typically much higher than those of Simon et al. (2005) at similar salinity and T. A possible explanation for this difference would be if the HCl/NaCl ratios in Hanley’s experiments were higher. The values reported by the authors are against this hypothesis, however, one must note that very different methodologies were used to estimate HCl concentrations in these two studies. The silicate melt present at 800 ºC in Hanley’s experiments is highly peraluminous, which would be in equilibrium with an HCl-rich fluid according to Williams et al. (1997), whereas the melt in Simon’s experiments is slightly peralkaline. The solubility values of Hanley at al. (2005b) fall into the same range with the most HCl-rich fluids of Frank et al. (2002) at identical T, further supporting this hypothesis. Considering the speciation of Au in chloride-bearing volatiles, Frank et al. (2002) proposed AuCl and HAuCl2 as the dominant species at low and high HCl concentrations, respectively, whereas Zajacz et al. (2010) argued for the dominance of AuCl and NaAuCl2 even in HCl-rich fluids.

Experimental Fractionation of the HSE

a

Frank et al. (2002)

Hanley et al. (2005)

Simon et. al. (2005)

Zajacz et al. (2010)

Simon et. al. (2007)

Zajacz et al. (2010) - S-bearing

Simon et. al. (2007) - S-bearing

o

700 C

cAu (ppm)

600 oC

800 oC

o

700 C

600 oC

o

700 C

o

600 C

mCltotal

cAu (ppm)

b

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mHCl/mCltotal

Figure 26. (a) The solubility of Au in magmatic volatiles as a function of total dissolved chloride concentration. All data from Table A5. Most data is from experiments done at 800 ºC and fO2 = FMQ + 0.7, except those of Zajacz et al. (2010), which are at 1000 ºC and FMQ + 0.2. Sulfur-bearing experiments are represented by orange half-filled symbols to be easily distinguishable from S-free experiments. Some experiments of Hanley et al. (2005b) were conducted at different temperatures, which are shown as labels next to the symbols. Data of Zajacz et al. (2010) with dissolved LiCl, KCl, and CaCl2 are not shown. (b) The same data shown as a function of mHCl/mCltotal. Note that some data of Zajacz et al. (2010) are not shown because they are nominally at mHCl = 0.

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The gold solubility data in S-free silicate melts are also somewhat controversial, which can likely be attributed to different handling of Au nuggets. Studies that integrated Au nuggets into the glass composition obtained Au solubilities for haplogranite melts with ASI of 0.9– 1.1 ranging from 0.5 to 4.7 ppm (Frank et al. 2002; Simon et al. 2005, 2007). Therefore, the volatile/SilLiq values are around 10−20 for vapors, in the high tens for low-HCl brines obtained DAu and in the hundreds to low-thousands for high-HCl brines (Table 6). Hanley et al. (2005b) and Zajacz et al. (2013) chose to exclude Au nuggets from the integrated signal, whereas Zajacz et al. (2012b) effectively eliminated nugget formation. Yet, the Au solubility determined by Hanley et al. (2005b) in strongly peraluminous haplogranite melt (ASI = 1.33−1.49) is 0.82 ppm as opposed to the 0.02 ppm reported by Zajacz et al. (2013) for less peraluminous (ASI =1.02), Cl-free granite melts. This apparent inconsistency may partially be resolved by correcting for the difference in fO2, and by noting the difference in the ASI and dissolved Cl concentrations in the silicate melts in these two studies. For example Zajacz et al. (2012b) showed positive correlation between Cl concentration and Au solubility in andesite melts, whereas Zajacz et al. (2013) showed that Cu and Ag solubilities increase rapidly with ASI in peraluminous melts, which may also apply to Au. Nevertheless, Au solubilities at T = 1000 ºC in hydrous andesite melts with 1 wt% dissolved Cl, determined on nugget-free glasses, are only about 0.2 ppm at FMQ + 0.2 (Zajacz et al. 2012b). This would scale to about 0.27 ppm at FMQ + 0.7. As the solubility of Au in metaluminous rhyolite melts at 800 ºC would likely be lower (Zajacz et al. 2013), it is probable that Au concentrations reported with inclusion of nuggets give an upper estimate for volatile/SilLiq values the true Au solubility in silicate melts. If this statement is valid, the reported DAu in these studies are minimum estimates; however, this needs further experimental confirmation. Considering S-bearing systems, to our knowledge, only four studies investigated the solubility of Au in S-bearing aqueous fluids at magmatic temperatures (Loucks and Mavrogenes 1999; Simon et al. 2007; Zajacz et al. 2010, 2011), and only one of these included silicate melt in the same experiment (Simon et al. 2007). Loucks and Mavrogenes (1999) determined Au solubility values of 30–1180 ppm in aqueous volatiles at 625–725 ºC and in the presence of magnetite-pyrrhotite and magnetite-pyrrhotite-pyrite assemblages constraining fO2 and fS2. Based on the relationship between the calculated fH2S and the measured Au solubility as determined over a fairly narrow range of fH2S (22−46 bars), they proposed a 4-coordinated AuHS(H2S)3 complex in the fluid. Later studies could not identify the same complex in aqueous hydrothermal solutions or magmatic volatiles (Pokrovski et al. 2009b; Williams-Jones et al. 2009; Zajacz et al. 2010). Experiments at hydrothermal conditions and those of Zajacz et al. (2010) at 1000 ºC proposed S/Au ratios of 1 to 2 in the stable Au complexes (e.g., AuHS, Au(HS)2−, AuHSH2S). In addition, Zajacz et al. (2010) observed that the addition of even small concentrations of alkalichlorides to H2O–H2S vapors increases the solubility of gold by about an order of magnitude in low-density vapors at 1000 ºC. This effect was the most pronounced with KCl, somewhat weaker with NaCl, and minor with LiCl. Feasible explanation may be the presence of mixed alkali–gold–hydrosulfide (e.g., KAu(HS)2) or alkali–gold–chloride–hydrosulfide (NaAuClHS) complexes in such low-dielectric fluids. The stability of such complexes was later predicted by ab initio static quantum chemistry calculations and molecular dynamics simulations as well for both Au and Cu (Zajacz et al. 2011; Mei et al. 2014). The solubility of Au in hydrous silicate melts shows a strong positive linear correlation with the concentration of dissolved reduced sulfur and this needs to be accounted for when modeling volatile/melt partition coefficients (Botcharnikov et al. 2010; Zajacz et al. 2012b). For example, in andesite melts at fO2 = FMQ − 0.1, Au solubility increases from 0.06 ppm to 1.27 ppm when the melt composition is changed from S-free to pyrrhotite saturated with 354 ppm dissolved S (Zajacz et al. 2013). This effect is even more pronounced in the presence of dissolved Cl and in peralkaline melts, but more subtle in peraluminous rhyolites (Zajacz et al. 2012b, 2013). Interestingly, Simon et al. (2007) found the solubility of Au to decrease

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with the addition of S to the silicate melt, however, this may relate to artificially higher Au concentrations in S-free experiments due to more extensive nugget formation in the glass (Zajacz et al. 2012b). The Au solubilities determined for S-bearing rhyolites by Simon et al. (2007) (0.18−1.10 ppm) still exceed the value of Zajacz et al. (2013) for similar slightly peraluminous rhyolite melt at 800 ºC (0.1 ppm corrected for difference in fO2). Zajacz et al. (2012b) modeled the partition coefficients of Au between andesite melts and magmatic volatiles using Au solubility data in both phases as well as S and Cl partition volatile/SilLiq reaches values coefficients determined in the same system. The results showed DAu around 200 in pyrrhotite saturated systems, but are significantly lower (~ 20) in S-free systems. volatile/SilLiq in S-free andesites is mainly due to the weak partitioning of Cl into the The lower DAu volatile/SilLiq of 11–50 volatile phase from intermediate melts. Simon et al. (2007) determined DAu for reduced S-bearing slightly peraluminous rhyolites, which may be treated as minimum values because Au nuggets were included in the glass analysis. All in all, we conclude that exsolving volatiles will most efficiently extract Au from mafic to felsic silicate melts if significant reduced S is present in the system (i.e., at or slightly below pyrrhotite saturation). However, even in the absence of reduced S-species, gold extraction from felsic melts can be highly efficient as the fluids derived from such evolved compositions are typically chloride-rich. This may further be promoted in peraluminous systems, as HCl/metal chloride ratios will be high, and therefore Au solubilities in the volatile phase will increase with increasing ASI of the melt.

The volatile/melt partitioning of PPGE To our knowledge only two studies investigated directly the volatile/melt partitioning behavior of Pt and data on other PPGE are still lacking. Simon and Pettke (2009) conducted experiments at T = 800 ºC and P = 100 and 140 MPa in a rhyolite melt–vapor– brine assemblage. The reported average Pt solubility values in the vapor and brine are 0.15–1.67 ppm, and 2.33–45 ppm, respectively. The fluids were HCl-rich corresponding to the peraluminous nature of the silicate melt (ASI = 1.19–1.31). Platinum nuggets were observed in the silicate glass, but their contribution to the laser ablation signal was judged insignificant. The so-determined Pt solubilities in the melt ranged over 0.17–0.53 ppm. The corresponding vapor/melt partition coefficients were between 0.88 and 6.0, whereas the brine/melt partition coefficients were between 6.7 and 149. Hanley et al. (2005b) found a negative correlation between the total chloride concentration and the solubility of Pt in supercritical brines, with Pt solubilities dropping from ~ 1700 ppm to ~ 20 ppm as mCltotal was increased from 4.3 to 39.9. This is similar to the effect observed for Au in the same study, however, the Pt concentrations are less consistent and scatter over a range of 1–2 orders of magnitude from individual FI in the same experiments. Platinum solubilities were reported in two haplogranite melts, one with an ASI of 1.33 contained 0.115 ± 0.078 (1σ) ppm, whereas the other one with an ASI of 1.49 contained 0.035 ± 0.018 (1σ) ppm Pt. The corresponding DPtvolatile/SilLiq values are 21000 ± 14000 (1σ) and 3300 ± 5700 (1σ), respectively. These are orders of magnitude higher than those of Simon et al. (2009). As the P, T, silicate melt and fluid compositions in these two studies are rather similar; there is no simple explanation for this inconsistency. It is apparent that the fluid inclusions in Hanley et al. (2005b) are much more Pt-rich, whereas the silicate glass is lower in Pt. The latter may be due to the fact that Hanley et al. did not integrate the Pt nuggets into the laser ablation signal, whereas the more Pt-rich fluids of Hanley et al. (2005b) could be an artifact of heterogeneous entrapment of Pt nuggets in the FI or Pt re-precipitation under a temperature gradient as discussed by the authors. Similar processes may have played a role in the experiments of Ballhaus et al. (1994), who measured erratic Pt concentrations of up to several wt% in SFI sampling brines in the presence or absence of sulfides at 900 ºC and 1 GPa. These experiments did not contain silicate melt.

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Unfortunately, to our knowledge there is no direct solubility or partitioning data published for other PPGE at magmatic temperatures. However, Fleet and Wu (1993, 1995) assessed the relative mobility of PGE in low-density vapors at 1000 ºC by monitoring the efficiency of metal transport along temperature gradients in vacuum-sealed silica tubes in the presence of various ligands. They found that PGE, in particular Pt and Pd were most mobile in the simultaneous presence of sulfide and NaCl species in water-free vapors, but less mobile in sulfide-bearing but chloride-free systems. This observation is consistent with that of Zajacz et al. (2010) on Au in low-density aqueous vapors at the same temperature. Fleet and Wu (1993) also pointed out that the mobility of PGE in the vapor is significantly reduced in the presence of pnictogens and chalcogens that form refractory compounds with them. For example, Pt mobility was reduced in the presence of As, due to the formation of refractory PtAs2. In addition, numerous studies were conducted on the solubility of Pd and Pt at hydrothermal conditions, mostly in the temperature range of 100–500 ºC. The detailed discussion of these studies is beyond the scope of this review and the reader is referred to Wood (2002). However, in brief summary, these studies agreed that Pd and Pt are most likely transported in the form of chloride complexes in strongly acidic and oxidizing hydrothermal solutions, whereas bisulfide complexes dominate under slightly acidic to neutral and reducing conditions. The equilibrium constant of the dissolution reactions to produce chloride complexes is decreasing with increasing temperature, however, if the fO2 of the system is tied to one of the common geologic redox buffer assemblages a net solubility increase is expected with increasing T due to parallel increase in fO2. The stability of Pt and Pd bisulfide complexes were found to have a maximum at around 150 ºC. These characteristics are in qualitative agreement with those observed for Au at hydrothermal temperatures (Williams-Jones et al. 2009), however, the equilibrium constants of the dissolution reactions for Pt and Pd are generally much smaller than those for Au, with Pt being slightly more soluble than Pd (Wood et al. 1992; Pan and Wood 1994). The predicted solubilities at most typical geological conditions at 300 ºC are in the sub-ppb range for both metals (Wood et al. 1992, 1994; Gammons and Bloom 1993; Gammons 1995, 1996), which is several orders of magnitude smaller than the values proposed for Pt at magmatic temperatures as discussed above. The highest temperatures were reached by Xiong and Wood (2000) and Hsu et al. (1991) using hydrothermal autoclaves to study Pd solubility. Xiong and Wood (2000) reported Pd solubilities of 40 ppb at T = 500 ºC, P = 55 MPa, a pH of ~ 5.5 and fO2 = Re–ReO2 (FMQ + 3.2) in 0.1 m KCl solution. Interestingly, Hsu et al. (1991) determined a much higher Pd solubility of 6.5 ppm, at lower fO2 (FMQ + 0.4) and slightly higher pH (6.5) and pressure (100 MPa), but otherwise identical conditions. Hsu et al. (1991) also identified retrograde Pd solubility with increasing temperature, which was still as high as 12.5 ppm at 700 ºC in 3 m NaCl solution at a pH of 6.5. The rather high solubilities determined by Hsu et al. (1991) were proposed to be an experimental artifact by Wood and Mountain (1991).

The volatile/melt partitioning of IPGE and Re Out of all HSE, the least is known about IPGE solubilities in magmatic volatiles. To our knowledge no volatile/melt partitioning experiments were conducted on these metals. The experiments of Fleet and Wu (1993) included Os and Ir in the set of PGE investigated. These metals showed approximately identical mobility in the silica tubes, which was however, lower than that of Pt and Pd by about a factor of 40. Xiong and Wood (2000) studied the solubility of Os in KCl-bearing hydrothermal fluids at 400 and 500 ºC, 80 MPa and pH corresponding to the K-feldspar–quartz–muscovite buffer. They determined Os solubilities of up to about 1.7 ppm in 1.5 m aqueous KCl solution at 500 ºC and fO2 = Re–ReO2. volatile/SilLiq of 7.5–356 between haplobasalt melts MacKenzie and Canil (2011) determined DRe and chloride-rich (5−30 M CaCl2 + MgCl2) aqueous fluids at T = 1300–1400 ºC and P = 1 GPa; however, no clear systematics were observed as a function of fluid composition. Amongst other elements, Johnson and Canil (2011) investigated the kinetics of the degassing of Re and Au

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from silicate melts at T = 1200–1430 ºC and atmospheric pressure. Gold was the most efficiently degassed element with a log diffusivity (D) of −10.7 m2/s, whereas Re was lost to the air more slowly (log D = −12.2 m2/s). Volatile/melt partition coefficients were not determined in that study. Within the temperature regime of hydrothermal systems, Xiong and Wood (1999) and Xiong and Wood (2002) identified 4+ to be dominant oxidation state of Re in solution, and noted a steep positive correlation between Re solubility and chloride concentrations in the fluid phase. Rhenium solubilities of about 36 ppm were determined in 1 m KCl solution at T = 500 ºC, P = 55 MPa, fO2 = Re–ReO2 (FMQ + 3.2) and pH of the K-feldspar–muscovite–quartz buffer. Xiong and Wood (2002) proposed that the most stable chloride complexes of Re at 400−500 ºC are ReCl40 and ReCl3+. They also pointed out that the solubility of ReS2 is about two orders of magnitude lower than that of ReO2 at these conditions, limiting the mobility of Re in the presence of reduced S-species for these elements.

CONCLUDING REMARKS Despite considerable efforts, our understanding of the geochemistry of the HSE in magmatic systems is still rather limited. Whereas the effects of temperature and oxygen fugacity are now reasonably well documented as to their impact on metal–silicate melt partitioning, most results have been obtained in experiments done at relatively low confining pressure. At such conditions, oxygen is only sparingly soluble in the Fe melt, however, so its influence on partitioning of the HSE is poorly known. Results of very high pressure experiments (i.e., >35 GPa) have shown that oxygen can influence the activity of the moderately siderophile elements in the metallic melt (Siebert et al. 2013), so this is an area that needs further exploration for the HSE. The zeroth order conclusion that sulfide is important to concentrate the HSE has been known for several decades. However, the absolute magnitude of DSulfLiq/SilLiq for the HSE is only now being constrained accurately, with the variation in DSulfLiq/SilLiq with fO2 and fS2 reasonably well determined for Re, but results are scattered for Au, and largely unknown for the PGE. The recent discovery that alloys are important during mantle melting (and subsequent magma solidification) is still difficult to incorporate into petrogenetic models, as solubility data in both sulfide and silicate melt are incomplete. Moreover, the effect of dissolved sulfur (as well as other potential complexing agents, such as As) on HSE solubility in silicate melt is an area of considerable uncertainty, but initial results suggest this could be an important effect. With regard to low pressure degassing in magmatic systems, the majority of past experimental work has focussed only on the behavior of Au, Pt, and Pd, and in the hydrothermal regime. The capacity of low density orthomagmatic fluids to dissolve and transport the other HSE is less well known, but of importance to our understanding of the formation of magmatic ores, their possible remobilization, as well as the transfer of the HSE to the oceans and atmosphere.

ACKNOWLEDGMENTS The authors are grateful for the time spent by Raul Fonseca, Kate Kiseeva and Dave Walker to read and provide detailed and helpful reviews on an earlier version of this chapter. Editors James Day and Jason Harvey also provided detailed comments, and patiently nudged us along to chapter completion. Funding for Brenan’s research comes from Equipment, Discovery and Discovery Accelerator Grants from the Natural Sciences and Engineering Research Council of Canada. Bennett acknowledges post-doctoral support from the Carnegie Insitution of Washington, as well as previous graduate student funding from a Mineralogical Society of America Grant for Student Research in Mineralogy and Petrology, and a Geological Society of America Graduate Student Research Grant. Zajacz is supported by a Discovery Grant from the Natural Sciences and Engineering Research Council of Canada. Special thanks to Josée Normand and Yanan Liu who helped draft figures, compile the data tables and format the final document.

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Meteoritics and Planetary Science 46:199−217 Williams TJ, Candela PA, Piccoli PM (1997) Hydrogen-alkali exchange between silicate melts and two-phase aqueous mixtures: an experimental investigation. Contrib Mineral Petrol 128:114–126 Williams-Jones AE, Heinrich CA (2005) 100th Anniversary special paper: Vapor transport of metals and the formation of magmatic-hydrothermal ore deposits. Econ Geol 100:1287−1312 Williams-Jones AE, Bowell RJ, Migdisov AA (2009) Gold in Solution Elements 5:281−287 Willis J, Goldstein JI (1982) The effects of C, P, and S on trace element partitioning during solidification in Fe–Ni alloys. Proc Lunar Planet Sci Conf 13th Part I J. Geophys Res 87, Suppl: A435–A445 Wißmann S, Wurmb VV, Litterst FJ, Dieckmann R, Becker KD (1998) The temperature-dependent cation distribution in magnetite. J Phys Chem Solids 59:321−330 Wood SA (2002) The aqueous geochemistry of the platinum-group elements with applications to ore deposits. In: The Geology, Geochemistry, Mineralogy and Mineral Beneficiation of PGE. Cabri, LJ (ed). CIM Spec Vol 54, Geol Soc of the Can Inst Min, Metall Petrol, Montreal Quebec Wood M (2003) Arsenic in igneous systems: an experimental investigation: BASc Thesis, University of Toronto, Toronto, Ontario Wood SA, Mountain BW (1991) Hydrothermal solubility of palladium in chloride solutions from 300 ºC to 700 ºC—Preliminary experimental results—A discussion. Econ Geol 86:1562−1563 Wood SA, Mountain BW, Pan P (1992) The aqueous geochemistry of platinum, palladium and gold—Recent experimental constraints and a reevaluation of theoretical predictions. Can Min 30:955−982 Wood SA, Pan P, Zhang Y, Mucci A (1994) The solubility of Pt and Pd sulfides and Au in bisulfide solutions. 1. Results at 25 ºC–90 ºC and 1 bar pressure. Mineral Dep 29:309−317 Wood BJ, Walter MJ, Wade J (2006) Accretion of the Earth and segregation of its core. Nature 441:825−833 Xiong YL, Wood SA (1999) Experimental determination of the solubility of ReO2 and the dominant oxidation state of rhenium in hydrothermal solutions. Chem Geol 158:245−256 Xiong Y, Wood SA (2000) Experimental quantification of hydrothermal solubility of platinum-group elements with special reference to porphyry copper environments. Mineral Petrol 68:1−28 Xiong YL, Wood SA (2002) Experimental determination of the hydrothermal solubility of ReS2 and the Re–ReO2 buffer assemblage and transport of rhenium under supercritical conditions. Geochem Trans 3:1−10 Yokoyama T, Walker D, Walker RJ (2009) Low osmium solubility in silicate at high pressures and temperatures. Earth Planet Sci Lett 279:165–173 Yu G, Jacobsen SB (2011) Fast accretion of the Earth with a late Moon-forming giant impact. PNAS 108:17604–17608 Yudovskaya MA, Tessalina S, Distler VV, Chaplygin IV, Chugaev AV, Dikov YP (2008) Behavior of highlysiderophile elements during magma degassing: A case study at the Kudryavy volcano. Chem Geol 248:318−341 Zaccarini F, Garuti G, Cawthorn RG (2002) Platinum-group minerals in chromitite xenoliths from Onverwacht and Tweefontein ultramafic pipes, eastern Bushveld Complex, South Africa. Can Min 40:481−497 Zajacz Z, Halter WE, Pettke T, Guillong M (2008) Determination of fluid/melt partition coefficients by LAICPMS analysis of co-existing fluid and silicate melt inclusions: Controls on element partitioning. Geochim Cosmochim Acta 72:2169−2197 Zajacz Z, Seo JH, Candela PA, Piccoli PM, Heinrich CA, Guillong M (2010) Alkali metals control the release of gold from volatile-rich magmas. Earth Planet Sci Lett 297:50−56 Zajacz Z, Seo JH, Candela PA, Piccoli PM, Tossell JA (2011) The solubility of copper in high-temperature magmatic vapors: a quest for the significance of various chloride and sulfide complexes. Geochim Cosmochim Acta 75:2811−2827

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Zajacz Z, Candela PA, Piccoli PM, Sanchez-Valle C (2012a) The partitioning of sulfur and chlorine between andesite melts and magmatic volatiles and the exchange coefficients of major cations. Geochim Cosmochim Acta 89:81−101 Zajacz Z, Candela PA, Piccoli PM, Sanchez-Valle C, Wälle M (2012b) Gold and copper in volatile saturated mafic to intermediate magmas: solubilities, partitioning and implications for ore deposit formation. Geochim Cosmochim Acta 91:140−159 Zajacz Z, Candela PA, Piccoli PM, Sanchez-Valle C, Waelle M (2013) Solubility and partitioning behavior of Au, Cu, Ag and reduced S in magmas. Geochim Cosmochim Acta 112:288−304 Zhang L, Fei Y (2008) Effect of Ni on Fe–FeS phase relations at high pressure and high temperature. Earth Planet Sci Lett 268:212–218 Zhang SX, Wu SY, Xu P, Li LL (2010) Theoretical investigations of the spin Hamiltonian parameters and the local structure for Rh2+ in rutile, Can J Phys 88:49−53 Zientek ML, Likhachev AP, Kunilov VE, Barnes S-J, Meier AL, Carlson RR, Briggs PH, Fries TL, Adrian BM (1994) Cumulus processes and the composition of magmatic ore deposits: examples from the Talnakh District, Russia, in: PC Lightfoot, AJ Naldrett Eds, Proc Sudbury-Noril’sk Symposium, Ont Min Northern Develop Mines, Spec Vol 5, p 373−392 Zoller WH, Parrington JR, Kotra JMP (1983) Iridium enrichment in airborne particles from Kilauea volcano: January 1983. Science 222:1118–1121

2

Reviews in Mineralogy & Geochemistry Vol. 81 pp. 89-106, 2016 Copyright © Mineralogical Society of America

Analytical Methods for the Highly Siderophile Elements Thomas Meisel General and Analytical Chemistry Montanuniversität 8700 Leoben Austria [email protected]

Mary F. Horan Department of Terrestrial Magnetism Carnegie Institution of Washington Washington, DC 20015 USA [email protected]

INTRODUCTION The highly siderophile elements (HSE) include the fifth-period transition metals ruthenium (Ru), rhodium (Rh), palladium (Pd), and the sixth-period transition metals rhenium (Re), osmium (Os), iridium (Ir), platinum (Pt) and gold (Au). In addition to being iron-loving, these elements are also resistant to oxidation, have high melting temperatures and are important as industrial catalysts. HSE abundances in geologic materials vary significantly, ranging from ~1 mg/g in ore materials down to a few pg/g in basalts (Table 1). These elements comprise two long-lived radiometric decay schemes: 187Re decays to 187Os, and 190Pt decays to 186Os. The HSE have been targeted to address a wide variety of geochemical and cosmochemical questions. Early work suggested HSE concentrations can constrain Hadean mantle evolution (Chou 1978) and showed the geochronologic potential of the Re/Os isotope system (Herr and Merz 1955; Herr et al. 1961; Markey et al. 1998). More recent applications combine

Table 1. Range of HSE abundances, in ng/g, in selected rock types and in model upper mantle. Re

Os

Ir

Ru

Pt

Pd

Refs.

38.1

449

456

651

858

563

1

Primitive upper mantle

0.35

3.9

3.5

7.0

7.6

7.1

2

Peridotites

0.051–0.30

0.8–3.4

1.5–4.9

2.1–9.1

1.6–8.3

0.6–6.5

2, 3, 4

0.1–2

0.03–0.7

0.05–0.9

0.08–3.7

2.3–22

1.4–33

5

200 °C) under high pressure conditions (> 50 bar) in sealed borosilicate or quartz glass containers. For isotope dilution analyses, a mixed isotopic spike is added to the sample powder along with the dissolution acid. Both high temperature and closed system dissolutions are necessary for full isotopic equilibration between sample powder and the isotope spike (Shirey and Walker 1995). One technique uses Carius tubes, a borosilicate or quartz glass tube with a narrow neck for easy sealing with a torch. Carius tubes were originally developed for elemental analysis (C, H, O, N, S, P) of organic substances (Carius 1860, 1865), described for digestion of platiniferous materials (Gordon 1943; Gordon et al. 1944), then applied to Re-Os isotopic analysis (Shirey and Walker 1995). This technique does not involve expensive instrumentation, but requires some skill with a torch to produce leak-proof seals. Most studies use borosilicate Carius tubes, but modified cleaning techniques for borosilicate and quartz Carius tubes can be used to better assure lower blanks for Pt (Day and Walker 2016). Up to 3 g of finely ground sample powder is reacted in a Carius tube in reverse aqua regia. Each Carius tube is placed in a steel protection vessel prior to heating in an oven at 220–345 °C for up to several days. Use of temperatures above about 270 °C usually requires an air-tight protection vessel and counterpressure supplied by dry ice (e.g., Becker et al. 2006). After heating, the Carius tubes are cooled to room temperature then chilled in a mixture of dry ice and alcohol, or in ice, to minimize overpressure during opening. Tubes are opened behind a protective shield. The other method of acid digestion uses a high pressure asher (HPA, Anton Paar, Graz) which employs reusable quartz glass containers to maintain high temperature and high pressure under controlled conditions. Up to 5 g of sample powder can be treated in reverse aqua regia (HNO3 and HCl volume ratio 5:2) in the HPA vessels at 300–320 °C maintained for several hours. However, the manufacturer recommends not to use the HPA for prolonged periods above 280 °C when HCl is present. Some level of user skill is required to assemble the quartz tubes so that a leak-proof seal is produced. The main drawbacks to this dissolution technique are the high initial expense of the HPA and maintenance expenses caused by corrosion of the autoclave and tubing in the HPA by HCl fumes. This system, however, can yield especially low blanks for HSE (Meisel et al. 2001b). Low-temperature acid attack utilizing concentrated HBr and HF in closed Teflon®PFA developed by the Paris group (Birck et al. 1997) shows severe underestimation for Os results for whole rock analysis as shown by Meisel et al. (2003b) but can serve as low blank alternative for individual base metal sulfide and aggregates of silicate grains. Systematic studies comparing results using different digestion methods and conditions have yielded contradictory results. For example, Meisel et al. (2003b) found that for a serpentinized peridotite reference material, UB-N, temperatures between 230 to 240 °C, attained in Carius tubes, were insufficient to completely digest its HSE-bearing spinels and HSE alloys; the highest yields of HSE were obtained at temperatures of ≥ 300 °C in the HPA. By contrast, Puchtel and Humayun (2005), Harvey et al. (2010) and Ishikawa et al. (2014) found little difference between HPA results for UB-N and those obtained by dissolution in Carius tubes at 230 °C. Other rock types yield different comparative results. For example, Becker et al. (2006) found that spinel grains from fresh peridotites remained undigested at 220–240 °C in Carius tubes but were completely dissolved or disaggregated at 345 °C. The latter protocol utilized counter pressure to achieve the highest temperature in Carius tubes. The contrast in these results with those for UB-N suggests that “the process of serpentinization may be helpful in pre-digesting peridotites” (Becker et al. 2006). At least two studies indicate that some basalts may benefit from treatment with HF after aqua regia dissolution. Dale et al. (2012) showed that de-silicification with HF, after dissolution using aqua regia in a HPA, more efficiently extracted Ir, Ru Pt, Pd and Re. Basalt reference material

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sample preparation

sample preparation

TDB-1 was used by Ishikawa et al. (2014) to study the effectiveness of different digestion techniques. Improved results in TDB-1 were obtained by additional treatment with hydrofluoric acid after Carius tube processing, a result confirmed by the Na2O2 sintering technique (Bokhari and Meisel 2014) (Fig. 2). In TDB-1, prolonged heating in aqua regia at elevated temperatures also completely releases Re from the silicate minerals, without additional treatment in HF, but does not completely release Ru (Figs. 2 and 3). Like TDB-1, the use of HF during dissolution

Figure 2. Literature survey of HSE concentration data in basalt TDB-1 separated by digestion technique. AD refers to digestions that use only HCl–HNO3 and were performed in Carius tubes and by HPA with temperatures between 240 and 320 °C, and in a microwave oven using PFA vials at 150 °C. AD–HF refers to the same variety of digestion methods as AD, that also included an additional HF dissolution step. FAID: NiS fire assay using isotope dilution. SI: Na2O2 sintering. The additional HF treatment always yields the highest values for Re and Ru, independent of digestion instrument and temperatures. Sintering also fully releases those elements from the silicate matrix while the fire assay technique seems to be problematic. Data from (Peucker-Ehrenbrink et al. 2003; Meisel and Moser 2004a; Mungall et al. 2006; Shinotsuka and Suzuki 2007; Qi and Gao 2008; Dale et al. 2009; Paquay et al. 2009; Simonson et al. 2009; Park et al. 2012; Goderis et al. 2013; Xie et al. 2013; Bokhari and Meisel 2014; Honda et al. 2014; Ishikawa et al. 2014; Marchesi et al. 2014; Yang et al. 2014; Chu et al. 2015; Li et al. 2015).

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Figure 3. Re and Ru concentration data in TDB-1and for BIR-1 obtained using HCl–HNO3 digestion in Carius tubes and by HPA at temperatures between 240 and 320 °C. The duration of this HCl–HNO3 step is given by the size of the symbol. AD refers to digestions that used only HCl–HNO3. AD–HF refers to those methods that included HF treatment either prior or after HCl–HNO3 digestion. For basalt TDB-1, these results show that release of Re can be achieved by treatment with HF or by digestion in HCl–HNO3 for > 60 h. Ru from TDB-1 is not fully liberated during HCl–HNO3 digestions for >60 h, but instead appears to require treatment with HF. Data for BIR-1 yield similar results to those for TDB-1. Data for BIR-1 from: (Meisel and Moser 2004a; Ishikawa et al. 2014; Chu et al. 2015; Li et al. 2015) and for TDB-1 from (Mungall et al. 2006; Dale et al. 2009; Simonson et al. 2009; Park et al. 2012; Goderis et al. 2013; Ishikawa et al. 2014; Chu et al. 2015; Li et al. 2015).

appears essential to fully release all Ru in some basalts (Fig. 3). These results suggest that some proportion of HSE in some basalts is hosted in phases not accessed by aqua regia. On the other hand Day et al. (2015) showed that the use of HF can alter measured Re/Os and Pt/Os ratios in some high-MgO basalts and increase blanks. Thus, the effectiveness of the digestion technique of choice needs to be validated by applying different techniques to the samples of interest and by the use matrix matched reference materials (see later).

Chemical separation of HSE For samples from which more than about 100 ng of HSE are obtained, interferences and matrix effects become less important during ICP-MS analysis. In this case direct ICPMS measurement may be possible without further chemical isolation of these elements. For lower abundances and for more interference-free analysis, preconcentration of the HSE is necessary. Isolation of Os is usually by solvent extraction from aqua regia into CCl4 or CHCl3 and back extraction into HBr (Cohen and Waters 1996). In this process, Os, oxdized as OsO4, is extracted from aqua regia through liquid-liquid extraction into the solvent by shaking. This extraction is repeated for a total of 3 times (as each step extracts about 60–70 %), with the Oscontaining solvent removed to a clean beaker after each extraction. Osmium is then recovered from the solvent by addition of concentrated HBr which is shaken and emulsified. During this step, Os is reduced and partitioned into HBr. After removing and discarding the solvent, the HBr solution containing the Os is dried. Alternatively, OsO4 can be distilled directly from the solution after acid digestion (Morgan and Walker 1989; Nägler and Frei 1997). Further cleanup of Os is accomplished by microdistillation (Roy-Barman 1993; Birck et al. 1997). The Os is transferred to the cap of a 5 mL Teflon®PFA conical beaker, covered with in a solution of CrO3 in 6 mol/L H2SO4. A small drop of HBr (20 µL) is placed in conical tip of the beaker.

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The beaker is capped in an upside down position and heated to 80 °C for 2–3 h. During the microdistillation, the Os is oxided by the Cr solution, volatilized and trapped in the HBr. The solution of Os in HBr can then be dried to 1–2 µL, and is ready for mass spectrometry. For purification of the remaining HSE, each laboratory appears to have developed its own unique procedure. Two general approaches are used: separations using anion exchange resin (e.g., Rehkämper and Halliday 1997; Pearson and Woodland 2000; Meisel et al. 2001b; Horan et al. 2003; Chu et al. 2015) or using cation exchange resin (e.g., Meisel et al. 2003a; Fischer-Gödde et al. 2010). The HSE typically form anion complexes in HCl (most HSE) or HNO3 (Re). Use of anion exchange resin, typically AG1-X8, and high molarity acids allows separation of HSE into groups, such as Re + Ru and Pt + Ir, that are especially well-suited for analysis by magnetic sector ICP-MS. Potential isobaric interferences on Ru from chromium oxides can be minimized by treatment of the sample with H2O2 prior to anion exchange chromatography (e.g., Dale et al. 2012). The main drawback to separation of the HSE + Re by anion exchange is that some HSE, particularly Pd, may have relatively low recoveries. In the other approach, the HSE are little adsorbed by cation exchange resin and therefore elute quickly as a group in weak acid, while the major and most trace elements are retained on the resin. This technique can be applied off-line in batches or on-line by coupling the column directly to the ICP-MS. In both cases near-quantitative recovery of the HSE is obtained, and therefore can be modified for analysis of unspiked monoisotopic Rh and Au (Meisel et al. 2003a; Qi et al. 2004; Fischer-Gödde et al. 2010). Interferences from Cd and Zr must be carefully monitored or separately removed. For cosmochemical applications in which the isotope compositions of the HSE are to be measured to ε-level (1 part in 10,000) or better precision, additional purifications steps are necessary (Yokoyama and Walker 2016, this volume, and references therein). Quantification by mass spectrometry. Osmium is usually measured as OsO3− by negative thermal ionization mass spectrometry (N-TIMS) on Pt-filament material using Ba(OH)2 as an electron emitter (Creaser et al. 1991; Völkening et al. 1991). Alternatively, OsO4 may be sparged from a solution directly into an ICP-MS (Hassler et al. 2000). Quantification of the other HSE is by ICP-MS, either single collector analysis using a magnetic sector field (ICP-SFMS) or quadrupole (ICP-QMS) mass spectrometer (e.g., PeuckerEhrenbrink and Jahn 2001; Meisel et al. 2003a; Wang and Becker 2014) or by magnetic sector multicollector ICP-MS (e.g., Puchtel et al. 2007). The ICP-SFMS and ICP-QMS allow the possibility of measuring all of the HSE in a single aliquot. Magnetic sector mass spectrometers such as the single collector ICP-SFMS “ELEMENT” and MC-ICP-MS (e.g., Nu Plasma and Thermo Neptune) offer higher sensitivity analysis. For isotope dilution analysis, the high mass range elements, i.e., Ir, Os, Re, Pt, and Au are nearly interference-free. Lower mass HSE (Ru, Rh, and Pd) have spectral interference either from molecules produced by the argon plasma, doubly charged species, or molecular species of elements inadequately removed during chemical purification (e.g., Meisel et al. 2001b; Puchtel and Humayun 2005; Fischer-Gödde and Becker 2012). Higher precision measurement results of the full isotope compositions of the HSE for nucleosynthetic or cosmogenic applications require more stringent monitoring of interferences are discussed in the references cited in Yokoyama and Walker (2016, this volume). Procedural blanks. The need for lower analytical blanks is driven by the interest in the HSE distribution in geologic samples having HSE levels at sub-ng/g levels and also for the study of individual grains. The HSE belong to the least abundant group of elements in the Earth’s crust, but may be elevated through anthropogenic influence on the environment, for example by the use of automobile catalytic converters. Even in a dust-free environment, contamination occurs from mineral acids, alkaline flux, resins, metals, and glassware used during sample preparation. The relatively low concentrations of HSE in most meteorites and terrestrial rocks require the use of highest-purity reagents. Sub-pg/mL blank levels of HSE in mineral acids usually can be achieved by subboiling distillation.

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Total procedural blanks for solvent extraction of Os of less than a few pg are typical, with an 187Os/188Os ratio of 0.15–0.20 (e.g., Puchtel et al. 2007; Dale et al. 2012; Wang and Becker 2014). A major contributor this blank is Os in HNO3 that is not removed by subboiling distillation. Further purification of HNO3 can be achieved by careful treatment with H2O2 (R. Creaser, personal communication). Procedural blanks levels for Re are similar to those for Os, while those for Pt, Pd, and Ru may be higher (e.g., Puchtel et al. 2007; Dale et al. 2012; Wang and Becker 2014). Chromatographic resin can be a major contributor to these blanks; anion resin should be cleaned using high molarity HNO3 and/or HCl before use. For samples having low abundances of HSE, the contribution of the procedural blank to their measurement uncertainties cannot be neglected. Figure 4 shows how the procedural blank and its own uncertainty is propagated into the resulting uncertainty for sample Os concentrations at various levels. For details see Moser et al. (2003). The uncertainty in Os concentrations at the lowest levels can be improved not only by a decrease in the absolute amount of the procedural blank, but also by reduction in the variance of the average of the blank determinations. Some borosilicate Carius tubes have been found to have high-Pt blanks; more aggressive cleaning procedures (e.g., Puchtel et al. 2008) or substitution of quartz Carius tubes (e.g., Day et al. 2010) can help. The lowest and most reproducible blanks for the HSE can often be obtained by dissolution using the HPA. In the analysis of HSE for Apollo lunar samples, Day et al. (2007) pointed out that Pt was the element that showed the most discrepancy between HPA and Carius tube dissolutions, possibly resulting from underestimation of Pt-blank variability for the latter method. They concluded that Pt-blank measurements for the HPA were highly reproducible.

blank 0.001 +/− 0.004 ng

uncertainty %RSD

12

Figure 4. A comparison of the measurement uncertainty with sample Os concentrations for a given procedural blank. The measurement uncertainty is dramatically increased for samples having Os concentrations that are close to the procedural blank.

8

4 0.05

0.10 sample Os [ng/g]

0.15

REFERENCE MATERIALS FOR HSE ANALYSIS Reference materials (RM) play a key role in method development and for quality control in routine analyses. Further, they can serve as calibration standards and can be used to establish a traceability chain (BIPM et al. 2008). Reference materials should be homogenous, meaning that the contribution of sample heterogeneity must be very small compared to the overall measurement uncertainty. The homogeneity of a RM can only be guaranteed for sample sizes that are larger than some minimum sample size. Such a minimum should be given in a certificate of analysis for a reference material. HSE abundance and isotopic data for RM ideally should have small uncertainties in order to better resolve results produced by different measurement procedures or in different laboratories. This goal is currently only achievable with relatively easy- to-digest and relatively homogeneous materials such as fertile peridotites.

Comment

**Contaminated with Henderon Molybdenite

*GeoReM,

USGS

certified*

Major and trace elements content

Source

all bottled

basalt

USA, Oregon

Bottling

Availability

Petrography

Origin

BCR-1 and BCR-2**

*GeoReM

USGS

certified*

all bottled

basalt

USA, Hawaii

BHVO-1 and BHVO-2

out of stock

Comment

Mafic Rocks

IAGeo

SARMCRPG/CRNS

out of stock

Source

*GeoReM, successor BIR-2 not available

USGS

certified*

all bottled

< 5 kg

dolerite

Iceland, Reykjavik

BIR-1

contaminated with lead glass

Proficiency testing results available

400 g sachets

characterized

no

> 100 kg

> 100 kg

prepared in batches

harzburgite

Albania, Devolli

HARZ-01

Major and trace elements content

Bottling

< 500 g

Availability

lherzolite

France, Voges

Marocco, Beni Bousera

lherzolite

UB-N

GP13

CanmetMINING

characterized

all bottled

CanmetMINING

characterized

all bottled

gabbro

Canada, Wellgreen Complex

Canada, Tremblay Lake diabase

WGB-1

Japan. Geol. Survey

certified

TDB-1

IAGeo

ISO certified

all bottled

?

>100 kg all bottled

peridodite

Japan, Horomon

JP-1

harzburgite

Austria, Kraubath

MUH-1

*Depleted but WPR1a is now available

CanmetMINING

REE certified

all bottled

out of stock*

mineralized peridotite

Canada, Wellgreen Complex

WPR-1

IAGeo

ISO certified

all bottled

> 100 kg

dunite

Mongolia

GAS

Table 2. Most common HSE reference materials used for method validation and quality control (cont’d on next page).

Petrography

Origin

Ultramafic Rocks

IAGeo

ISO certified

all bottled

> 100 kg

komatiite

Canada, Abitibi

OKUM

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99

Table 2 (cont’d). Most common HSE reference materials used for method validation and quality control. Sediments Origin Petrography Availability Bottling Major and trace elements content Source Comment

SCo-1

SDO-1

JCh-1

shale

black shale

chert

Wyoming, USA < 1 kg

out of stock

all bottled

all bottled

all bottled

well characterized

well characterized

well characterized

USGS

USGS

Japan. Geol. Survey

out of stock, SCo-2 will be available in future

out of stock

Certified reference materials produced following ISO guides are widely used in commercial gold, silver, palladium, and platinum analysis. Examples include WMG1 (mineralized gabbro, NRCan), SARM-7 (Merensky Reef platinum ore, SA Bureau of Standards). In contrast, few well-characterized and matrix-matched reference materials exist for samples having non-commercial levels of HSE. We briefly discuss some of these low-level HSE RM used in the geochemistry community. Mafic reference materials include TDB-1 (basalt) and WGB-1 (gabbro) which have been certified for Pd, Pt, and Au concentrations (NRCan 1994a,b). These materials were characterized during method validation studies and are now among the best studied nonmineralized, low abundance RM with a silicate matrix (Enzweiler et al. 1995; Plessen and Erzinger 1998; Meisel et al. 2001b; Bédard and Barnes 2002; Qi et al. 2003; Meisel and Moser 2004a; Qi et al. 2004; Boulyga and Heumann 2005; König et al. 2012; Li et al. 2013). Icelandic basalt BIR-1, a coarse-grained olivine tholeiite provided by the USGS, has proven to be homogeneous for HSE even for test portion sizes of less than 2 g (Ishikawa et al. 2014) The USGS RM basalts BHVO-1, BHVO-2 and BCR-2 have also been characterized for HSE concentrations (Meisel and Moser 2004a; Li et al. 2013; Chu et al. 2015) but fewer systematic studies of digestion techniques are available. An initiative to develop a komatiite sample, KAL-1 from a single Alexo lava flow, was unsuccessful, as only a small amount of material was produced and differences in the HSE distribution between batches became apparent (J. Carignan, CRPG, personal communication). Instead, an existing komatiite RM named OKUM collected from Abitibi, Canada, by the Ontario Geological Survey is being characterized (Wang and Becker 2014). More than 100 kg of material were specially prepared for the certification of major and trace elements by the International Association of Geoanalysts (IAG) following ISO Guidelines, and is being distributed to laboratories for development as a RM for HSE. Lherzolites are especially useful as RM for HSE. The bulk of the HSE budget of fertile peridotites, i.e., sulfur-rich lherzolites, typically is hosted in base metal sulfides. These sulfides are easily dissolved during acid digestion of the peridotites. Lherzolite RM such as UB-N and GP13 have been well characterized for HSE contents, as well as for their 187Os/188Os compositions (Pearson and Woodland 2000; Meisel et al. 2003b, 2004; Pearson et al. 2004; Day et al. 2012). GP13, a fertile lherzolite from the Beni Bousera massif, Morocco, was initially developed as in-house RM but is now no longer available. UB-N, a serpentinized, fertile lherzolite, is still available from the SARM at the CRPG but only in coarse grained batches that are further pulverized to powder on demand. Differences in Cr content and in HSE contents are present in different batches, probably as a result of variations in the spinel abundances (J. Carignan, CRPG, personal communication; H. Becker, FU Berlin, personal communication).

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Satisfactory RM for harzburgites are lacking. A harzburgite RM, MUH-1, from the Preg Quarry, Kraubath, Austria, was prepared for the certification of major and trace elements, and is being distributed for interlaboratory comparisons of HSE and Os isotopic compositions. This sample, however, appears to be subject both to incomplete recovery of HSE during dissolution and the nugget effect at 5 g sample size, as discussed in an earlier section. Table 2 gives an overview of RM that have been used for method validation and quality control. Several of these RM were not developed for the intended use as HSE RM. As such more RM that have been analyzed for HSE content but were not included in this table as too few data have been published.

APPENDIX Terminology and definitions. A reported result without quantitative indication of its quality is of limited use, and makes comparison among samples difficult and evaluation of results obtained by different methods or in different labs complex. A common language for scientific measurements has been defined in the VIM3 (BIPM et al. 2008). Because some of these terms are little used in the geologic community, here we provide definitions and recommendations to avoid misunderstandings of metrological concepts and make data easier to compare. Italicized text are taken from VIM3 (BIPM et al. 2008) unless otherwise noted. Measurand. “Quantity to be measured”. For example the measurand can be defined as: (a) The mass fraction of Pt in a 100 mg test portion; (b) The mass fraction of Pt in the laboratory sample; (c) The mass fraction of Pt in particular geological formation. “Analyte”, by contrast, refers to the name of a substance or compound. Measurement precision. “The closeness of agreement between indications or measured quantity values obtained by replicate measurements on the same or similar objects under specified conditions, usually expressed numerically by measures of imprecision, such as standard deviation, variance or coefficient of variation.” While the term precision is correctly used within the earth science community, the conditions under which the measurement precision was obtained also should be reported. These conditions can be, for example, repeatability conditions of measurement, intermediate precision conditions of measurement or reproducibility conditions of measurement (see ISO 5725-3:1994 and vide infra ). The most common kinds of precision are listed below and shown in Appendix Figure 1. Measurement repeatability. Measurement precision obtained using the same operators, same measuring system, same operating conditions and same location, and replicate measurements on the same or similar objects over a short period of time. This term is preferred over “internal reproducibility” or “internal precision”. Intermediate measurement precision. This term includes the same measurement procedure, same location and replicate measurements on the same or similar objects over an extended period of time. It may include other changes, including new calibrations, calibrators, operators and measuring systems, and should specify which conditions are changed and unchanged, to the extent practical. This term is preferred over “external reproducibility” or “external precision”. Appendix Figure 1a and b show intermediate measurement precisions obtained in two laboratories for the Ir content of the komatiite powdered reference material OKUM. Measurement reproducibility. Measurement precision under reproducibility conditions that includes different locations, operators, measuring systems, procedures and replicate measurements on the same or similar objects. This precision can usually only be approximated, and will be the largest all types of precision. It is thus the opposite of measurement repeatability. Information on the measurement conditions that changed and remained unchanged, to the extent practical, should be provided. Appendix Figure 1c shows the reproducibility precision for Ir in OKUM, as measured in 7 laboratories.

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Ir in OKUM lab #1 intermediate precision

lab #2

all labs n=7

intermediate precision

reproducibility precision

Appendix Figure 1. A comparison of two common types of measurement precisions is shown. (a, b) Intermediate precision is given by data produced on independent digestions within a single lab. (c) Reproducibility precision refers to measurements performed in different labs, with different sample preparation methods, different calibration strategies etc. Data shown are taken from an unpublished interlaboratory comparison of Ir measured in komatiite candidate reference material OKUM. The vertical solid line gives the mean of the data set; the dotted lines show the one standard deviation boundaries. For clarity, Ir concentrations > 2.5 ng/g are not shown. Unpublished data provided by Marcus Burnham (OGL), Iouri Borin (VSGEI) and Thomas Meisel (Montanuniversität Leoben).

ng/g

Measurement accuracy. This term refers to the “closeness of agreement between a measured quantity value and a true quantity value of a measurand.” Measurement accuracy is not a numerical quantity, but a measurement is said to be more accurate when it offers a smaller measurement error. Quantification and detection limit. The use of the term “detection limit” is not preferred, as its usage is inconsistent and provides little information on the actual uncertainty of a given measurement. More useful information is given by the “quantification limit” which has no rigorous definition, but is generally defined as the lowest limit above which the measurement uncertainty of the result is fit for purpose. Measurement uncertainty. This term refers to “the dispersion of the values that could reasonably be attributed to the measurand” (BIPM et al. 1993). Measurement uncertainty (Appendix Fig. 2a and b), quantified by standard deviations, comprises the statistical distribution not only of a series of measurements, but also includes other components that arise from probability distributions based on experience or other information (BIPM et al. 1993). A best estimate for the value of the measurand, therefore, includes a measurement uncertainty that incorporates all components of uncertainty that contribute to the dispersion, including systematic effects, for example those associated with corrections and reference standards, (BIPM et al. 1993). An older but less general definition of measurement uncertainty is “an estimate characterizing the range of values within which the true value of a measurand lies” (BIPM et al. 1984, definition 3.09). This definition is not inconsistent with the latest VIM3 definition but it makes clear that the measurement uncertainty is actually a range of values large enough to encompass the true, but unknown, value (Appendix Fig. 2c). Thus, information on precision and accuracy is not sufficient as it does not provide a range. More information on the correct use of “measurand”, “error” and “uncertainty” is given in Annex D of the Guide in estimation uncertainty (BIPM et al. 2008).

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reference quantity value

Systematic Error

Measurement Error

Random Error

True Value within this range

Precision

u a)

b)

U

c)

Appendix Figure 2. Three different expressions of the quality of a measurement result. a) Two parameters describe the quality of a measurement: precision is a statistical quantity that describes repeatability or intermediate precision; accuracy refers to its deviation from the true value. As the true value is unknown and only approximated, the concept of accuracy is not rigorous. b) The deviation of a measurement result from a reference or certified value gives the “measurement error” but, here, the error bar does not encompass the reference value. c) According to current metrological concepts, the use of measurement uncertainty (u or 2u = U), in which all possible systematic and non-systematic errors are accounted for, is preferred. The range of the measurement uncertainty should include the true value.

Measurement error. This term refers to a measured quantity value minus a reference quantity value (Appendix Fig. 2b), and is only applicable in the rare case when the reference quantity value has a negligible measurement uncertainty.

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Yang AY, Zhou M-F, Zhao T-P, Deng X-G, Qi L, Xu J-F (2014) Chalcophile elemental compositions of MORBs from the ultraslow-spreading Southwest Indian Ridge and controls of lithospheric structure on S-saturated differentiation. Chem Geol 382:1–13, doi:http://dx.doi.org/10.1016/j.chemgeo.2014.05.019 Yokoyama T, Walker RJ (2016) Nucleosynthetic isotope variations of siderophile and chalcophile elements in the Solar System. Rev Mineral Geochem 81:107–160

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Reviews in Mineralogy & Geochemistry Vol.81 pp. 107-160, 2016 Copyright © Mineralogical Society of America

Nucleosynthetic Isotope Variations of Siderophile and Chalcophile Elements in the Solar System Tetsuya Yokoyama Department of Earth and Planetary Sciences Tokyo Institute of Technology Ookayama, Tokyo 152-885 Japan [email protected]

Richard J. Walker Department of Geology University of Maryland College Park, MD 20742 USA [email protected]

INTRODUCTION Numerous investigations have been devoted to understanding how the materials that contributed to the Solar System formed, were incorporated into the precursor molecular cloud and the protoplanetary disk, and ultimately evolved into the building blocks of planetesimals and planets. Chemical and isotopic analyses of extraterrestrial materials have played a central role in decoding the signatures of individual processes that led to their formation. Among the elements studied, the siderophile and chalcophile elements are crucial for considering a range of formational and evolutionary processes. Consequently, over the past 60 years, considerable effort has been focused on the development of abundance and isotopic analyses of these elements in terrestrial and extraterrestrial materials (e.g., Shirey and Walker 1995; Birck et al. 1997; Reisberg and Meisel 2002; Meisel and Horan 2016, this volume). In this review, we consider nucleosynthetic isotopic variability of siderophile and chalcophile elements in meteorites. Chapter 4 provides a review for siderophile and chalcophile elements in planetary materials in general (Day et al. 2016, this volume). In many cases, such variability is denoted as an “isotopic anomaly”; however, the term can be ambiguous because several preand post- Solar System formation processes can lead to variability of isotopic compositions as recorded in meteorites. Here we strictly define the term “isotopic anomaly” as referring to an isotopic deviation from the terrestrial composition resulting from the incorporation of varying proportions of elements with diverse nucleosynthetic origins into a meteorite component or parent body. The term will not be used here to refer to isotopic variations that result from massdependent isotopic fractionation, radioactive decay in the Solar System, or spallation effects. Based on astronomical observations and physical modelling, the formation of the Solar System has generally been thought to have initiated by the collapse of a dense molecular cloud core composed of gas and dust grains, which originated from diverse stellar environments (e.g., Boss 2003). This scenario regarding the birth of the Solar System has largely been corroborated by the discovery of presolar grains in primitive meteorites; presolar grains are 1529-6466/16/0081-0003$10.00

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remnant circumstellar dust grains which survived various destructive dynamic processes in the early Solar System, such as heating and collision, and were eventually incorporated into the parent bodies of primitive meteorites. Presolar grains are identified by their having isotopic compositions that can be drastically different from terrestrial, and from one another, depending on their mineral types and mode of origin (Zinner 2014). It was initially believed that presolar grains had been thoroughly mixed within the molecular cloud and/or in a turbulent protoplanetary disk, thereby resulting in homogeneous stable isotope compositions for all elements within the solar nebula, excluding some light elements such as hydrogen, carbon, nitrogen, and oxygen (e.g., Kerridge 1985; Clayton and Mayeda 1999). Consistent with this, most early studies of the isotopic compositions of numerous elements failed to find resolved isotopic anomalies on the scale of bulk aliquots of chondrites and differentiated meteorites (e.g., Murthy 1963; McCulloch and Wasserburg 1978a). The advent of multi-collector inductively coupled plasma mass spectrometers (MC-ICPMS), as well as improvements to thermal ionization mass spectrometers (TIMS) in the 1990s has enabled very high precision measurement of isotope ratios for a variety of elements. These mass spectrometers can efficiently generate ion beams of the elements of interest, which are accelerated and focused by high voltage potentials, then separated by an electromagnet into individual beam courses based on the mass/charge ratio of ions (m/z). Ionization is achieved either by the heat of a metal filament (TIMS), or plasma (ICP-MS). Both types of instruments require the chemical separation and purification of the element of interest, prior to isotopic measurement. This is necessary in order to avoid spectroscopic interferences from isobars and molecular ions. In ICPMS analysis, the purified element is dissolved in diluted acid and aspirated into the plasma ion source. The advantage of MC-ICP-MS is that ionization efficiency is high compared to thermal ionization for most of the elements in the periodic table. This enables isotopic analyses of elements with high ionization potentials (e.g., Zr, Hf), that are difficult or impossible to measure with TIMS. Also, although the plasma ionization source tends to strongly fractionate the isotopes of an element, the fractionation effects are generally stable and can be corrected for by interspersal of standards of known isotopic composition with samples, or by monitoring the isotopic ratio of an element of similar mass and known isotopic composition that is added to the sample. This means that MC-ICP-MS can be used to precisely measure the isotopic compositions of elements that consist of only two isotopes. By contrast, in TIMS analysis, the purified element is loaded on a metallic ribbon (e.g., Re, W, Ta, Pt, Ir), which is heated to cause ionization of the atoms. The advantage of TIMS is that the energy distribution of the ion beam is so small that the fluctuation of the ion beam intensity and the background noise are minimized compared to MC-ICP-MS. Correction for fractionation requires that at least three isotopes be measured. One ratio is used for fractionation correction. Modern MC-ICP-MS and TIMS instruments are now capable of making sub-5 ppm measurements (2 SD) for some isotopic ratios of some elements. This capability has enabled the detection of small but significant isotopic variability among bulk meteorites for some elements (typically at a level of ± 200 ppm or less). One of the first elements for which isotopic anomalies were identified in bulk planetary materials is Cr. Chromium is characterized by excesses and deficits of the neutron-rich isotope 54 Cr, compared to terrestrial standards: ε54Cr values (parts per 104 deviation of 54Cr/52Cr from the terrestrial ratio) in meteorites ranges from −0.7 ε for differentiated meteorites (eucrite, diogenite, mesosiderite, pallasite, angrite, and SNC) to +1.6 ε for CI chondrites (Shukolyukov and Lugmair 2006; Trinquier et al. 2007; Qin et al. 2010). With respect to siderophile and chalcophile elements, planetary scale isotopic variability has been most thoroughly documented for Mo and Ru (Dauphas et al. 2002a; Chen et al. 2010; Burkhardt et al. 2011). By contrast, no isotopic anomalies have yet been documented in bulk meteorites for Os, Te, and Cd, despite extensive searches (Fehr et al. 2005; Yokoyama et al. 2007, 2010; Wombacher et al. 2008). This contrast in the magnitude of isotopic anomalies among different elements may imply that Solar System

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precursors were not thoroughly homogenized by the dynamic processes operating within the solar nebula, and suggests instead that individual elements acted differently during the nebular processes. The cause of the inconsistency regarding isotope variability has not yet been resolved. In this chapter we review how nuclear reactions in a variety of stellar environments in the Galaxy produced the elements in the periodic table. We then discuss the nature of presolar grains, the carriers of isotopically anomalous components in the solar nebula at the time of planetesimal formation. Readers seeking more detailed overviews of stellar nucleosynthesis and presolar grains are referred to several recent reviews (Meyer and Zinner 2006; Nguyen and Messenger 2011; Heger et al. 2014; Zinner 2014). In subsequent sections, we review isotopic variability evidenced among siderophile and chalcophile elements, not only in bulk meteorites, but also in chondrite components such as acid leachates and residues, and calcium and aluminum-rich inclusions (CAIs). Finally, we discuss the origin of planetary scale isotopic variability in the solar nebula, specifically focusing on processes that may have led to the decoupling of isotopic anomalies among different elements. Such information is key for understanding the behavior of materials derived from diverse stellar sources in the solar nebula, and decoding how individual planets or planetesimals obtained diverse nucleosynthetic components in varying proportions.

ORIGIN OF ELEMENTS: STELLAR NUCLEOSYNTHESIS The Big Bang and subsequent nuclear reactions in various stellar environments have produced all of the elements in the universe. Hydrogen (1H, 2H) is the dominant product of Big Bang nucleosynthesis, while helium (3He, 4He) and trace amounts of 7Li and 7Be were also produced (Fields and Olive 2006). The theory of subsequent stellar nucleosynthesis was established by seminal studies, such as Burbidge et al. (1957) (referred to as the B2FH paper) and Cameron (1957). These studies recognized that elements not produced by the Big Bang were synthesized in tandem with the evolution of stars from birth to death. The pathway of nucleosynthesis is determined by the stellar mass, which can be divided into two general groups; low to intermediate mass stars (< 8 M; M refers to the mass of the Sun), and massive stars (> 8 M). In the following, we first summarize how stars at the hydrostatic equilibrium stage synthesized elements from He to Fe. We then focus on the nucleosynthesis of elements heavier than Fe, by which most of the siderophile and chalcophile elements are produced.

Production of elements from He to Fe via hydrogen to silicon burning Hydrogen Burning. The first step of stellar nucleosynthesis begins with H burning, which occurs in main sequence stars. Hydrogen burning is a process in which four 1H nuclei are converted to 4He via three proton-proton reaction chains (pp chains). In stars like the Sun, the burning is via the pp chains, which all commence with the reaction 1

H +1H → 2 H + e + + ν e

In this reaction, one of the two protons converts to a neutron, a very unlikely event in the scattering of two protons, which makes the reaction slow and allows the Sun to have a long lifetime (~1010 years). The next reaction is 2

H +1H → 3He + γ

The pp-1 chain (Fig. 1a), which dominates H burning in the Sun, is completed by 3

He +3He → 4 He + 21H

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(a)

νe

1H 1H

e+

1H

3He

1H

2H

e+

γ

(b)

pp-1 chain

γ

12C

1H

1H

3He

4He

1H

1H 4He

1H

νe

1H

γ

2H

1H

13N

15N

νe

e+

νe

16O 1H

γ

γ

13C

15O

γ 17F

1H

νe e+

e+

1H

14N

γ 4He

17O 1H

CNO cycle

Figure 1. Schematic view of hydrogen burning. (a) pp-1 chain and (b) CNO cycle.

Fig. 1 temperatures, the pp-2 and pp-3 chains can dominate in which 4He is synthesized At higher via 7Be and 7Li (pp-2) and 7Be, 8B, and 8Be (pp-3). The pp-1 chain is currently dominant (85%) in the interior of the Sun, and the pp-2 (15%) and pp-3 (0.02%) chains are much less active because they require temperatures higher than that of the center of the Sun. Besides the pp chain, the CNO cycle is another type of H burning which occurs in stars heavier than the Sun with core temperatures higher than 2 × 107 K. This is a dual cycle consisting of the CNand NO-cycles, in which four 1H nuclei are consumed to produce 4He, with reactions catalyzed by C, N, and O atoms co-existing in the star (Fig. 1b). The catalytic C, N, and O must be inherited from pre-existing stars.

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Helium Burning. As H burning proceeds, 1H atoms in a stellar center are exhausted and converted into 4He, forming a He core surrounded by a H-rich shell. No nuclear fusion initially occurs in the He core while the core begins to gravitationally contract, resulting in increasing pressures and temperatures. This contraction leads to H burning at the bottom of the H-rich shell, increasing the star’s luminosity. At this point the H-rich shell starts to expand significantly, and surface temperatures decrease. During this phase, termed the red giant phase, the star begins to ascend to the upper right on the Hertzsprung–Russell (H–R) diagram (red giant branch; Fig. 2). The He core increases its mass by the ash fall of He from the H-rich shell, and the core continues to contract. As a result, the core temperature rises sufficiently (108 K) to result in the fusion of accumulated 4He, i.e., He burning. In He burning, three α (4He) particles react and produce 12C (triple-α process) as follows: 4

He + 4 He + 4 He →12 C + γ

In some cases, especially in massive stars, 12C reacts with 4He and produces 12 C(α, γ)16O. As a result, a C/O core is formed in the stellar center.

16

O via

The subsequent scenario of nucleosynthesis is separated into two pathways depending on stellar mass. For low- to intermediate-mass stars (0.5–8 M), the C/O core is lighter than the critical mass necessary to initiate the nuclear fusion of 12C (1.07 M). The C/O core, thus, becomes increasingly dense and is supported by the pressure of degenerate electrons. Ultimately, He burning starts at the bottom of the He shell outside the C/O core, and the star ascends the asymptotic giant branch (AGB) pathway on the H–R diagram (Busso et al. 1999). The AGB phase is essential for producing many elements heavier than iron via the slow neutron capture process (s-process; see below). In the initial AGB phase, H burning occurs at the bottom of the H shell outside the He shell. As the He ashes accumulate on the He shell, the bottom of the He shell ignites explosively in a thermal pulse. The thermal pulse drives convection in the He shell and transfers materials in the He shell to the stellar surface, including the s-process nuclides.

Post -AGB TP-AGB First Thermal Pulse

Mbol

Early-AGB

Core He Flash

Core He exhaustion

RGB First dredge-up

He-burning

Core H exhaustion Zero age Main Sequence

Log (Teff) Figure 2. Schematic evolution of a 1 M star with solar metallicity in the H-R diagram. Figure modified from Busso et al. (1999).

Fig. 2

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As He and H shell burning continues in AGB stars, outer layers are eventually blown off by a strong stellar wind emitted by the star itself, leaving the central C/O core behind. The remaining electron-degenerate C/O core is called a “white dwarf”. When the white dwarf forms in a binary system, it can acquire H and He gas from the companion star and initiate explosive H burning on the surface, which is observed as a nova. The typical nucleosynthesis occurring in novae is explosive H burning, which produces 15N, 22Na, and 26Al via proton captures on CNO nuclei (e.g., Wiescher et al. 1986). If the white dwarf gains enough mass from the companion star to the level of the Chandrasekhar limit (~1.4 M), then the stellar temperature and pressure increase sufficiently to initiate the fusion of C and O. This leads to a runaway reaction, releasing enormous energy and leading to stellar disruption (Type Ia supernova). Type Ia supernovae are extremely energetic such that C and O are fused into isotopes of Fe and Ni. Consequently, Type Ia supernovae are an important contributor of 56Fe and other Fe-peak isotopes in galaxies. Carbon Burning. In the case of massive stars (>8 M), the mass of the C/O core formed after central He burning becomes heavier than the critical mass to fuse 12C (> 1.07 M). The C/O core continues to contract without electron degeneracy, and C burning initiates when the temperature reaches ~ 0.8 × 109 K: C +12C →24 Mg* 24 Mg* →20 Ne + 4He

12

Mg* →23 Na +1H 24 Mg* →24 Mg + γ 24

where 24Mg* is an excited state of 24Mg. The dominant reaction of 24Mg* decay produces 20Ne, so that C burning is the principal contributor of 20Ne in galaxies. The C shell burning starts when C in the core is exhausted. The C burning is the end stage of nuclear fusion for stars with masses ranging from 8 to 10 M. In such cases, the core is lighter than the critical mass to fuse 20 Ne (1.37 M), resulting in the formation of AGB stars with an electron degenerate O/Ne/ Mg core, comparable to the case of low- to intermediate-mass stars. When AGB stars lose the envelope by stellar wind, the remaining core becomes a O/Ne/Mg white dwarf. The O/Ne/Mg white dwarfs can also produce novae in binary systems. Neon Burning. For stars with masses of > 10 M, the O/Ne/Mg core continues to contract. When the core temperature reaches 1.5 × 109 K, 20Ne, the main product of C burning, starts to react as follows (Ne burning): 20 20

Ne + γ →16 O + 4 He Ne + 4 He →24 Mg + γ

After consuming all of the 20Ne, the 16O- and 24Mg-rich core contracts and is characterized by increasing temperature until the next nuclear fusion process initiates (O burning). Oxygen Burning. The O burning occurs when the temperature of a stellar core reaches 1.5 × 109 K following Ne burning. The nuclear reactions of O burning include: O +16O → 32 S* 32 * S → 28Si + 4He 32 * S → 31 P + 1H 32 * S → 32S + γ

16

where 32S* is an excited state of 32S. The primary product of the O burning is 4 He (α particle), as well as 32S, forming a Si/S core after consumption of 16O.

28

Si and

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Silicon Burning. As the Si/S core continues to contract after O burning, photodisintegration of 28Si occurs, which generates a supply of light particles (neutrons, protons, and 4He). Remaining 28Si captures these light particles to make heavier species up to nickel. These nuclei all come into a state of equilibrium under exchange of the light particles (a quasi-statistical equilibrium). Due to their relatively high binding energies, the so-called alpha nuclei (32S, 36Ar, 40Ca, 48Cr, 52Fe, and 56Ni) dominate the abundances in equilibrium. Eventually, radioactive 44Ti, 48Cr, and 52Fe, as well as 56Ni (T1/2 = 6.1 d), decay to 44Ca, 48Ti, 52 Cr, and 56Fe, respectively, via electron capture. The binding energy per nucleon (Eb) as a function of mass number, Eb increases from low to intermediate mass nuclides with a peak top around mass number = 56 (Fig. 3; Ghahramany et al. 2012). The binding energy then decreases towards heavy nuclides, such as 238U. This implies that the nucleosynthesis of heavier nuclides, via α-capture, proceeds spontaneously, releasing their binding energies until the formation of 56Ni. By contrast, the next reactions in which 56Ni captures light particles (neutrons, protons, and 4He) are non-spontaneous as they do not release binding energy but consume energy. The silicon burning is sufficiently slow that at least some of the 56Ni has time to decay into 56Fe. Consequently, 56Fe is the final nucleosynthetic product of the hydrostatic stellar evolution of massive stars. During Si burning in the stellar core, O-burning is proceeding in the shell enclosing the core, followed by Ne-, C-, He-, and H-burning proceeding from inner to outer shells. As a result, the star is made up of an “onion-skin” structure with several shells containing ashes of different nuclear fusion reactions in individual shells.

Binding energy per nucleon (MeV)

At the end of hydrostatic stage of massive stars, all fuel is consumed in the core and the core begins to contract. At this point, Fe in the core disintegrates into protons, neutrons, and α-particles by absorbing gamma rays, leading to the collapse of the stellar core. Because of the high density in the collapsed core, protons capture electrons to convert to neutrons. Finally, the stellar center is crushed into either a neutron star, which is supported by the degeneracy pressure of neutrons, or a black hole when the star is sufficiently massive (> 30 M). The outer layers of the star are blown off by the shock generated by the collapse of the core. This gravitational explosion is known as a core-collapse supernova (Type II, Ib, Ic). It should be noted that the designations Type II, Ib, Ic are spectroscopic. Type II supernovae are characterized by the presence of hydrogen in their spectra, while Type I’s do not. Of Type I supernovae, Ib’s are characterized by helium in their spectra, while neither helium nor silicon is present in the spectra of Ic’s. The shock also leads to some explosive burning in the inner layers of the original star. For instance, shock heating of the Si-rich layer causes explosive silicon burning, which produces Fe and Ni isotopes. Further, a neutron burst can occur during the explosive helium burning (Meyer et al. 2000). 10 56Fe

9 8

16O

238U

12C

7 6

4He

5 4 3 2 1 0 0

50

100

150

Mass number

200

250

Figure 3. Binding energy per nucleon (experimental values). Data source: Ghahramany et al. (2012).

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As mentioned above, Fe and Ni are the final products of nucleosyntheses during hydrostatic stellar evolution. Because nuclear fusion of elements heavier than Fe is an endothermic reaction, more energetic and explosive nuclear burning is required for synthesizing heavier elements, including most of the siderophile and chalcophile elements. In the following, we introduce three important mechanisms for synthesizing elements heavier than Fe; the s-process, r-process and p-process. The s-process and r-process are neutron capture reactions which play a critical role in producing heavy elements. By contrast, the p-process is a reaction of either proton addition or photodisintegration, which produces minor proton-rich isotopes. The s-process. The s-process (slow neutron capture process) is a chain reaction of neutron capture and β– decay, which proceeds relatively slowly with a timescale of thousands of years per single neutron capture. In the s-process, a seed nucleus with atomic number Z and mass number A captures a neutron and increases the mass number without changing the atomic number:

( Z, A ) + n

→

( Z, A+1)

The next neutron capture occurs in 103–104 years, when the synthesized nucleus, (Z, A+1), is sufficiently stable:

( Z, A+1) + n

→

( Z, A+2 )

If the nucleus (Z, A+2) is radioactive, with a half-life significantly shorter than the timescale of neutron capture, then β– decay occurs to produce a nucleus of higher atomic number, without changing the mass number:

( Z, A+2 ) + n

→

( Z+1, A+2 ) + e – +νe

where νe represents electron antineutrino. Examples of s-process pathways in the region from Mo to Ru, and from W to Os are shown in Figure 4. A box with a solid line refers to stable nuclides, and the number in the box is the Solar System isotopic abundance of the element (Böhlke et al. 2005). As shown in Figure 4, neutron capture repeats from 94Mo to 99Mo, the latter of which which decays to 99Tc with a half-life of 66 h. The pathway is branched at the point of radioactive 99Tc generation, because its half-life is sufficiently long that capture of another neutron is possible to create 100 Tc, while some portion of 99Tc decays to 99Ru. The pathways then meet at 100Ru and neutron capture continues until the formation of the short-lived nuclide 103Ru. In the s-process pathway, 96 Mo and 100Ru are pure s-process nuclides because the formation of these nuclides, via the r-process, is shielded by stable isobaric nuclides of 96Zr and 100Mo (see below). By contrast, there exist some stable nuclides of Mo and Ru that cannot be produced by the s-process. Of such nuclides, proton-rich nuclides of 92Mo, 96Ru, and 98Ru are produced only by the p-process, while the neutron-rich nuclides 100Mo and 104Ru are pure r-process isotopes. The s-process occurs in locations where the neutron density is relatively low such that the neutron capture process repeats every 103–104 years. To date, two stellar environments have been recognized as the sites for two different s-processes; the main s-process and the weak s-process (Käppeler et al. 1991, 2011). The main s-process occurs in the He-burning shells of low-mass AGB stars, where free neutrons are produced via the following reaction to trigger the s-process: 13

C + 4 He → 16O + n

Nucleosynthetic Isotope Variations in the Solar-System

184Os

Z = 76

0.02

186Os

187Os

188Os

189Os

185Re

186Re

187Re

188Re

1.59

Z = 75

37.4

1.96

3.7d

13.2

62.6

16.1

180W

182W

183W

184W

185W

186W

187W

Z = 52

120Te

122Te

123Te

124Te

125Te

126Te

127Te

121Sb

122Sb

123Sb

26.3

2.60

0.10

Z = 51 Z = 50

57.2

112Sn

0.97

114Sn

0.66

115Sn

0.34

96Ru

Z = 44

5.54

14.8

94Mo

9.23

119Sn

120Sn

121Sn

122Sn

98Ru

99Ru

100Ru

101Ru

102Ru

103Ru

104Ru

14.5

7.68

12.8

15.9

96Mo

16.7

24.2

12.6

8.59

17.1

4.2Myr

213ky

99Tc

100Tc

97Mo

98Mo

99Mo

9.56

24.2

32.6

31.6

27h

39.3d

4.63

18.6

7.14

28.6

19.0

191Os

15.4d

192Os

40.8

66h

70m

128Te

31.7

130Te

33.8

124Sn

5.79

126Sn

235ky

p-nuclide s-nuclide

15s

66h

75d

26.3

42.8

118Sn

1.87

95Mo

2.8d

4.82

117Sn

98Tc

92Mo

0.91

30.7

116Sn

Z = 43 Z = 42

14.3

190Os

17h

Z = 74

0.12

115

100Mo

9.67

r-nuclide

Figure 4. s-process pathways around Mo–Tc–Ru, Sn–Sb–Te, and W–Re–Os. Filled, bold, and double

boxes4are nuclides produced dominantly by the p-, s-, and r-processes. Dashed boxes are radioactive nuFig. clides with their half-lives. Numbers in boxes are terrestrial abundances. Broken lines indicate minor path which may occur in specific conditions.

The primary nuclide as the starting material of the s-process in AGB stars is 56Fe, which was inherited from previous generations of stars. The main s-process pathway proceeds to form heavier nuclides and eventually reaches the final cycle, starting from 206Pb as follows: 206

Pb + 3n → 209 Pb → 209 Bi + e – + νe

209

Bi + n → 210 Bi → 210 Po + e – + νe

210

Po → 206 Pb + 4He

In this cycle, 206Pb captures neutrons and produces 209Bi via radioactive 209Pb (T1/2 = 3.3 h). Further neutron capture forms radioactive 210Bi (T1/2 = 5 d) which decays to 210Po, followed by the α decay of 210Po (T1/2 = 138 d) which loops back to 206Pb. Therefore, 209Bi is the heaviest, long-lived nuclide (T1/2 = 1.9 × 1019 y) that can be generated by the s-process, implying that another mechanism is required to produce heavier isotopes, such as 232Th and 238U. By contrast, the weak s-process occurs as a result of He burning in the cores of massive stars, where free neutrons are supplied by the following process (Raiteri et al. 1993; Pignatari et al. 2010): 22

Ne + 4 He → 25 Mg + n

The weak s-process synthesizes s-process nuclides from an iron group seed nucleus up to the mass number A = 90 (Sr, Y). The r-process. The r-process (rapid neutron capture process) is a chain reaction in which successive neutron captures occur faster than the timescale of β-decays of neutron-rich nuclei generated, followed by the cascade of β-decays towards the zone of stable nuclei. The r-process occurs in stellar environments where the neutron density is sufficiently high to allow the

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formation of neutron-rich heavy nuclei via rapid neutron capture. This process is a major source of neutron-rich isotopes of various elements outside the s-process pathway (e.g., 100Mo, 104Ru). Also, the r-process generates long-lived heavy radioactive nuclides of 232Th, 235U, and 238U. The production site of r-process nuclei is still debated. There are two promising sites for r-process synthesis: core-collapse supernovae (ccSN), and neutron star mergers. Of the ccSN models, nucleosynthesis in the neutrino-driven winds from protoneutron stars has been favored by many astrophysicists (e.g., Takahashi et al. 1994; Woosley et al. 1994; Wanajo et al. 2001). However, the neutrino-driven wind models contain the fundamental problem of unacceptable overproduction of some elements (e.g., Sr, Y, and Zr), relative to the Solar System elemental abundances. Moreover, recent numerical simulations have demonstrated that the neutrino-driven wind becomes proton-rich because of the interaction between neutron and neutrino (Fischer et al. 2009). Wanajo (2013) concluded that the neutrino-driven wind from protoneutron stars can be the source of light, trans-iron elements, but not for more than 10% of nuclei heavier than A = ~110 found in the Solar System. In comparison, the merger of two neutron stars in a binary system has been proposed as an alternative scenario for r-process nucleosynthesis (e.g., Freiburghaus et al. 1999). Until recently, the neutron star mergers have not been accepted as a dominant r-process site because of their rarity of occurrence, and because the high yield of r-process elements per event would lead to r-process enrichment that is not consistent with observations of very low metallicity stars (Argast et al. 2004). However, Tsujimoto and Shigeyama (2014) observed a constant [Eu/H] (N.B., bracket represents the logarithm of the ratio of Eu/H value for a star compared to that of the Sun) in faint dwarf-spheroidal (dSph) galaxies, irrespective of their [Fe/H] values, whereas massive dSph galaxies are characterized by an increase of the [Eu/H] as [Fe/H] increases. Note that [Eu/H] and [Fe/H] are indices of r-process/hydrogen ratio and galactic chemical evolution, respectively. This observation implies that the r-process nucleosynthesis did not occur in faint dSph galaxies while supernovae frequently took place, as indicated by the increase of [Fe/H]. This would support neutron star mergers as the main source of r-process nuclides in our galaxy, especially for nuclides with A > 130 (Tsujimoto and Shigeyama 2014). Unlike red giants and supernovae, however, neutron star mergers do not condense dust grains as carriers of the r-process nuclides synthesized (cf. presolar grains). The mechanism how the r-process nuclides, synthesized by the neutron star mergers, were incorporated into the Solar System is still unclear. Therefore, one must keep in mind that the astrophysical site(s) for the r-process is very much an open question that will require further astronomical and theoretical investigation. The p-process. Proton-rich nuclei (e.g., 92Mo, 96Ru, 98Ru, 120Te, 184Os) that depart from the s-process pathway cannot be produced by either the s-process or the r-process. The nucleosynthesis of such proton-rich nuclides is referred to as the p-process, in which three different types of processes occurring in various sites have been proposed; the rp-process, the γ-process, and the νp-process. Rauscher et al. (2013) provide a recent comprehensive review of the p-process. The rp-process (rapid p-process) is a sequential reaction of successive proton capture and subsequent β+ decay. The rp-process requires an extremely proton-rich environment because (γ,p) reactions become faster than proton captures at high temperatures. The rp-process is thought to occur in explosive H- and He-burning on the surface of a mass-accreting neutron star, i.e., X-ray bursts (Schatz et al. 1998). However, there exists a definitive end point of the rp-process around A = 110, which prevents the synthesis of p-nuclei heavier than Te (Schatz et al. 2001). The γ-process is a photodisintegration of pre-existing intermediate nuclei which produces lighter, proton-rich nuclei, either via successive (γ,n) reactions, or via (γ,p) or (γ,α) reactions, followed by β+ decays. The γ-process may have dominated over the rp-process in the

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production of p-nuclides for the Solar System (Rauscher et al. 2013). The primary site of the γ-process is core-collapse supernovae of massive stars, during which the shock front reaches the O/Ne-shell of the star where photodisintegration of preexisting s- and r-process materials occurs (e.g., Woosley and Howard 1978; Rauscher et al. 2002). One problem with the γ-process is that it underproduces the most abundant p-nuclei, 92Mo, 94Mo, 96Ru, and 98Ru, in the Solar System. To account for the deficits of 92,94Mo and 96,98Ru in the Solar System, an additional site for the γ-process, the Type Ia supernovae, has been proposed. Travaglio et al. (2011) explored the calculation of p-process nucleosynthesis for high-resolution two dimensional SNIa models, considering two types of explosions; delayed detonation and pure deflagration. The authors demonstrated that they could produce almost all p-nuclei at the same level as 56 Fe including the debated nuclides 92,94Mo and 96,98Ru, by assuming strong enhancements of s-process seed nuclides synthesized during the recurrence of thermal pulses during the white dwarf mass accretion phase. Travaglio et al. (2011) used a single degenerate (SD) scenario for the occurrence of Type Ia supernovae, in which mass accretion from a companion star onto a white dwarf reaches the Chandrasekhar mass and causes a supernova. On the other hand, a double degenerate (DD) scenario considered that the merger of two white dwarfs causes a supernova. The fraction of SD and DD types in actual type Ia supernovae remains unclear. The third proposed p-process, the νp-process, is a neutrino-driven rp-process which occurs in the innermost ejected layers of a core-collapse supernova (ccSN) when intense neutrino fluxes create a proton-rich environment (Fröhlich et al. 2006; Pruet et al. 2006; Wanajo 2006). The standard rp-process is hindered during the ccSN, due to the existence of a number of “waiting points” such as 64Ge (T1/2 = 1.06 m), where the timescale of β+ decay is sufficiently longer than the timescale of a ccSN (< 1 s). In the νp-process, the problem of waiting points can be resolved by accelerating the reactions via neutron capture and proton release, using neutrons produced by the absorption of antineutrino by protons (Wanajo 2006): p + νe → n + e + 64

Ge + n → 64 Ga + p

The reactions occur within 1 s, so the waiting points are bypassed. The νp-process is a possible candidate for the origin of the solar abundances of 92,94Mo and 96,98Ru that cannot be explained by the γ-process.

PRESOLAR GRAINS Since their discovery in the 1980s, a variety of presolar grains have been documented in primitive chondrites and interplanetary dust particles (IDPs). Different types of presolar grains have diverse isotopic compositions for various elements which are drastically different from one another, evidently indicating that they originated from multiple stellar environments where diverse nucleosyntheses occurred. In the last two decades, significant efforts have been made for direct isotopic measurement of single presolar grains, which has gradually clarified the linkage between individual presolar grains and their nucleosynthetic sources. Another approach to understanding the isotopic characteristics of presolar grains is the analysis of acid resistant residues and leachates extracted from primitive chondrites. The latter method is specifically useful when analyzing trace elements with concentrations of a few ppm in presolar grains. Collectively, these studies of presolar grains have revealed that much of the isotopic variability observed among bulk meteorites is nucleosynthetic in origin, and evidently caused by heterogeneous distribution of presolar grains in the solar nebula, prior to the onset of planetesimal formation. In this section, we review the history of presolar grain studies as well as the types of presolar grains discovered in chondrites and interplanetary dust grains.

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Yokoyama & Walker Historical Background: Discovery and Isotope Analysis

Presolar grains are interstellar dust grains that existed before the formation of the Solar System. Presolar grains reside in primitive meteorites, especially in the matrix of carbonaceous chondrites that have been minimally affected by thermal metamorphism on their parent bodies. In addition to carbonaceous chondrites, presolar grains have also been found in ordinary and enstatite chondrites. Presolar grain abundances comprise typically no more than 0.1% of the mass of an individual meteorite (e.g., Nguyen and Messenger 2011; Zinner 2014). Presolar grains contained in meteorites and interplanetary dust grains are remnants of the ingredients that made up the Solar System. These ingredients formed in a variety of stellar environments. Further, the existence of presolar grains in chondrites implies that these grains have survived dynamical processes in the early solar nebula, such as evaporation, condensation, mixing and collision. Since these grains are labile at high temperature, chondrites that were affected by higher degrees of thermal metamorphism (e.g., petrologic grade greater than 3.6), and differentiated meteorites that experienced melting on their parent bodies (e.g., achondrites, iron meteorites, stony-iron meteorites) no longer carry presolar grains. The existence of presolar materials was predicted before their discovery (Clayton 1978). As summarized in the previous section, current nucleosynthetic models suggest all elements heavier than He were synthesized by a variety of stellar environments as stellar evolution in the universe proceeded. Since nucleosynthesis in different stellar environments would produce elements with different isotopic compositions, it was projected that such putative “presolar components” should have isotopic compositions drastically different from those of terrestrial composition (Clayton 1978). However, due to a lack of evidence for resolvable nucleosynthetic anomalies among bulk planetary materials, it was originally concluded that the inner part of the early solar system experienced sufficiently high temperatures that all solid materials were evaporated and ultimately mixed, such that the isotopic compositions of all elements in accessible planetary materials were thoroughly homogenized. The improvement of mass spectrometers and associated techniques in the 1960–70s led to the discovery of isotopic anomalies in meteorites and their components (Reynolds and Turner 1964; Black and Pepin 1969; Lewis et al. 1975; McCulloch and Wasserburg 1978a,b; Papanastassiou and Wasserburg 1978). Because these measurements were made on chemically processed, bulk planetary materials, however, these studies were incapable of identifying the carriers of the isotopically anomalous components. Finally, in the 1980s, the group of E. Anders from the University of Chicago examined the residues of stepwise acid leaching from primitive chondrites, and measured their rare gas isotopic compositions. They discovered that the fraction showing an anomalous Xe isotope composition (Xe-HL) is composed of acid insoluble, presolar nanodiamonds. This is considered the first discovery of presolar grains (Lewis et al. 1987). Subsequently, the group discovered presolar silicon carbide (carrier of Xe-S and Ne-E(H)), and presolar graphite (carrier of Ne-E(L)) (Bernatowicz et al. 1987; Amari et al. 1990). The study of individual presolar grains has dramatically improved since the 1980s as a result of the development of secondary ion mass spectrometry (SIMS), which enables direct isotopic analysis of an area of 10 μm diameter or less. In SIMS analysis, the sample of interest is put in a sample chamber that is under vacuum, without prior chemical separation. The targeted area for isotopic analysis is sputtered by a focused primary beam (e.g., Cs+, O–) to generate secondary ions that are focused and separated by their mass/charge ratios, using the electrostatic analyzer and magnetic sector that follow the ion source. The analytical targets for SIMS measurements are typically major element constituents of presolar grains, such as C, O, and Si, as well as some minor elements lighter than Fe (e.g., N, Mg, Ca, and Ti). The analytical uncertainties depend on various conditions; however, those of major elements are normally less than a few % in general.

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A major drawback of SIMS analysis is the difficulty of isotope analysis for elements heavier than Fe. Resonance ionization mass spectrometry (RIMS) is an alternative analytical technique which has played an important role in obtaining the isotopic compositions of some heavy elements in presolar grains, including some siderophile elements (e.g., Mo, Ru: Nicolussi et al. 1998a,b; Savina et al. 2004). In RIMS analysis, a sample volume is sputtered using a conventional ion probe gun. The cloud of neutral and charged particles resulting from the sputtering is then interrogated by irradiating the cloud with laser beams with wavelengths tuned to achieve the excitation energy level necessary to selectively photoionize the target element. The isotopic composition of the element is then measured by time of flight mass spectrometry. This method enables selective ionization of the target element without chemical separation from unwanted elements which, if ionized, could result in isobaric interferences with the element of interest. (e.g., CHARISMA, Argonne National Laboratory: Savina et al. 2003).

Types of Presolar Grains and Their Origin Table 1 summarizes the types of presolar grains discovered in meteorites and interplanetary dust particles (IDPs) to date. Microscopic images of representative presolar grains are shown in Figure 5. In the following, we briefly review the nucleosynthetic origin of each type of presolar grain. For the isotope data presented in some Figures, we take data from the database for presolar grains created by (Hynes and Gyngard 2009). Table 1. Abundances, sizes, and stellar sources of presolar grains. Grain type

"Abundance (ppm)"

"Size (µm)"

Stellar sources

1000

0.002

SNe

SiC

40

0.1–20

Graphite

2

1–20

SNe, AGB

100

0.1–3

RG, AGB, SNe

15000

0.1–1

RG, AGB, SNe

200

0.1–1

RG, AGB

0.002

0.3–1

SNe

Nanodiamonds

Oxides Silicates (IDPs) Silicates (meteorites) Si3N4

AGB, SNe, Novae, J-stars

Presolar grain abundances vary with meteorite type.

Nanodiamond. As described previously, nanodiamonds (Fig. 5a) were the first type of presolar grain discovered in chondrites. The existence of nanodiamond was confirmed by using X-ray diffraction (XRD) and transmission electron microscopy (TEM) analysis. The abundance of nanodiamonds in chondrites can be as much as 0.1% of the mass. Direct isotope analysis of single nanodiamond grains is currently impossible because of their minuscule grain size. Therefore, isotope compositions of nanodiamond grains have been obtained by group analysis of nanodiamond-enriched acid residues extracted from chondrites. However, the origin of this type of presolar grain is still controversial. As described above, rare gas isotope results on the nanodiamond-enriched fractions provided the first evidence that at least some portion of the nanodiamonds present are interstellar dust grains, which record the collective isotopic compositions of the materials present in their formation environment (Lewis et al. 1987). Excesses of Te and Pd isotopes (128Te, 130Te, and 110Pd) have also been reported for nanodiamond-enriched fractions extracted from a carbonaceous chondrite (Richter et al. 1998; Maas et al. 2001). These characteristics can be produced by an intense neutron-burst when a supernova shock heats the He-burning shell and liberates neutrons via 22Ne(α,n)25Mg (Meyer et al. 2000). More recently, Stroud et al. (2011) analyzed nanodiamond separates from the Allende and Murchison meteorites

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2.5 nm




1µ µm







1µ µm

0.1µ µm

Figure 5. Images of presolar grains. (a) nanodiamond (b) SiC (c) graphite, and (d) corundum. Used by the following permission. (a) Tyrone L. Daulton, Washington University in St. Louis (b) Scott Messenger, NASA (c) Sachiko Amari, Washington University in St. Louis (d) Aki Takigawa, Kyoto University.

using electron microscopy with sub-nanometer resolution. The authors demonstrated that the separates were a two-phase mixture of nanodiamonds and glassy carbon, which was most likely the product of supernova shock-wave transformation of pre-formed organics in the interstellar medium. By contrast, Dai et al. (2002) found that fragile, carbon-rich IDPs of cometary origin are nearly free of nanodiamonds, suggesting that nanodiamonds in meteorites formed within Fig.inner 5 Solar System and are not presolar. The inconsistency may imply that the majority the of nanodiamonds originated in the Solar System. Nevertheless, at least a small proportion of nanodiamonds are undoubtedly presolar and record nucleosynthetic anomalies. Silicon carbide. The maximum abundance of SiC in chondrites is 30–40 ppm (Davidson et al. 2014). SiC (Fig. 5b) is strongly acid resistance, however, so that the separation of presolar SiC from other meteorite components is possible by a stepwise acid leaching procedure. Since SiC does not form by nebular processes, SiC grains observed in chondrites are all presumed to be presolar in origin. Typical SiC grains occur at a μm scale. Therefore, it is possible to perform direct isotopic analysis of this type of grain using SIMS. Consequently, SiC is the most studied presolar grain among all types of presolar materials. Comprehensive analysis of Si, C, and N isotope compositions in presolar SiC grains has revealed that presolar grains can be divided into several groups of different nucleosynthetic origin. More than 90% of presolar SiC is categorized as mainstream SiC, and the rest are separated into subgroups of A + B, X, Y, Z, and Nova. The mainstream SiC grains are characterized by higher 29,30Si/28Si and 14N/15N ratios, and lower 12C/13C ratios, compared to those of terrestrial values (Figs. 6–7). Carbon isotope ratios, as revealed by the analysis of mainstream SiC grains, are consistent with the 12C/13C ratio observed in carbon stars, which are AGB stars as discussed in earlier section (Lambert et al. 1986). In addition, an infrared emission feature around 11.3 μm observed in the outflows of carbon stars points to the existence of SiC particles (Speck et al. 1997). Based on these observations, and the comparison of C and N isotopic compositions in SiC with nucleosynthetic theories, the

Nucleosynthetic Isotope Variations in the Solar-System

121

200

100

0

200

-100 -100

Mainstream, A+B 0

100

200

300

δ29Si/28Si

0

-200

-400

Mainstream X Y Z A+B nova

-600

-800 -800

-400

0

400

δ30Si/28Si

800

1200

1600

Figure 6. Silicon isotopic compositions of presolar silicon carbide grains. The presolar grain database by Hynes and Gyngard (2009) was used to plot the data. The original sources of representative data are Alexander (1993), Alexander and Nittler (1999), Amari et al. (2001a, 2001b, 2001c), Barzyk et al. (2007), Besmehn and Hoppe (2003), Hoppe et al. (1994, 1996), Huss et al. (1997), Lin et al. (2002), Marhas et al. (2008), Nittler and Alexander (2003), Virag et al. (1992), and Zinner et al. (2003, 2007). The full references for the entire data suite are available in Hynes and Gyngard (2009).

mainstream SiC grains, as well as Y and Z grains, are concluded to originate from low-mass Fig. 6 carbon stars (1–3 M) (e.g., Zinner et al. 1989; Alexander 1993; Hoppe et al. 1994; Nittler and Alexander 2003). Isotope analyses of heavy elements in individual mainstream SiC grains using RIMS or SIMS discovered large excesses of s-process nuclides, such as 86,87Sr, 96Mo, 90,91,92,94 Zr, 100Ru, and 134,136Ba (Fig. 8) (Nicolussi et al. 1998a; Savina et al. 2004; Barzyk et al. 2007; Liu et al. 2014). Likewise, enrichments of s-process nuclides in chondritic acid residues containing presolar SiC grains were observed for a variety of elements, including the siderophile elements Mo, Ru, W, and Os (Dauphas et al. 2002b; Yokoyama et al. 2007, 2010, 2011; Burkhardt et al. 2011, 2012a). These results are consistent with the nucleosynthetic theory that s-process occurs in the He-burning shell of low-mass AGB stars. Type X SiC grains are rare, comprising only ~1% of the abundance of all SiC grains analyzed. These grains can be distinguished from the mainstream grains with their lower 29,30Si/28Si and 14 N/15N ratios and higher 12C/13C ratios (Figs. 6–7). The origin of X grains is most likely the

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Mainstream X Y Z A+B nova

10000

14N/15N

1000

Figure 7. Carbon and nitrogen isotopic compositions of presolar silicon carbide grains. Dashed lines are those of solar system average. Data are from the same sources as for Figure 6.

100

10

1

(a)

1

10

100

1000

12C/13C

200

10000

SiC Maingroup

0

δiMo

‐200

Fig. 7‐400 ‐600 ‐800 ‐1000 91

92

93

94

95

96

iMo

97

98

(b)1600

99

100

101

SiC X‐grains

1400

Figure 8. Representative data for Mo isotope compositions in (a) mainstream SiC grains and (b) type X SiC grains. Data are from Barzyk et al. (2007) and Pellin et al. (2006).

1200

δiMo

1000 800 600 400 200 0 ‐200 ‐400 91

92

93

94

95

96

iMo

97

98

99

100

101

ejecta of core collapse supernovae, because some X grains are enriched in 44Ca and 49Ti, both of which are daughter nuclides of short-lived radioactive nuclides synthesized only in ccSNe 49 Ti and V, respectively). By contrast, RIMS measurements of X grains discovered excesses (44Fig. 8 95 of Mo and 97Mo (Fig. 8) (Pellin et al. 2006). Interestingly, the Mo isotopic pattern observed in the X grains does not match any patterns produced by conventional three nucleosynthetic models (s-, r-, and p-processes). Rather, the pattern is consistent with a formation mechanism

Nucleosynthetic Isotope Variations in the Solar-System

123

via the neutron burst that takes place in a He-shell during passage of the supernova shockwave (Meyer et al. 2000; Rauscher et al. 2013). The importance of these X-grain measurements is that they show the existence of additional nucleosynthetic components, other than those synthesized via the s-, r-, and p-processes, the main nucleosynthetic sources for trans-iron elements. Such components may not play a significant role in the overall composition of the Solar System, but would be present in individual presolar carriers, and thus, could influence isotopic variations in early Solar System materials. The neutron burst model also accounts for the isotopic composition of Te and Pd in nanodiamonds (see above). Isotopic analyses of the other minor types of SiC grains are limited. The Nova grains, originated from the explosive H burning of novae, have lower 12C/13C and 14N/15N ratios and higher 30Si/28Si and 26Al/27Al ratios relative to the terrestrial values (Figs. 6–8) (Amari et al. 2001a). The type A + B grains possess Si isotope compositions similar to the mainstream grains but have extremely low 12C/13C ratios and variable 14N/15N ratios (Amari et al. 2001b). These isotopic signatures cannot be explained by nucleosynthesis in normal AGB stars. Amari et al. (2001b) proposed that a possible source of A + B grains with enhanced s-process elemental abundances was born-again AGB stars, which are post AGB stars in late thermal pulse phase due to the re-ignition of the He shell surrounding the C/O core (Sakurai’s object). An alternative source of the A + B grains are the J-type carbon stars, which have extremely low 12C/13C ratios, although the details of such a model are still unclear (Amari et al. 2001b). In addition, excesses of p-process isotopes 92,94Mo and 96,98Ru were reported in a type B SiC grain (Savina et al. 2007), which would require at least one more nucleosynthetic source for this type of grain. Graphite. The abundance of presolar graphite in chondrites (Fig. 5c) reaches a maximum of ~2 ppm. The grain size of the presolar graphite ranges from 1 to 20 μm. The density of presolar graphite is bimodal; the more grain sizes increase, the more the density of grains decreases. Because of their distinctive isotopic characteristics, the stellar origin of the low density (LD) graphite and high density (HD) graphite is thought to be different from each other. The 12C/13C ratios of HD graphite grains range from 2 to 4000 with bimodal peaks around C/13C = 10 and 300–400 (Fig. 9). Amari et al. (2014) investigated presolar graphite grains separated from the Murchison meteorite and concluded that HD grains of 2.10–2.15 g·cm–3 with 12 13 C/ C ≥ 100 formed in the outflow of low-mass (1.5–3 M) and low-metallicity AGB stars, and those of 2.15–2.20 g·cm–3, with 12C/13C ≥ 60 formed in the same stars, as well as in 5M and solar to half-solar metallicities. Such environments are markedly different from the site where the mainstream SiC grains formed. In contrast, the HD grains with 12C/13C ≤ 20 have multiple origins including the ejecta of ccSN, J stars, and born-again AGB stars. Another intriguing signature is that some of the HD graphite grains contain tiny refractory carbide grains (e.g., ZrC, MoC, and RuC) at the center as a subgrain (Bernatowicz et al. 1996; Croat et al. 2005). The sub-grains evidently existed prior to the crystallization of graphite grains, which provide a clue to understanding the formation process of presolar grains in the stellar envelope. Because Zr, Mo, and Ru are dominantly s-process elements, the existence of refractory subgrains supports the AGB origin of the HD graphite grains. Nicolussi et al. (1998b) measured Zr and Mo isotope compositions in 32 individual HD graphite grains from the Murchison meteorite using RIMS. Although most of the grains showed close-to-terrestrial Mo isotopic compositions, five grains had Mo isotopic patterns with excess s-process nucleosynthetic signatures. In addition, three out of eight graphite grains for which Zr and Mo isotope compositions were measured simultaneously presented correlated s-process isotopic characteristics for both Zr and Mo, suggesting low-mass, thermally pulsed AGB stars as their origin. 12

On the other hand, the LD graphite grains ( 2.05 g·cm–3) possess trace element abundances higher than HD graphite. LD graphite is characterized by excesses of 15N, 18O, 28Si, as well as high 26Al/27Al ratios. The distribution of 12C/13C shift to lower ratios compared to HD graphite, with a peak around 12C/13C = 90–200. These isotopic signatures resemble those of SiC X

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High density Low density

16O/18O

1000

100

10

1

10

1

10

100

1000

10000

100

1000

10000

12C/13C

14N/15N

1000

100

10

12C/13C

Figure 9. Carbon, nitrogen, and oxygen isotopic compositions of presolar graphite grains. Dashed lines are those of solar system average. Fig. 9 The presolar grain database by Hynes and Gyngard (2009) was used to plot the data. The original sources of representative data are Amari et al. (1993), Hoppe et al. (1995), and Jadhav et al. (2013). The full references for the entire data are available in Hynes and Gyngard (2009).

grains, the origin of which is most like ccSNe. The ccSN origin of LD graphite is supported by excesses of daughter nuclides of the short-lived radionuclides 44Ti and 41Ca. These nuclides are produced by neutron capture in the C/O-shell of type II supernovae, although 44Ti and 41Ca are also produced by the oxygen and silicon burning, and by the weak s-process in convective core helium burning, respectively. Oxides and Silicate. Presolar corundum (Al2O3) is the grain first discovered as a carbon-free mineral (Fig. 5d). Presolar oxides identified other than corundum are spinel (MgAl2O4), hibonite (CaAl12O19), TiO2, FeO, and FeCr2O4. The maximum abundance of presolar oxides in chondrites is ~100 ppm. The grain size of presolar oxides ranges from 0.1 to 3 μm in diameter. Unlike SiC, oxide grains can be formed by a variety of nebular processes and are abundant in the Solar System. Therefore, the only way to identify presolar oxides in chondrites is the direct isotope analysis for oxygen using SIMS. Identification of presolar silicates has been hampered by the fact that silicates are not acid resistant, especially to HF. In addition, silicates are the most abundant types of minerals in the Solar System that formed in the nebula. The first discovery of presolar silicates was in interplanetary dust particles (Messenger et al. 2003). The abundance of presolar silicates in IDPs reaches ~1%, which is the greatest among all types of presolar grains. Subsequently, presolar

Nucleosynthetic Isotope Variations in the Solar-System

125

silicates were discovered in chondrites by SIMS analysis (Nagashima et al. 2004; Nguyen and Zinner 2004). The abundance of presolar silicates in chondrites is a maximum of ~ 250 ppm. The presolar silicates that have been identified include forsterite, enstatite, Ca–Al-rich phases, and non-stoichiometric amorphous silicates with grain size ranging from 0.1–1 μm. Based on O isotopic compositions, Nittler et al. (1997) categorized presolar oxide grains into four groups with different stellar origins. The same groupings are applied to presolar silicate grains. As shown in Figure 10, Group 1 grains are enriched in 17O and have 18O/16O ratios slightly lower than the Solar System average. This is consistent with those observed in red giant and AGB stars. Group 2 grains have significant deficits in 18O (18O/16O < 0.001) with 17 O/16O ratios slightly higher than the Solar System average. The deficits of 18O are difficult to explain with first and second dredge-ups in red giant stars, however, the mixing of materials from the envelope into the H-burning shell (cool bottom processing) that occurs in low mass (< 1.65 M) AGB stars may explain the observed isotopic compositions. Group 3 has 17O/16O and 18O/16O ratios lower than the Solar System average. This can be explained by the gradual

1E‐1

Group 1

1E‐2

17O/16O

oxide‐gr1 oxide‐gr2 oxide‐gr3 oxide‐gr4 oxide‐ung silicate‐gr1 silicate‐gr2 silicate‐gr3 silicate‐gr4 silicate‐ung

Group 2

1E‐3

Group 4

Group 3

1E‐4

L2011B10 7-5-7

T84

1E‐5 1E‐5

1E‐4

1E‐3

18O/16O

1E‐2

1E‐1

Figure 10. Oxygen isotopic compositions of presolar oxide and silicate grains. Dashed lines are those of solar system average. The presolar grain database by Hynes and Gyngard (2009) was used to plot the data. The original sources of representative data are Bose et al. (2010), Choi et al. (1998, 1999), Floss and Stadermann (2009), Gyngard et al. (2010), Mostefaoui and Hoppe, (2004), Nagashima et al. (2004), Nguyen et al. (2003), Nguyen and Zinner (2004), Nguyen et al. (2007, 2010), Nittler et al. (1997, 2008), Vollmer et al. (2009a, 2009b), and Zinner et al. (2003, 2005). The full references for the entire data are available in Hynes and Gyngard (2009).

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increase of secondary nuclides 17O and 18O, relative to the primary nuclide 16O, along with galactic chemical evolution (GCE), indicating that Group 3 grains originated from low mass and low metallicity stars. The origin of Group 4 grains has been debated. Although these grains are enriched in 17O and 18O relative to 16O, they do not show excesses of 29Si and 30Si relative to 28Si, which is inconsistent with the GCE model. An alternative interpretation for the origin of Group 4 grains is a single supernova (Choi et al. 1998; Nittler et al. 2008). To date, no isotopic data are available for trace elements heavier than iron in presolar oxide and silicate grains. New generations of mass spectrometers with extremely high sensitivity will be required to shed light on the isotopic characteristics of trace elements, including siderophile and chalcophile elements, in this group of presolar grains. Silicon Nitride. Silicon nitride (Si3N4) is one of the rarest presolar grains. The isotopic characteristics of Si3N4 grains resemble those of SiC X grains, suggesting a supernova origin. However, one Si3N4 grain was found in the enstatite chondrite Indarch. It was characterized by excesses of 13C and 14N, and likely originated from an AGB star (Zinner et al. 2007).

ISOTOPE ANOMALIES OF SIDEROPHILE AND CHALCOPHILE ELEMENTS IN BULK METEORITES The extent and causes of isotopic heterogeneity in the early Solar System have been long-standing issues since the establishment of nucleosynthetic models and the discovery of presolar grains. The existence of presolar grains in chondrites and interplanetary dust particles (IDPs) implies that the original dust grains present in the solar nebula were not completely evaporated before the onset of planetesimal formation. Heterogeneous distribution of isotopically anomalous presolar grains in the solar nebula would cause isotopic anomalies among planetary bodies that formed separately in time and space. Therefore, investigating nucleosynthetic isotopic anomalies in a variety of meteorites is of great importance for understanding large-scale material transport and subsequent nebular and planetary processes. In the following section, we review recent achievements in the precise measurement of the isotopic compositions of siderophile and chalcophile elements in bulk aliquots of meteorites, including chondrites and differentiated meteorites. We specifically focus on transiron elements which are generally synthesized by the stellar nucleosyntheses of the s-, r-, and p-processes. A specific emphasis is made here on studies after ~2000, in which high precision isotope measurements were employed with analytical uncertainties of epsilon level (part per 10,000) or better. Such high precision isotope analyses are not available for all siderophile and chalcophile elements, and are presently limited to Mo, Ru, Te, W, and Os. The variation of isotope compositions for these elements in bulk meteorites are summarized in Table 2 for representative references. We will review observations regarding these elements, and also highlight advancing work on Pt and Cd isotopes. Note that in the study of nucleosynthetic isotope anomalies in extraterrestrial materials, the magnitude of an isotopic anomaly is commonly expressed as ε or μ notation;  ( i M / j M)sample  ε iM = − 1 × 10 4  i j  ( M / M)standard 

(1)

i j  ( M / M)sample  µ iM = − 1 × 106  i j ( M / M)  standard 

(2)

where i M and j M are target and reference isotopes of an element M, respectively.

Table 2. Variation of isotope compositions for Mo, Ru, Te, W, and Os in bulk meteorites and terrestrial standards. Chondrites

Non-chondrites

Terre -strial Std

Carbonaceous

Ordinary

Enstatite

Achondrites

Irons

Stony irons

min–max

min– max

min–max

min–max

min–max

min–max

± 2s

–0.14 to 0.52

–0.23 to 1.99

1.06 to 1.14

0 ± 0.75 0 ± 0.50

Molybdenum ε92Mo/96Mo

1.12 to 6.44

0.58 to 0.94

0.36 to 0.65

ε Mo/ Mo

0.00 to 4.82

0.43 to 1.01

0.30 to 0.61

0.09 to 0.69

–0.31 to 1.42

0.85 to 0.85

ε95Mo/96Mo

0.53 to 3.17

0.12 to 0.38

0.18 to 0.19

–0.12 to 0.06

–0.19 to 1.02

0.38 to 0.80

0 ± 0.41

ε97Mo/96Mo

0.12 to 1.66

0.09 to 0.18

0.11 to 0.14

0.02 to 0.20

–0.15 to 0.53

0.16 to 0.41

0 ± 0.23

ε100Mo/96Mo

0.35 to 2.28

0.04 to 0.37

0.12 to 0.31

–0.14 to 0.40

0.00 to 0.91

0.20 to 0.79

0 ± 0.44

[1]

[1]

[1]

[1]

[1]

[1]

[1] 0 ± 1.19

94

96

References

Ruthenium ε96Ru/101Ru

0.13 to 1.41

–0.01 to 1.38

–0.58 to 0.82

0.2 to 0.45

ε98Ru/101Ru

0.34 to 2.94

–0.84 to 2.44

–0.42 to 1.79

0.25 to 1.56

0 ± 1.98

ε100Ru/101Ru

–1.70 to 0.60

–0.26 to 0.44

–1.08 to 0.00

–0.58 to –0.36

0 ± 0.31

ε102Ru/101Ru

–0.29 to 0.81

–0.43 to 0.45

–0.53 to –0.05

–0.03 to 0.15

0 ± 0.64

ε104Ru/101Ru

0.58 to 1.65

–0.45 to 0.55

–0.44 to 0.52

–0.19 to 0.37

0 ± 0.58

References

[2]

[2]

[2]

[2]

[2]

ε100Ru/101Ru

–3.37 to –1.00

–0.30

–1.1 to –0.33

ε102Ru/101Ru

–1.25 to –0.39

–0.14

–0.462 to –0.04

References

[3]

[3]

[3] Tellurium

ε120Te/128Te

–26 to 39

–49 to 16

2

–39 to 49

0 ± 45

ε122Te/128Te

–2.2 to 1.9

–2.3 to –0.5

–1.4

–1.2 to 4.5

0 ± 1.4

ε124Te/128Te

–1.0 to 0.3

–1.9 to –0.4

–0.10

–0.5 to 1.8

0 ± 1.0

ε126Te/128Te

–0.3 to 0.4

–0.4 to 0.0

–0.3 to 0.0

–0.3 to 0.3

0 ± 0.3

ε130Te/128Te

–0.4 to 0.9

0.3 to 0.9

0.5 to 0.9

–1.5 to 0.7

0 ± 0.6

References

[4]

[4]

[4]

[5]

[4]

ε183W/184W

–0.47 to 0.54

–0.43 to 0.43

–0.16 to 0.09

0 ± 0.77

References

[6]

[6]

[6]

[6]

–0.15 to 0.01

0 ± 0.05

[7]

[7]

Tungsten

ε184W/183W References Osmium ε186Osi/189Os*

–0.73 to 0.49

–0.07 to 0.18

–0.39 to 0.33

0.03

ε188Os/189Os

–0.17 to 0.10

0.00 to 0.13

–0.16 to 0.23

0.06

0 ± 0.07

ε190Os/189Os

–0.09 to 0.07

–0.03 to 0.04

–0.03 to 0.03

–0.07

0 ± 0.10

[8,9]

[8,9]

[8,9]

[10]

[9]

References

2 ± 0.36

ε Os / Os

–0.06 to 0.79

–0.12

2 ± 0.25

ε188Os/188Os

–0.55 to 0.06

–0.17 to 0.03

0 ± 0.17

ε190Os/188Os

–0.03 to 0.31

–0.10 to 0.11

0 ± 0.13

References

[11]**

[11]**

[11]

186

i 188

*

References: [1]Burkhardt et al. (2011) [2] Chen et al. (2010) [3] Fischer-Gödde et al. (2013) [4] Fehr et al. (2005) [5] Fehr et al. (2004) [6] Kleine et al. (2004) [7] Qin et al. (2008) [8] Yokoyama et al. (2007) [9] Yokoyama et al. (2010) [10] van Acken (2011) [11] Walker (2012). * ε186Osi/188,189Os are the part per 104 deviation of the calculated initial 186Os/188,189Os of meteorites from the initial ratios of the solar system. Note that ε186Osi/188,189Os of the terrestrial standard are not identical to zero because the 186Os/188,189Os of the standard used (UMd Os) are elevated relative to chondrites. ** Os isotope data for iron meteorites and pallasites reported in Walker (2012) contain samples affected by long-term cosmic ray exposures.

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Isotope anomalies of siderophile elements in bulk meteorites Molybdenum. Molybdenum is a moderately siderophile element with a CI chondrite abundance of 1.02 ± 0.11 µg/g (1s) (Lodders 2003). Most meteorites, excluding some achondrites, have Mo abundances that are substantially higher than that of CI chondrites. In addition, recent development of the negative-thermal ionization mass spectrometry (N-TIMS) technique, and refinement of multi-collector-inductively coupled plasma mass spectrometry (MC-ICP-MS) techniques have dramatically reduced the amount of Mo necessary for performing high-precision isotope analysis compared to earlier studies utilizing positive-thermal ionization mass spectrometry (P-TIMS) techniques, making this element highly useful for the study of nucleosynthetic isotopic anomalies in meteorites.

εiMo

Molybdenum has seven stable isotopes 92Mo, 94Mo, 95Mo, 96Mo, 97Mo, 98Mo, and 100Mo, with averaged terrestrial abundances of 14.8%, 9.23%, 15.9%, 16.7%, 9.56%, 24.2%, and 9.67%, respectively (Böhlke et al. 2005). Molybdenum isotopes are synthesized via stellar nucleosyntheses of the s-process (trace 94Mo, 95Mo, 96Mo, 97Mo, 98Mo), the r-process (95Mo, 97 Mo, 98Mo, 100Mo) and the p-process (92Mo, 94 Mo). In addition, the p-process nuclide 97 Tc decays by electron capture to 97Mo, (a) 1.4 s‐deficit 1.2 with a half-life of 2.6 × 106 yr. Because 1.0 four out of seven Mo isotopes are produced 0.8 almost entirely by a single nucleosynthetic 0.6 process (92,94Mo: p-process, 96Mo: s-process, 100 0.4 Mo: r-process), the excess or deficit 0.2 of a specific nucleosynthetic component, 0.0 relative to the terrestrial component, makes ‐0.2 for a distinctive Mo isotopic composition 91 92 93 94 95 96 97 98 99 100 101 iMo which is useful for diagnosing the origin of (b) 1.4 Mo isotope anomalies in meteorites. r‐excess 1.2 1.0

εiMo

0.8 0.6 0.4 0.2 0.0 ‐0.2 91

(c)

92

93

94

95

1.4

96

iMo

97

98

99

100

101

p‐excess

1.2 1.0

εiMo

0.8 0.6 0.4 0.2 0.0 ‐0.2 91

92

93

94

95

96

iMo

97

98

99

100

101

Fig. 11 Figure 11. Molybdenum isotope patterns representing (a) s-deficit, (b) r-excess, and (c) p-excess relative to the terrestrial component. The model patterns were calculated by subtracting an s-process endmember component (Arlandini et al. 1999) from the solar system composition and scaled to a mixing ratio of (p + r)/s yielding ε92Mo = 1.

The pattern of εMo values for cases representing an excess or deficit of a pure nucleosynthetic component relative, to the terrestrial component, is shown in Figure 11. Note that ε96Mo and ε98Mo values are defined as zero because, for this diagram, the data acquired by mass spectrometry techniques are corrected for massdependent fractionation during isotopic measurement using 98Mo/96Mo = 1.453171 (Lu and Masuda 1994). The excess of a pure p-process component affects only ε92Mo and ε94Mo, while the deficit of a pure s-process component results in positive εMo values, excluding ε96Mo and ε98Mo, generating a W-shaped pattern (Fig. 11). By contrast, the excess of a pure r-process component produces a pattern akin to the case of s-process deficit, although a characteristic kink is observed in ε94Mo. From the Mo isotopic patterns found in meteorites, therefore, it is theoretically possible to resolve the proportions of individual nucleosynthetic components within a sample.

Nucleosynthetic Isotope Variations in the Solar-System

129

Review of three early studies of Mo isotope anomalies in meteorites (Dauphas et al. 2002a; Yin et al. 2002; Becker and Walker 2003a) is useful for highlighting the issues that must be considered when examining meteorites for nucleosynthetic isotope anomalies. Each of these studies reached different conclusions regarding the isotopic composition of Mo in meteorites. Yin et al. (2002) used N-TIMS to determine the Mo isotope compositions (excluding 92Mo) in bulk aliquots of two carbonaceous chondrites, Murchison (CM2) and Allende (CV3). Dauphas et al. (2002a) utilized MC-ICP-MS for determining the Mo isotope composition of Allende. Both studies obtained W-shaped isotopic patterns (Fig. 12a) which resemble the case for s-process deficit (Fig. 11). Notably, the patterns they reported for Allende generally match the model pattern for p- and r-process excesses. By contrast, Becker and Walker (2003a) measured Mo isotope compositions in Allende and the H4 ordinary chondrite Forest Vail using N-TIMS, and found no resolvable Mo isotope anomalies. The inconsistencies between these studies were also evident for iron meteorites. Yin et al. (2002) analyzed five iron meteorites from four groups (IAB, IIAB, IIIAB, and IVB), and observed normal terrestrial Mo isotopic compositions for each. Similarly, Becker and Walker (2003a) observed no Mo isotope anomalies in two IIAB irons. In contrast to these two studies, Dauphas et al. (2002a) measured Mo isotopic compositions in fourteen iron meteorites from various groups (IAB-IIICD, IIAB, IIE, IIIAB, IVA, IVB, and ungrouped), as well as one mesosiderite and two pallasites, and detected anomalies in all of these meteorites. The anomalous patterns were similar to, but of lesser magnitude than observed in Allende (Fig. 12b). As with the study of isotopic anomalies of any element in bulk planetary materials, at least three fundamental questions regarding the apparent Mo isotope anomalies must be addressed.

Figure 12. (a) Molybdenum isotope patterns for bulk aliquots of chondrites. Data sources: [1] Dauphas et al. (2002a), [2] Yin et al. (2002), [3] Becker and Walker (2003a). For [1] and [2], error bars are uncertainties of individual measurements estimated by 2 SE (standard error) of multiple lines of data acquisition in a single isotopic measurement. For [3], error bars are 2 SE of multiple data obtained by different isotopic measurements. (b) Mo isotope anomalies for iron meteorites (Dauphas et al. 2002a). Error bars are 2 SE of multiple isotopic measurements.

Fig. 12

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First, are the anomalies reflective of the sample processed, or the result of an analytical artifact? For example, mass independent fractionation during TIMS/ICP-MS measurements has been well documented, as described below. Numerous factors can cause anomalous fractionation, which is a major concern for the determination of precise isotope ratios. For TIMS analysis, temperature gradients across the ionizing filament can lead to the generation of multiple sample domains on the filament that fractionate at different rates, resulting in the mixing of isotopically distinct reservoirs in the detectors. Such an isotope mixing from vigorously and weakly fractionated reservoirs produces isotopic compositions that deviate from the exponential law curve that is commonly used for correcting mass fractionation (Upadhyay et al. 2008). In the case of MC-ICP-MS analysis, the cross section of ion beams generated by plasma ionization is so broad, relative to thermal ionization, that the beams may be partially clipped in the flight tube and cause mass independent fractionation (Albarède et al. 2004). Furthermore, an odd–even isotope separation, dependent on the shape of the sampler and skimmer cones, was reported for precise W isotope analysis using MC-ICP-MS (Shirai and Humayun 2011). Most of these problems are not reproducible and are difficult to evaluate with the mass spectrometers used in the studies noted above. Analytical artifacts can also result from isobaric interferences from other elemental or molecular species that are ionized along with the element of interest, usually as a result of inadequate chemical purification. Second, if the measurements are accurate and precise to the level of precision reported, do anomalous isotopic compositions reflect the incomplete digestion of the sample, with complementary anomalies residing in materials that remained un-accessed by the dissolution method? For instance, presolar SiC is very acid-resistant, so incomplete dissolution of s-processenriched mainstream presolar SiC grains would result in an apparent Mo isotope pattern for a bulk meteorite with an s-process deficit (W-shaped), assuming the true bulk sample had a terrestrial Mo isotopic composition (no anomalies). This is a very real possibility for some elements in some meteorites that contain relatively abundant presolar SiC, because dissolution of SiC using normal acids, especially in non-pressurized digestion vessels, is quite limited. In the case of Murchison which contains 9 ppm of presolar SiC (Huss et al. 2003), the incomplete digestion of SiC could result in ε94Mo and ε100Mo values of ~ + 0.45, if it is assumed that bulk Murchison has no Mo isotopic anomalies. By contrast, Allende contains very little SiC (0.01 ppm) due to the destruction of presolar grains by thermal metamorphism on the parent body (Huss et al. 2003). Consequently, anomalies of the magnitude reported by Yin et al. (2002) and Dauphas et al. (2002a) for this meteorite are unlikely to reflect incomplete sample dissolution. It should also be noted that iron meteorites solidified from liquid metal at conditions under which presolar grains cannot survive. Hence, anomalies present in iron meteorites, as well as the metal portions of stony irons, cannot be explained as a result of incomplete dissolution. As a corollary, because individual members of a so called “magmatic” iron meteorite group (e.g., IIAB, IID, IIIAB, IVA, and IVB) formed as part of a crystal-liquid fractionation sequence of genetically identical materials (e.g., Pernicka and Wasson 1987), irons of a given group should be expected to be characterized by identical anomalies. A third question relates to whether or not a portion of a meteorite characterized by nucleosynthetic anomalies is representative of the bulk meteorite. For example, lithological heterogeneity among pieces of a single meteorite can lead to varying conclusions regarding the nature and magnitude of isotope anomalies present in the meteorite. This is because some chondritic components, such as the irregularly-distributed calcium-aluminum rich inclusions (CAIs) can contain both high abundances of an element such as Mo, as well diverse isotopic compositions (e.g., Yin et al. 2002). Thus, the analysis of different portions of a chondrite characterized by differing proportions of CAIs, such as Allende, may yield inconsistent results, especially when processing the typically ≤1 g aliquots used for such studies. Nevertheless, in the case of Allende, most CAIs present in this meteorite are characterized by Mo isotope

Nucleosynthetic Isotope Variations in the Solar-System

131

anomalies with r-process excesses (see later section for CAI), making it difficult to produce a W-shaped pattern of s-process deficits without a kink in ε94Mo. More recently, a comprehensive analysis of Mo isotope compositions for a wide range of meteorite samples was made using a newer generation MC-ICP-MS with analytical precisions several times better than early studies (Burkhardt et al. 2011). To overcome the problem of incomplete digestion of refractory presolar grains, the authors conducted complete melting of carbonaceous chondrites using CO2 laser fusion in ultra-pure graphite capsules under oxidizing (atmospheric) or reducing (7% H2–93% Ar) conditions prior to sample digestion with acids (Pack et al. 2010). They discovered that the Mo isotope composition of a CO2-fused Murchison sample was indistinguishable from those of acid-digested Murchison, suggesting that incomplete digestion of presolar grains in chondrites via conventional acid digestion with HF–HNO3–HClO4 would not have a measurable effect on the Mo isotope composition. With the CO2 fusion method, Burkhardt et al. (2011) analyzed Mo isotope compositions for bulk aliquots of five carbonaceous chondrites (CI, CM2, CR2, CO3, CV3) and one enstatite chondrite (EH4). In addition, bulk samples of two ordinary (H3, H6) and two enstatite chondrites (EH4, EL6), as well as metal fractions from two carbonaceous (CR-an, CB) and two ordinary chondrites (L6, LL6) were examined using conventional acid digestion. As shown in Figure 13, all chondrites display positive εMo values with W-shaped patterns indicative of s-process Mo deficits. This observation is even more pronounced when the data are plotted in εiMo–ε92Mo space (Fig. 14), where mixing lines between terrestrial Mo and either a pure s-process or a pure r-process component are clearly discriminated from each other. Excluding one CM chondrite, the chondrite data follow the theoretical s-process mixing line, of which one end-member component is defined by data for SiC grains measured by RIMS (Nicolussi et al. 1998a). The magnitude of Mo isotope anomalies is in the order ofindicating no meaningful correlation between Mo isotope anomaly and lithological and chemical properties. More importantly, enstatite chondrites have Mo isotope compositions that are marginally resolvable from the terrestrial composition. This is comparable to recent high-precision observations for a variety of other elements in enstatite chondrites that have stable isotope compositions similar or identical to those of Figure 13. Molybdenum isotope patterns for (a) terrestrial (e.g., O, N, Cr, Ru, Os: Javoy et al. bulk 13 chondrites, (b) iron meteorites, and (c) palFig. 2010; Herwartz et al. 2014). lasites, angrite and Martian meteorite. Data are from Burkhardt et al. (2011). Error bars are 2 SE of multiple data obtained by different isotopic measurements excluding angrite and Martian meteorite for which 2 SD (standard deviation) from repeated analyses of a terrestrial standard is applied.

The authors also analyzed Mo isotope compositions in differentiated meteorites, including iron meteorites from various groups (IAB-IIICD, IC, IID, IIE, IIIAB, IIIE, IIIF,

6

(c)

Chondrites, CAI

5

ε94Mo

4 3

ce r-ex

2 1

ss

3

1

2

Type B CAI

2

s-de

ficit

0

3

4

ε92Mo

5

6

7

-1

(d)

Irons, Pallasites, Mars, Angrite

2

3

ss

-1

s-d

it

c ef i

ess

c r-ex

1

3

0

0

1

2

3

4

ε92Mo

Irons, Pallasites, Mars, Angrite

2

ε95Mo

(b)

0

e xc r-e

1

-1 -1

Chondrites, CAI

4

Type B CAI

0

6 5

cit ef i s-d ε95Mo

(a)

ε94Mo

Fig. 14

Yokoyama & Walker

132

e xc r-e

5

6

7

ss

s-de

1

ficit

0

-1

-1 -1

0

1

2

ε92Mo

3

4

-1

0

1

2

3

4

ε92Mo

Figure 14. Plots of Mo isotope anomalies in εiMo–ε92Mo space for meteorites and CAI. Data are from Burkhardt et al. (2011). Symbols and errors are the same as Figure 13.

IVAB, ungrouped), two pallasites, two Martian meteorites, and one angrite. Of these samples, all magmatic irons (IIAB, IID, IIIAB, IVA, IVB), IIE irons, and pallasites exhibit positive Mo isotope anomalies with W-shaped patterns, without an ε94Mo kink, whereas the rest do not have any resolvable Mo isotope anomalies. These results are in general agreement with the result of Dauphas et al. (2002a). No meteorites studied have Mo isotope anomalies with negative εMo values. This implies that the Earth, Mars, and that the parent bodies of angrite and non-magmatic IAB-IIICD irons, accreted from precursor material with the most s-process enriched Mo of any cosmochemical materials yet studied, presumably in the inner region of the early Solar System, where the building blocks of meteorite parent bodies formed. The most likely cause of the discrepancies between the early studies and Burkhardt et al. (2011) is simply the issue of analytical precision. Becker and Walker (2003a) and Yin et al. (2002) did not detect the smaller anomalies because of large analytical uncertainties. Ruthenium. Ruthenium is a highly siderophile element with a CI chondrite abundance of 0.692 ± 0.044 µg/g (Lodders 2003). It has seven stable isotopes 96Ru, 98Ru, 99Ru, 100Ru, 101Ru, 102 Ru, and 104Ru, with averaged terrestrial abundances of 5.54%, 1.87%, 12.8%, 12.6%, 17.1%, 31.6%, and 18.6%, respectively (Böhlke et al. 2005). Ruthenium isotopes are synthesized via stellar nucleosyntheses of the p-process (96Ru, 98Ru), the s-process (100Ru), and the r-process (104Ru). Three isotopes are synthesized both by the s- and r-processes (99Ru, 101Ru, 102Ru). In addition, the short-lived 98Tc and 99Tc decay to 98Ru and 99Ru; 98Tc is synthesized by the p-process, while 99Tc lies along the s-process path (Fig. 4). The half-life of 98Tc is poorly constrained, with estimates ranging from 4.2 to 10 Myr (Kobayashi et al. 1993; Parrington et al. 1996). By contrast, the half-life of 99Tc is only 213 kyr (Parrington et al. 1996). High precision Ru isotope analyses have been conducted by both N-TIMS and MC-ICP-MS. Precise Ru isotopic data with MC-ICP-MS for five IIAB and two IIIAB iron meteorites, as well as one carbonaceous chondrite (CV3) and one ordinary chondrite (H5) indicated that the

Nucleosynthetic Isotope Variations in the Solar-System (a)

98

5

Chondrites, CAIs

4

CM2 CV3 LL5 H4

εiRu

3

Coarse‐grained CAI Fine‐grained CAI

2 1 0 ‐1 ‐2 95

(b)

4

133

96

97

98

99

Irons and Pallasites

100

iRu

101

102

IAB IIAB IIIAB IVA

103

104

105

IVB ungrouped Pallasite

Ru/101Ru and 99Ru/101Ru ratios in these meteorites overlapped with the terrestrial values within the ± 0.8 and ± 0.3ε levels of precision reported (2sM), respectively (Becker and Walker 2003b). This implies that these meteorites present no sign of early Tc/Ru fractionation as detectable in Ru isotope shifts, nor do they possess nucleosynthetic Ru isotope anomalies above the stated levels of uncertainty.

Chen et al. (2010) subsequently conducted a 2 comprehensive study of Ru isotopes in two carbonaceous 1 chondrites (CM2, CV3), 0 four ordinary chondrites (LL5, three H4), thirteen ‐1 iron meteorites from varying groups (IAB, IIAB, IIIAB, ‐2 IVA, IVB, ungrouped), 95 96 97 98 99 100 101 102 103 104 105 iRu and three pallasites. This study utilized N-TIMS and Figure 15. Ruthenium isotope compositions for chondrites, reported external precision CAIs, and iron meteorites in εRu units. Data are from Chen et al. for 98Ru/101Ru and 100Ru/101Ru (2010). Error bars are 2SE of multiple data obtained by differFig. 15 at the ± 2.0 and ± 0.3ε levels, ent isotopic measurements excluding fine-grained CAI for which respectively (2s). The authors internal precision of individual isotopic run is applied. found no Ru isotopic effects in the ordinary chondrites and IAB iron meteorites, whereas significant deficits were observed in 100Ru/101Ru ratios for a bulk aliquot of Allende (CV3) (up to – 1.7ε), and the rest of the iron meteorites and pallasites (up to – 1.1ε) (Fig. 15). Incomplete sample digestion would not account for the observed anomalies because all the isotopically anomalous samples contain minuscule or no presolar grains. Rather, the results suggest a widespread and large-scale Ru isotopic heterogeneity in the early Solar System, resulting in a deficit in s-process nuclides or enhancements in both p- and r-process nuclides in the early solar nebula. Chen et al. (2010) argued that the inconsistency between their study and Becker and Walker (2003b) arose from different approaches for correcting massdependent fractionation; Chen et al. used 99Ru/101Ru for normalization whereas Becker and Walker used 96Ru/101Ru. They also presumed that the precision of the other Ru isotope ratios was inadequate to detect the expected isotope shifts in 96,98,102,104Ru/101Ru ratios, owing to an s-process deficit, which were smaller than the change on 100Ru/101Ru. This problem was resolved by recent high precision Ru isotope analyses with MC-ICP-MS for a series of chondrites and iron meteorites (Fischer-Gödde et al. 2012, 2013). When Ru isotope ratios were normalized with 99Ru/101Ru for mass fractionation, the data displayed negative ε100Ru values which clearly correlated with ε102Ru values, consistent with s-process deficit (Fig. 16) (Fischer-Gödde et al. 2012, 2013). εiRu

3

Yokoyama & Walker

134 0.5

CB

ε102Ru

0.0

Chinga

-0.5 -1.0 -1.5 -2.0 -4

IIIAB IVB

CV

IVA OC

CR-an

CM

s-d

-3

it efic

-2 ε100Ru

-1

0

Figure 16. Plot of Ru isotope anomalies for chondrites and iron meteorites in ε102Mo–ε100Ru space. Data are from Fischer-Gödde et al. (2013).

Tungsten. Tungsten is a moderately siderophile element with a CI chondrite abundance of 0.089 ± 0.007 µg/g (Lodders 2003). It has five stable isotopes 180W, 182W, 183W, 184W, and 186W with averaged terrestrial abundances of 0.12%, 26.5%, 14.3%, 30.6%, and 28.4%, respectively (Böhlke et al. 2005). One minor W isotope (180W) is synthesized via the stellar nucleosynthesis of the p-process, Fig. and 16 the rests are synthesized by the s- and r-processes (182W, 183W, 184W, 186 W). In addition, the short-lived 182Hf, which is a predominantly r-process isotope, decays to 182W with a half-life of 8.9 Myr (Vockenhuber et al. 2004) making the 182Hf–182W system useful for chronological applications regarding processes associated with the formation and earliest evolution of planetary bodies (e.g., Kleine et al. 2009). Compared to numerous studies of 182Hf–182W systematics, studies focused primarily on nucleosynthetic W isotope anomalies in meteorites are limited. There are at least two reasons for the scarcity. First, excluding very minor (180W) and radiogenic (182W) isotopes, there are only three W isotopes usable for the study of nucleosynthetic isotope anomalies. Second, exposure to galactic cosmic radiation (GCR) can modify the W isotope compositions of extraterrestrial materials, especially for iron meteorites. Third, most W isotopes were synthesized by two processes (s and r) in nearly equal proportions, making it unlikely for nature to generate large isotopic variations, as observed in Mo isotopes. Consequently, many studies have reported 183W/184W (or 184W/183W) as simply an additional quality check for 182W data. Kleine et al. (2004) measured the 183W/184W ratios in a variety of chondrites as well as eucrites and Martian meteorites, and found no detectable isotope anomalies. Irisawa et al. (2009) measured W isotope ratios for a suite of carbonaceous and ordinary chondrites and found a small deficit in 184W/183W at a level of ~0.4ε for the Allende meteorite (CV3), but not for the other chondrites. These authors argued that the observed anomalies could not be due to the incomplete digestion of presolar SiC grains because the amount of surviving presolar SiC in Allende is significantly lower than for most other carbonaceous chondrites. Rather, they attributed the anomaly to differing proportions of isotopically anomalous CAIs in the samples studied. Further, Qin et al. (2008) reported deficits of ~0.1ε in 184W/183W ratios for IVB magmatic iron meteorites relative to the terrestrial standard and IIAB iron meteorites. These authors concluded that the observed stable W isotope heterogeneity reflected incomplete mixing of the products of s- and r-process components in the solar nebula. Osmium. Osmium is a highly siderophile element with a mean CI chondrite abundance of 0.486 µg/g (Lodders 2003). The search for nucleosynthetic isotope anomalies of Os in meteorites has been achieved almost exclusively using the N-TIMS technique. Osmium has seven stable

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isotopes 184Os, 186Os, 187Os, 188Os, 189Os, 190Os, and 192Os, which are synthesized via stellar nucleosyntheses of the p-process (184Os), s-process (187Os), p- and s-processes (186Os), and sand r-processes (188Os, 189Os, 190Os and 192Os). In addition to these processes, 186Os and 187Os are produced by radioactive decay of 190Pt (T1/2 = 488 Gyr) and 187Re (T1/2 = 41.5 Gyr). Platinum-190 is a pure p-process isotope, while 187Re is predominantly produced by the r-process. Because of the radiogenic isotopes, the terrestrial Os isotope composition cannot be uniquely determined. One of the more widely distributed Os isotope reference materials (UMd Os) has isotopic abundances of 184Os = 0.017%, 186Os = 1.59%, 187Os = 1.51%, 188Os = 13.30%, 189Os = 16.22%, 190 Os = 26.38%, and 192Os = 40.99%, although this particular standard has slightly sub-chondritic 187 Os/188Os and elevated 186Os/188Os (Yokoyama et al. 2010). Of the seven isotopes, two (189Os, 192Os) have large (> 90%) contributions from the r-process. In order to resolve possible nucleosynthetic isotope anomalies, 189Os has been used as the denominator of isotopic pairs, rather than the commonly used 188Os. Further, isotopic fractionation during mass spectrometry is best corrected using a fixed 192 Os/189Os (= 2.527411: Yokoyama et al. 2007). The radiogenic isotope 187Os is not utilized in the study of nucleosynthetic isotope anomalies because of the difficulty in determining the contribution from 187Re decay over 4.567 Ga in meteorites with sufficient accuracy. However, 186 Os can be regarded as a stable isotope after the typically minor correction for 190Pt decay using the Pt/Os ratio in most cosmochemical samples. This is possible because of the long halflife and minor abundance (0.014%) of 190Pt relative to 187Re. Unlike Mo and Ru isotopes, the contribution of p-process nuclides to the total Os is minuscule (~ 0.02%), and Os is composed almost entirely of s- and r-process nuclides. Therefore, Os isotope anomalies with s-process deficits almost completely match the isotopic pattern given by r-process excesses. The first high precision stable Os isotope analysis for bulk meteorites was reported by Brandon et al. (2005). They determined Os isotope compositions in bulk aliquots of three carbonaceous chondrites (C2-ung, CV3, CO3), seven ordinary chondrites (H3.4, two H4, two H5, LL3.4, LL3.6), and three enstatite chondrites (EH4 and two EL6) by applying the Carius tube digestion using a mixture of HNO3 and HCl (Shirey and Walker 1995). One problem evaluating the magnitude of Os isotope anomalies in meteorites is the selection of a standard reference material, because terrestrial samples are known to have variable 186Os/189Os, due to the long-term evolution of 186Os with supra- or sub-chondritic Pt/Os ratios. Thus, reference standards made from these materials do not necessarily represent the bulk Earth composition. To resolve this issue, Brandon et al. (2005) used the mean Os isotope composition of five H-group ordinary chondrites measured in their study as an averaged Solar System component, which exhibited uniform Os isotope ratios identical to the terrestrial values, but have an 186 Os/189Os different from that of the standard reference material they used. Figure 17 shows Os isotope compositions in εOs units normalized to the mean of H chondrites. The ε186Osi indicates the initial ε186Os value corrected for 190Pt decay over 4.56 Ga using the measured Pt/Os ratio. As presented in this Figure, most of the ordinary and enstatite chondrites show no Os isotope anomalies, except for slightly negative ε186Osi values of Parnellee (LL 3.6) and Indarch (EH4). By contrast, significant, large negative Os isotope anomalies are evident for ε186Osi, ε188Os, and ε190Os values in the carbonaceous chondrite Tagish Lake (C2-ung). The εOs values for the rest of carbonaceous chondrites were not resolvable from the bulk Solar System, excluding a small negative ε186Osi value for Ornans (CO3). The magnitude of negative Os isotope anomalies in Tagish Lake is in the order of ε186Osi > ε188Os > ε190Os, suggesting that the anomalies are deficits of s-process origin. An important observation is that replicate measurements of four different Tagish Lake dissolutions resulted in varying εOs values, as clearly shown in ε186Osi. Brandon et al. (2005) interpreted the observed isotope anomalies to be a result of incomplete digestion of isotopically anomalous acid resistant carriers in chondrites, most likely s-process enriched presolar SiC grains. Such grains are abundant in Tagish Lake (~5 to 8 ppm), and are also present in less metamorphosed

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carbonaceous and ordinary chondrites (petrologic grade less than 3.8), as well as in enstatite chondrites of type 3-4 (Russell et al. 1997; Huss et al. 2003). The Tagish Lake data suggest that sample digestion with HNO3 and HCl in Carius tubes dissolve varying proportions of SiC, resulting in nonreproducible εOs values. As mentioned above, conventional acid digestion with HF-HNO3-HClO4 would not have a measurable effect on the Mo isotope composition of SiC bearing chondrites. This is clearly not the case of Os with Carius tube digestion. Therefore, the presence or absence of planetary scale Os isotope heterogeneity remained unclear from this study, and indicated a need for Os isotopic analyses utilizing total sample digestion.

The problem of incomplete sample digestion was resolved by Yokoyama et al. (2007, 2010) who applied an (c) 1 Enstatite alkaline fusion technique that can completely dissolve all chondrite constituents, including presolar SiC. 0 In two studies, the authors analyzed Os isotope compositions in eight carbonaceous chondrites (CI1, C2‐1 ung, two CR2, CM2, two CV3, CK4), EH chondrites EL chondrites three ordinary chondrites (H3.8, two H5), and five enstatite chondrites ‐2 185 186 187 188 189 190 191 192 193 (EH4, four EL6). Although most of iOs the carbonaceous chondrites analyzed contain abundant presolar SiC grains, Fig. 17Figure 17. Osmium isotope compositions in (a) carbona- they obtained uniform εOs values, ceous, (b) ordinary, and (c) enstatite chondrites. Data are which were not resolvable from the from Brandon et al. (2005). Error bars are 2 SE of individ- terrestrial values, excluding ε186Osi ual measurements. across all chondrite classes (Fig. 18). In addition, negative εOs values were observed when SiC bearing carbonaceous chondrites (C2-ung, CM2, CR2) were digested by the Carius tube method (Yokoyama et al. 2007). This implies that Os isotope anomalies in bulk chondrites observed in Brandon et al. (2005) were caused by incomplete sample digestion, as postulated by that study, and that all chondrites and the silicate portion of the Earth most likely possess a common Os isotope composition. ‐2

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The homogenous Os isotope distribution was further corroborated by two subsequent studies. Van Acken et al. (2011) conducted precise Os isotope analyses for a variety of enstatite and Rumuruti chondrites, as well as one differentiated meteorite (aubrite). Because these samples were decomposed by the Carius tube method, some low petrologic grade enstatite and Rumuruti chondrites exhibited small but resolvable Os isotope anomalies due to incomplete sample dissolution, whereas the rest of the samples, including high grade chondrites, show values indistinguishable from the Solar System average.

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Figure 18. μ188Os values for CAIs, chondrites, iron meteorites, and pallasites. Error bars are 2 SE for samples with multiple analyses or the internal precision of individual isotopic run for samples with single measurement. Data sources: Yokoyama et al. (2007, 2010, 2011) and Walker (2012).

Extending the search for nucleosynthetic Os anomalies to other planetary materials, Walker analyzed Os isotope compositions in iron meteorites from groups IAB, IIAB, IIIAB, IVA, and IVB, as well as the main group pallasites. A large proportion of these meteorites exhibit Os isotope anomalies in ε186Osi, ε189Os, and ε190Os values (note that the author used 188 Os as the denominator of isotope ratios). However, the observed anomalies differ from nucleosynthetic Os isotope anomalies in chondrite constituents (see below), implying that the Os isotope anomalies in iron meteorites most likely reflect long-term exposure of the meteorites to cosmic rays, rather than nucleosynthetic effects. In fact, the ε190Os versus ε186Osi and ε189Os trends defined by the iron meteorites and pallasites have end-member compositions defined by samples with limited or no cosmic ray exposure, of which the εOs values are normal and unresolved from the chondritic composition.

Fig. 18 (2012)

Overall, the uniform Os isotope composition in bulk chondrites, as well as in differentiated meteorites, strongly suggests that dust grains with disparate isotopic Os isotopic compositions were thoroughly mixed within the solar nebula at the time of the initiation of planetesimal accretion. This conclusion contrasts with evidence for heterogeneous isotope compositions among early Solar System materials for numerous other elements (e.g., Ti, Cr, Ni, Mo, Ru, Ba, Nd, and Sm: Dauphas et al. 2002a; Andreasen and Sharma 2006; Carlson et al. 2007; Trinquier et al. 2007; Regelous et al. 2008; Trinquier et al. 2009; Chen et al. 2010; Qin et al. 2010; Burkhardt et al. 2011). The contrast must reflect physical and chemical processes in the early Solar System that acted differently on individual elements. This issue will be discussed later in detail. Platinum. Platinum is a HSE with a CI chondrite abundance of 1.004 µg/g (Lodders 2003). It has five stable isotopes 192Pt, 194Pt, 195Pt, 196Pt, and 198Pt with averaged terrestrial abundances of 0.78%, 33.0%, 33.8%, 25.2%, and 7.16%, respectively (Böhlke et al. 2005). In addition, one minor radioactive isotope 190Pt (T1/2 = 488 Gyr) exists with a terrestrial abundance of 0.014%. Platinum isotopes are synthesized via stellar nucleosyntheses by the p-process (190Pt), the s-process (192Pt), and the r-process (198Pt). Three isotopes are synthesized by both the s- and r-processes (194Pt, 195Pt, 196Pt). Platinum is a typical r-process element because four major isotopes (194Pt–198Pt) are nearly all (> 90%) synthesized by the r-process (Arlandini et al. 1999).

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High precision Pt isotope analyses reported for meteorites are rare. The majority of the data are for iron meteorites examined with MC-ICP-MS (Kruijer et al. 2013; Wittig et al. 2013). These studies have utilized Pt isotopes as a tracer for quantitatively evaluating the effect of cosmic ray exposure in order to correct 182W/184W ratios for exposure. These studies discovered that nearly all samples from IID, IVA, and IVB iron meteorites exhibit significant isotope anomalies in 192 Pt/195Pt and 196Pt/195Pt ratios with anomalies as great as + 58.6ε and + 0.88ε, relative to those of a terrestrial standard, respectively. They interpreted, however, that the observed anomalies were not nucleosynthetic in origin but represent the effect of neutron capture on Pt isotopes via cosmic ray exposure. This is because: 1) Pt isotope compositions of different irons from a same group varied randomly as opposed to the observation of Mo and Ru isotopes, and, 2) no correlation was observed in a ε192Pt–ε196Pt diagram. These authors concluded that no nucleosynthetic Pt isotope anomalies were resolvable at the current level of analytical precision.

Isotope anomalies of chalcophile elements in bulk meteorites In the following, we review evidence for nucleosynthetic isotope anomalies in meteorites for three chalcophile elements, Cd, Sn, and Te. Cadmium, Sn and Te are classified as moderately volatile elements with 50% condensation temperatures of 652 K, 704 K, and 709 K, respectively (Lodders 2003). Most of the recent investigations regarding isotope anomalies in meteorites focused on refractory elements, including siderophile elements as reviewed above, whereas isotopic data for moderately volatile elements are limited. The abundances of Cd, Sn, and Te in CI chondrite are the highest relative to the other types of chondrites. Moderately volatile elements, including Cd, Sn, and Te, occur in much lower concentrations in differentiated meteorites, excluding some iron meteorites, with typical levels of 0.1 × CI or less (e.g., Laul et al. 1972). Low abundances in most achondrites make these elements problematic for highprecision isotopic analysis. The depletion of moderately volatile elements in chondrites and achondrites, relative to CI chondrites, has generally been interpreted to reflect incomplete stepwise accretion of nebular material (Humayun and Clayton 1995; Albarède 2009), rather than as a result of loss by outgassing from dust grains/chondrules/planetesimals. Cadmium. The CI chondrite abundance of Cd is 0.675 ± 0.006 µg/g (1s) (Lodders 2003). Cadmium has eight stable isotopes, 106Cd, 108Cd, 110Cd, 111Cd, 112Cd, 113Cd, 114Cd, and 116Cd with averaged terrestrial abundances of 1.25%, 0.89%, 12.5%, 12.8%, 24.1%, 12.2%, 28.7%, and 7.49%, respectively (Böhlke et al. 2005). Cadmium isotopes are synthesized by the stellar nucleosynthesis of the p-process (106Cd, 108Cd), the s-process (110Cd), and the r-process (116Cd). Four Cd isotopes are synthesized both by the s- and r-processes (111Cd, 112Cd, 113Cd, 114Cd). Because Cd is one of the most volatile elements among siderophile and chalcophile elements, Cd isotopes been primarily used for tracking processes associated with vaporization in geochemistry and cosmochemistry (e.g., Wombacher et al. 2003, 2004; Ripperger and Rehkämper 2007). These studies utilized MC-ICP-MS for high precision Cd isotope analysis. Wombacher et al. (2008) determined precise Cd isotope compositions for a comprehensive suite of carbonaceous, ordinary, enstatite, and Rumuruti chondrites, as well as achondrites and lunar samples (soils, breccias, pristine anorthosite). However, none of the samples showed evidence of nucleosynthetic anomalies for the s- and r-process nuclides 111Cd, 112Cd, 113Cd, and 114Cd, nor for the pure s-process isotope 110Cd. Although data for two minor p-process nuclides are not available (106Cd, 108Cd), the result is consistent with another moderately volatile chalcophile element, Te (see below). Tin. The CI chondrite abundance of Sn is 1.68 ± 0.04 µg/g (1s) (Lodders 2003). Tin has ten stable isotopes, 112Sn, 114Sn, 115Sn, 116Sn, 117Sn, 118Sn, 119Sn, 120Sn, 122Sn, and 124Sn with averaged terrestrial abundances of 0.97%, 0.66%, 0.34%, 14.5%, 7.68%, 24.2%, 8.59%, 32.6%, 4.63%, and 5.79%, respectively (Böhlke et al. 2005). This is the largest number of stable isotopes among all elements of the periodic table. Six Sn isotopes are synthesized via

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single stellar nucleosyntheses; the p-process (112Sn, 114Sn, 115Sn), the s-process (116Sn), and the r-process (122Sn, 124Sn). The rest are synthesized both by the s- and r-processes (117Sn, 118Sn, 119 Sn, 120Sn). The large number of isotopes produced by various nucleosynthetic processes make Sn a promising element for the study of nucleosynthetic isotope anomalies in meteorites. However, high precision Sn isotope data for meteorites are not currently available. An early study of Sn isotopes using TIMS discovered no significant isotope anomalies in four iron meteorites or in the metal phase of one mesosiderite (De Laeter and Jeffery 1965). Recently, high precision Sn isotope analyses have been made mainly by MC-ICP-MS (e.g., Lee and Halliday 1995; Moynier et al. 2009; Haustein et al. 2010; Yamazaki et al. 2013). However, sample digestion remains problematic for the analysis of Sn isotopes in meteorites. First, Sn is readily oxidized and precipitates in concentrated nitric acid as meta stannic acid. Second, when Sn is treated with hydrochloric acid, stannic chloride (SnCl4) forms which vaporizes at 114° C (Yamazaki et al. 2013). One possible option for complete sample digestion and subsequent recovery of Sn is the alkali fusion technique, however, application of this method will require the development of new techniques for isolating Sn from the resulting sample solution containing significant amounts of alkali elements. Tellurium. The CI chondrite abundance of Te is 2.33 ± 0.18 µg/g (1s) (Lodders 2003). Tellurium has eight stable isotopes, 120Te, 122Te, 123Te, 124Te, 125Te, 126Te, 128Te, and 130Te with averaged terrestrial abundances of 0.096%, 2.60%, 0.91%, 4.82%, 7.14%, 19.0%, 31.7%, and 33.8%, respectively (Böhlke et al. 2005). Tellurium isotopes are produced by stellar nucleosynthesis of the p-process (120Te), the s-process (122Te, 123Te, 124Te, 125Te, 126Te), and the r-process (125Te, 126Te, 128Te, 130Te). Six out of eight Te isotopes were produced by single nucleosynthetic processes (p-process: 120Te, s-process: 122Te, 123Te, 124Te, r-process: 128Te, 130 Te). Similar to Sn isotopes, the mix of nucleosynthetic production processes makes Te highly advantageous for studying the nucleosynthetic contributions to individual meteorites. The isotopic composition of Te is also affected by the decay of the short-lived nuclide 126Sn to 126Te, with a half-life of 234.5 kyr. Tin126 is generated mainly by the r-process. If 126 Te excesses correlating with Sn/Te ratios are observed in meteorites, it will provide a tight constraint on early Sn-Te fractionation in the protoplanetary disk. As is the case for Cd and Sn, high-precision Te isotope analysis became available with the development of MC-ICP-MS (Lee and Halliday 1995; Fehr et al. 2004). Compared to conventional TIMS techniques, Fehr et al. (2004) improved the precision of 122–130Te/128Te by about one to two orders of magnitude. A comparable precision was obtained by recent N-TIMS technique (Fukami and Yokoyama 2014).

Figure 19. μ126Te values for CAIs, chondrites, and iron meteorites. Error bars are 2 SE for samples with Fig. 19 multiple analyses or the external reproducibility (2 SD) of Te standard measurements for samples with single isotopic run. Note that the error bar for CAI is significant (~500 ppm) and omitted in this.

High precision Te isotope data for meteorites have been obtained using MCICP-MS (Fehr et al. 2004; Fehr et al. 2005). In these two studies, the authors examined bulk aliquots of seven carbonaceous chondrites (CI, CM1/2, CM2, two CV3, CO3.5, CK4), three ordinary chondrites (H5, L3.7, LL6), and one enstatite

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chondrite (EH4), as well as metal and sulfide phases present in iron meteorites (IAB, IC, IIAB). No isotope anomalies were found for any Te isotopes in any of these samples, even for CI and CM chondrites, which contain acid resistant presolar grains (Fig. 19). Although the authors digested chondrite samples using a high-pressure asher (HPA-S Anton Paar) with a mixture of HCl and HNO3, followed by normal acid digestion with HF, this method was likely not as effective as alkali fusion digestion with respect to dissolving presolar SiC. Thus, the Te data contrast with the Os isotopic data for which acid-digested bulk chondrites exhibited apparent Os isotope anomalies, whereas the same samples decomposed via alkali fusion presented no Os isotope anomalies (Brandon et al. 2005; Yokoyama et al. 2007, 2010). Therefore, the homogeneous Te isotope composition in chondrites and iron meteorites most likely indicate that Te isotopes were homogeneously distributed in the solar nebula, and the effect of incomplete dissolution of acid resistant presolar grains is negligible for the resulting Te isotope compositions. As discussed in the next section, this conclusion is reinforced by the observation that sequential acid leachates of carbonaceous chondrites also have shown no detectable Te isotope anomalies (Fehr et al. 2006).

INTERNAL ISOTOPE ANOMALIES PRESENT IN CHONDRITES As seen in the previous section, the extent of isotope anomalies for siderophile and chalcophile elements in bulk aliquots of meteorites, if present, are only marginally higher than the level of analytical precision (within a few ε) for individual isotope ratios. One important observation is that acid-digested bulk chondrites exhibited resolvable Os isotope anomalies, whereas the same samples processed using an alkali fusion total digestion method displayed no detectable Os isotope anomalies (Yokoyama et al. 2010). This strongly suggests the existence of siderophile element-rich, acid resistant, isotopically anomalous phase(s) within the chondrites. Thus, as shown previously via analysis of individual presolar grains, selective digestion of chondrite phases provides strong evidence that chondrites are composed of multiple phases with intrinsic isotope compositions which are dramatically different from those of bulk rocks. In addition to direct isotopic measurement of presolar grains using SIMS/RIMS, two approaches have been recognized as useful methods for detecting internal nucleosynthetic isotope anomalies in chondrites using TIMS and MC-ICP-MS; 1) physically separating CAIs from carbonaceous chondrites and analyzing the isotopic compositions of target elements after chemical purification, and 2) performing sequential acid dissolution of chondrites and subsequent examination of the resulting leachates and acid residues. The extent of isotope anomalies observed in CAIs and acid leachates/residues for many elements are projected to be much larger than observed for bulk samples. These materials, therefore can be highly useful for examining the characteristics of multiple nucleosynthetic components hidden in meteorites. Here we review the progress in the measurement of isotope anomalies in these types of materials for siderophile and chalcophile elements using TIMS and MC-ICP-MS. We also introduce the application of the correlated isotope anomalies for better constraining the composition of s-process nucleosynthesis.

CAIs CAIs are millimeter- to centimeter-sized lithic, refractory inclusions. CAIs are found in most types of carbonaceous chondrites, especially the CV and CO groups, in which the abundances of refractory inclusions can reach ~10% (Scott and Krot 2003). CAIs are the earliest known condensates formed from a hot nebular gas. Based on the mineral chemistry, CAIs were subdivided into two groups Type A and B in a classic study by Grossman (1975). Type A CAIs contain 80–85% melilite, 15–20% spinel, 1–2% perovskite, and minor minerals including plagioclase and hibonite. Type B CAIs contain 35–60% pyroxene, 15–30% spinel, 5–25% plagioclase, and 5–20% melilite. The timing of the oldest CAI formation has been

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dated to be 4567.2–4568.2 Ma (Amelin et al. 2010; Bouvier and Wadhwa 2010; Connelly et al. 2012). Since the 1970s, CAIs have been known to possess isotope compositions different from their host chondrites for a variety of refractory elements including Ca, Sr, Ba, Nd, and Sm, especially for the subset of rare CAIs known as Fractionation and Unknown Nuclear effect (FUN) inclusions (e.g., Lee et al. 1978; McCulloch and Wasserburg 1978a,b; Papanastassiou and Wasserburg 1978). For example, very low 84Sr/86Sr ratios (–0.4‰ relative to a terrestrial standard) are found in some FUN CAIs (Papanastassiou and Wasserburg 1978).

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More recently, high precision isotope analyses for some heavy elements revealed that nucleosynthetic isotope anomalies are ubiquitous in normal CAIs as well (e.g., Brennecka et al. 2013). A prominent example of the high precision isotope analysis of CAIs for siderophile and chalcophile elements is Mo. Yin et al. (2002) analyzed one type B CAI from the Allende meteorite and found an r-process-enriched pattern with ε95Mo ≈ ε100Mo = ~+2 (Fig. 20). As shown in Figure 20, very similar patterns were also obtained for another type B, and one type A CAIs from Allende (Becker and Walker 2003a). These patterns are, however, slightly different from that for presolar SiC X-grains from supernovae (Fig. 8b), which show relatively small anomalies in ε100Mo, leading these authors to conclude that carbonaceous chondrites contain multiple nucleosynthetic components inherited from diverse supernova sources (Fig. 20). Burkhardt et al. (2011) subsequently conducted Mo isotope analyses of five type B CAIs from Allende with higher precision, and obtained patterns with clear kinks at ε94Mo, indicative of the enrichment of an r-process component (Fig. 20). However, the CAI data deviate from mixing lines between the terrestrial component and a putative r-process component (Arlandini et al. 1999) in εiMo–ε92Mo space (Fig. 14). In addition, one peculiar CAI, most likely a fluffy type A inclusion, exhibited much larger positive Mo isotope anomalies (ε92Mo ~+22; Fig. 20), without a 25 (a) ε94Mo kink, indicating an s-process Type A CAI [1] Type A CAI [3] deficit rather than an excess r-process 20 component. The authors concluded 15 that the relatively homogeneous Mo isotope anomalies in five type B CAIs 10 most likely arose from a late injection 5 of freshly synthesized material to the nebula that was enriched in neutron0 rich isotopes, as is consistent with ‐5 other isotope systems, such as Ti, 91 92 93 94 95 96 97 98 99 100 101 iMo Cr, Ni, Zr, Ba, and Nd (McCulloch (b) 5 and Wasserburg 1978a; Niemeyer Type B CAI [1] Type B CAI [2] and Lugmair 1981; Birck and 4 Type B CAI [3] Allègre 1984; Birck and Lugmair 3 1988; Schönbächler et al. 2003). 2 The injected material must have had 1 a Mo isotopic composition either slightly different from that estimated 0 by the s-process model of Arlandini ‐1 et al. (1999), or enriched both in ‐2 r- and p-process components. By 91 92 93 94 95 96 97 98 99 100 101 iMo contrast, the exceptional signature of an s-process deficit in one CAI may Figure 20. Molybdenum isotope compositions in (a) Type A CAIs and (b) Type B CAIs. Data Sources: [1] be the result of condensation from a Becker et al. (2003a), [2] Yin et al. (2002), [3] Burkhardt nebular material which preserved an Fig. et al.20(2011). Error bars are the internal precision of indiearly isotopic heterogeneity. vidual isotopic run for samples with single measurement.

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Chen et al. (2010) determined Ru isotope compositions in three coarse-grained and one fine-grained type B CAIs from Allende. Although the fine-grained CAI did not show Ru isotope ratios resolvable from the terrestrial, three coarse-grained CAIs exhibited negative anomalies exclusively for ε100Ru (–1.6 ε on average; Fig. 15). The anomalies are due either to an s-process deficit or enrichments of r- and p-process components, which is generally consistent with Mo isotope compositions in CAIs. Burkhardt et al. (2008) conducted precise W isotope analyses for bulk samples and mineral separates from several type A and type B CAIs from Allende, some of which were also examined for Mo isotopes, as described above. Unlike the Mo results, most of the investigated CAIs have relative proportions of 183W, 184W, and 186W that are indistinguishable from those of bulk chondrites and the terrestrial standard, although a few samples exhibited 184W/183W ratios ~1 to 2.5 ε units lower than the terrestrial value. The anomalous data most likely reflect an overabundance of r-process relative to s-process isotopes, although the extent of W and Mo isotope anomalies in the same samples were not correlated with each other. Fehr et al. (2009) analyzed Te isotope compositions in bulk samples and acid leachates from five Allende CAIs. The authors discovered minor differences in the Te isotope compositions for these samples relative to the terrestrial standard and bulk Allende, indicative of the presence of small deficits in r-process or excess in s-process isotopes of Te. However, such nucleosynthetic anomalies were barely resolvable from the terrestrial values, given the analytical uncertainties, and will require further investigation with higher analytical precision. High precision stable Os isotope data for CAIs are also rare. Yokoyama et al. (2009) examined two coarse-grained Allende CAIs, and found that none of the CAIs had Os isotopic anomalies that were resolvable from the solar values (Fig. 18). Similarly, Archer et al. (2014) analyzed one group I and one group II CAI from Allende and found no resolvable anomalies. Overall, isotope anomalies of siderophile and chalcophile elements in CAIs have been unambiguously documented for only Mo and Ru (and most likely for W to some extent), both of which are characterized by potential r-process enrichments. This is in agreement with the isotope anomalies of Sr and Zr in CAIs, both of which exhibit excess r-process signatures. Brennecka et al. (2013) noted that the magnitude of anomalies present in Sr, Mo, Ba, Nd, and Sm for eleven CAIs from Allende strongly depend on the atomic mass. With the addition of literature data for Zr and Ru isotope compositions in CAIs, the authors argued that isotopes lighter than mass 140 had r-process excesses, whereas isotopes greater than mass 140 had r-process deficits. Such marked change at an A of ~ 140 suggests a difference in the formation process of the r-process nuclides where A < 140 and A > 140. As noted earlier, ccSNe are responsible for the origin of lighter r-process nuclides, whereas another process such as neutron star mergers could be the sources of heavier r-process nuclides. If true, however, a question remains why some elements such as Te, Os do not follow the rule proposed by Brennecka et al. (2013). More importantly, although isotopic variations for Mo and W are observed in CAIs with different lithologies, it is not clear whether isotopic variability exists within a single CAI for any elements. Precise isotopic measurements of multiple elements in individual CAIs with detailed mineralogical descriptions would shed light on the extent of isotope homogeneity/heterogeneity in the early Solar System at the time of CAI formation.

Acid leachates and residues Sequential acid leaching is known to be a useful approach for identifying internal nucleosynthetic isotope anomalies in chondrites. It is a form of chemical mineral separation and can concentrate diverse presolar grains with different acid resistance into individual leaching fractions. In fact, the first discovery of presolar grains was made by the analysis of rare gas isotopes in the acid residues of a primitive chondrite (Lewis et al. 1987), and various types of acid resistant presolar grains have been extracted from the acid digestion residues of individual

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Table 3. An example of sequential acid leaching procedure applied to ~16.5 g of Murchison meteorite in Reisberg et al. (2009). Step Step 1

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Figure 21. Molybdenum isotope compositions in acid

Fig. 21 leachates from (a) Allende, (b) Orgueil, and (c) Murchison. Data sources: Dauphas et al. (2002a, 2002b), Burkhardt et al. (2012b). Error bars are the internal precision of individual isotopic run for samples with single measurement.

chondrites (Huss et al. 2003). As summarized below, high precision isotope analyses of siderophile and chalcophile elements for chondritic acid leachates and residues have been conducted for Mo, Te, W, and Os since 2002. The typical approach of such studies has been to successively acid leach a chondrite sample in multiple steps proceeding from weak leaching conditions (dilute acid, lower temperature, shorter time) to harsh leaching conditions (concentrated acid, higher temperature, longer time), in order to concentrate as many distinct types of presolar phases as possible. Table 3 shows an example of a sixstep acid leaching procedure used to separate Os from a carbonaceous chondrite (Reisberg et al. 2009). In this case, the particular phases targeted by the stepwise leaching included sulfides, carbonates, and fine-grained matrix minerals, including amorphous glass in the first two steps (steps 1–2), metal and tiny olivine grains in step 3, and more acid-resistant phases (e.g., pyroxenes, oxides) in the later steps (steps 4–6). Most of the studies discussed used similar stepwise leaching procedures. Molybdenum. Dauphas et al. (2002b) performed precise Mo isotope analysis for acid leachates obtained by four-step acid leaching of two carbonaceous chondrites, Allende (CV3) and Orgueil (CI1). The authors found that the four leachates of Allende presented W-shaped Mo isotopic

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patterns which are essentially the same as that of bulk Allende, revealing no internal isotopic heterogeneity in this meteorite (Fig. 21a). By contrast, two out of the four leachates from Orgueil indicated the existence of two distinctive phases for Mo isotopes; a W-shaped pattern in step 2 indicative of s-process deficit, and an M-shaped pattern in step 3 which mirrors the pattern in step 2 (Fig. 21b). The authors concluded that one particular nucleosynthetic carrier phase cannot account for the Mo isotopic signature observed in the step 2 leachate, because the coincidental coupling of p- and r-process enrichments in a single presolar phase is unrealistic for generating the observed s-process deficit. The carrier of the M-shaped pattern in step 3 is most likely mainstream presolar SiC grains, although SiC is known to be significantly acid resistant. The authors speculated that the sequential acid leach did not destroy SiC grains but altered the surface morphology, which enabled leaching of MoC inclusions within SiC. However, it should be noted that SiC grains are not the only carriers of pure s-process Mo. As noted earlier, Nicolussi et al. (1998b) found s-process Mo isotope signatures in presolar HD graphite grains, which would be more susceptible to acid dissolution than SiC grains. Although presolar graphite is less abundant than SiC in carbonaceous chondrites, the contribution of s-process-enriched presolar graphite grains in the step 3 leachate is conceivable. Burkhardt et al. (2012b) carried out Mo isotope analyses for acid leachates from Murchison (CM2). The samples analyzed were the aliquots of leachates prepared by Reisberg et al. (2009) (Table 3), although the final residue remaining after step 5 was first fused by CO2 laser under a reducing atmosphere (Burkhardt et al. 2011), then digested in HNO3-HF-HClO4 to completely dissolve acid resistant presolar grains. Similar to the result of Orgueil, the Murchison leachates exhibited the patterns of s-process deficits in the earlier leaching steps (steps 1–3), and those with s-process excesses in the later leaching fractions (steps 4–6) (Fig. 21c). In particular, the largest positive and negative εMo values were obtained from the first leachate (ε92Mo = +30.5) and the final residue (ε92Mo = –79.3), respectively. The εiMo vs. ε92Mo plots clearly indicate that at least two components must be present in this meteorite to account for the observed covariations. The authors concluded that the carrier of an s-process-enriched component, which plotted somewhere on the correlation lines beyond the step 6 data, was presolar SiC, whereas the s-process-depleted component represented by the step 1 leachate data most likely reflected a homogenized nebular component. Tungsten. Burkhardt et al. (2012a) determined W isotope compositions in the same acid leachates of Murchison used for Mo and Os isotope analyses (Burkhardt et al. 2012b), and revealed significant internal W isotope variations within this meteorite that reflect a heterogeneous distribution of s- and r-process components (Fig. 22). As was the case of Mo, the first leachate 10

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(step 1) released elevated 182W/184W and 183W/184W ratios relative terrestrial values (ε182W = +2.81, ε183W = +3.62), indicating a deficit of s-process W isotopes in this fraction. Conversely, the later fraction was characterized by negative ε182W and ε183W values, especially for the final residue (ε182W = –25.48, ε183W = –15.28), which generally match the pattern of mainstream SiC grains (Ávila et al. 2012). As a result, a positive correlation is observed in a plot of ε182W versus ε183W. Burkhardt et al. (2012b) concluded that the Mo and W isotope compositions of the different leach steps of Murchison broadly correlated, as expected from s-process nucleosynthesis, indicating that presolar Mo and W are most likely hosted in the same carriers.

ε190Os

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Osmium. Unlike Mo and W, no Os isotope anomalies have been found in bulk aliquots of chondrites and differentiated meteorites. However, the existence of internal Os isotope heterogeneity has been identified for a variety of chondrites. The six-step leaching experiment of Reisberg et al. (2009) revealed large internal Os isotope anomalies of nucleosynthetic origin (ε184Os from −108 to +460; ε186Os from −14.1 to +12.6; ε188Os from −2.6 to +1.6; ε190Os from −1.7 to +1.1). The Os isotope anomalies are correlated and form nearly linear trends in εiOs versus ε188Os plots (Fig. 23), strongly suggesting the existence of at least two phases with anomalous nucleosynthetic compositions within the meteorite; the s-processenriched and r-process-enriched components, relative to the terrestrial. Much of the s-process rich Os, however, was released by relatively mild leaching (steps 2–4), as opposed to the results for Mo and W (Burkhardt et al. 2012b), which were characterized by the strongest s-process-enrichments in the final residue, most likely due to the incorporation of mainstream presolar SiC. Moreover, the enrichments of r-process isotopes were observed in both steps 1 and 5 leachates. These results indicate the presence of multiple distinct presolar phases, with anomalous nucleosynthetic compositions in the Murchison meteorite. The authors pointed out that the enrichment of s-process Os in early leaching fractions was not 15 caused by the dissolution of s-process 10 L3 L4 rich presolar SiC, but was due either to L2 5 the dissolution of a readily leachable, L6 0 unidentified presolar phase enriched ‐5 in s-process Os, or to the existence of L1 a complementary, chemically resistant ‐10 L5 r-process-rich phase that was not ‐15 dissolved by these leaching steps. In ‐20 addition, the authors also speculated ‐3 ‐2 ‐1 0 1 2 upon the existence of p-process rich ε188Os presolar grains to account for elevated 2.0 ε184Os values observed in some leach 1.5 fractions. 1.0 L4 L3

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Yokoyama et al. (2007, 2010) conducted Os isotope analyses for acid leachates and residues of not only carbonaceous, but also ordinary chondrites. These authors specifically focused on: 1) early leachates of bulk meteorites, 2) the residues enriched in insoluble organic matter (IOM) extracted via demineralization of chondrites, and 3) acid leachates and residues of the IOM-rich fraction. The IOM-rich residues were prepared by

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Figure 24. Osmium isotope anomalies in (a, b) acid residues and (c, d) acid leachates from carbonaceous and ordinary chondrites. Data are from Yokoyama et al. (2010). Error bars are uncertainties of individual measurements (2 SE). Circle, diamond, and square symbols are residues (filled) and leachates (open) from chondrite samples with petrologic types 1, 2, and 3, respectively. Black bold lines represent mixing lines between a solar component and a presumed s-process component, for which the slopes are 4.84 ± 0.34 for (a,c) and 0.647 ± 0.027 for (b,d). Bold and thin dashed lines in (a, b) are mixing trends for the solar and an s-process component calculated by Reisberg et al. (2009) using the Os MACS values in Mosconi et al. (2006) and Bao et al. (2000), respectively. Gray thin lines in (c, d) represent regressions of leachate samples with 2s error envelope (gray dashed curves).

a relatively mild, low temperature leaching of the bulk meteorites with CsF/HF (Cody and Alexander 2005), preserving a larger amount of presolar grains compared to the final residues of sequential acid leaching. As shown in Figure 24, nearly all IOM-rich residues exhibit s-process rich Os isotope signatures due to the incorporation of s-process-enriched carriers such as presolar SiC. The extent of s-process enrichment is generally in the order of the petrologic grade of host chondrites (CI1 > C2-ung > CM2 > CR2 ≈ C2/3 > CV3 ≈ L3.05). The change in the pattern of the Os isotope anomalies across the chondrite groups is evidently controlled by the abundance ratio of an s-process-rich phase vs a normal phase with solar Os isotope composition in the IOM-rich residues. Compared to type 3 chondrites, type 1 and type 2 chondrites well preserve presolar grains because of the lack of destructive thermal metamorphism on their parent bodies. On the other hand, relative to residues from CI chondrites, type 2 residues would be slightly, and type 3 significantly, enriched in high temperature components with a solar Os isotopic composition that survived acid dissolution (e.g., Os-alloy, chromites, and Mg-rich spinels). As opposed to IOM residues, early leachates of bulk chondrites are characterized by negative εOs values that are indicative of the enrichment of r-process Os isotopes, which complement the s-process enrichments in IOM-residues (Fig. 24). By contrast, the leachates of IOM-residues exhibit both positive and negative εOs values, suggesting the existence of multiple presolar phases within IOM-rich residues. Such phases are variably enriched in s-, r-, and possibly p-process Os isotopes, because the trend generated by leachate samples in a ε186Osi vs. ε188Os plot is clearly resolved from that of residue samples.

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This approach was further extended to evaluate the effects of nebular and parent body processes on presolar components in chondrites (Yokoyama et al. 2011). The authors determined Os isotope compositions in IOM-rich residues separated from type 1 and type 2 carbonaceous chondrites that were subjected to varying degree of aqueous alteration on their parent bodies. They also examined residues from enstatite chondrites, which formed under markedly reduced redox conditions, compared to carbonaceous and ordinary chondrites. All the IOM-rich residues analyzed were characterized by positive εOs values, indicative of excess s-process isotopes (Fig. 25). Similar to the previous study, the magnitude of positive εOs values in the IOM-rich residues from carbonaceous chondrites was in the order of petrologic grade of the host meteorites (type 1 > type 2 > type 3). More importantly, this trend existed even within a single chondrite group; i.e., CR1 > CR2 and CM1 > CM2. The authors concluded that the isotopic variation in IOM-rich residues from type 1–2 chondrites was caused by the selective destruction, during aqueous alteration on their parent bodies, of presolar grains enriched in r-process Os, most likely either presolar silicates or other unidentified, reduced presolar phases, such as metal alloys, carbides, and silicides. On the other hand, the IOM-rich residues from type 3–4 enstatite chondrites have larger positive εOs values than those of carbonaceous and ordinary chondrites that experienced the same degree of thermal metamorphism. This argues for the preservation of s-process-rich presolar grains in enstatite chondrites during heating events under reduced conditions, either while components formed in the nebula, or on their parent bodies. Overall, chondrites possess significant internal Os isotope anomalies as documented in their acid leachates and residues. To account for the uniform, terrestrial Os isotope composition in bulk chondrites, any destructive processes of isotopically anomalous presolar components in the

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nebula or on the parent bodies did not lose the released component but re-incorporated into a new phase(s) so as not to modify the bulk isotopic composition. However, this might not be the case for elements which have different volatility and/or chemical characteristics compared to Os. Tellurium. The features of Te isotope compositions in chondritic acid leachates are markedly different from what are observed for Mo, W, and Os. Fehr et al. (2006) analyzed Te isotope compositions with MC-ICP-MS in sequential acid leachates from three carbonaceous chondrites, Orgueil, Murchison, and Allende. The leaching procedure sequence was acetic acid, HNO3, HCl, HF-HCl, and HF-HNO3. Unlike Mo, W, and Os, no sign of isotope anomalies were detected in the abundances of p-, s-, and r-process nuclides for leachates from the three chondrites, excluding a nitric acid fraction of Murchison, which showed an elevated 130Te/126Te ratio (ε130Te = +3.5 ± 2.5) indicative of a small excess of the r-process nuclide. The general absence of isotope anomalies in chondritic acid leachates is inconsistent with the results not only for Mo, W and Os, as noted above, but also for Cr, Zr, and Ba (Rotaru et al. 1992; Hidaka et al. 2003; Schönbächler et al. 2005). The authors concluded that the reason for this mismatch was unclear, but it might reflect volatility and more efficient mixing of Te in the solar nebula. By contrast, Fukami et al. (2013) performed Te isotope analyses of acid leachates and residues with N-TIMS for Allende, Murchison, and Tagish Lake (C2-ung). The leaching method was the same as described by Reisberg et al. (2009) (Table 3), except that that the final residue remaining after step 6 was first combusted at 1000 ºC in a sealed quartz glass tube, then digested with HF-HNO3 in a high pressure digestion system to completely dissolve refractory presolar grains such as SiC. All but one of the leachate samples were characterized by Te isotopic compositions that were indistinguishable from the terrestrial values. An exception was the final residue of Allende which showed a low 130Te/128Te ratio compared to the terrestrial value (ε130Te = –33±17). Presolar SiC does not account for the observed anomaly because the final residues from Murchison and Tagish Lake showed no Te isotope anomalies resolvable from the terrestrial composition, and these meteorites are known to contain more presolar SiC than Allende. Fukami and Yokoyama (2014) concluded that the Te isotope composition of the final residue of Allende could be reproduced either by depletion of the r-process component (Maas et al. 2001), or by incorporation of an anomalous component represented by that in nanodiamond (Richter et al. 1998). The final residue (nanodiamond) fraction of Allende does not exhibit distinctive isotope anomalies for Mo (Dauphas et al. 2002b), Sr (Yokoyama et al. 2015), or Os (Yokoyama et al. 2007), but evidently shows Xe isotope anomalies (Lewis and Anders 1988). Again, the reason for the discrepancy may be related to the volatility of individual elements, although the details remain unclear.

Isotopic constraints on the s-process nucleosynthetic component As described above, precise isotopic analyses of Mo, W, and Os in acid leachates, acid residues, and CAIs from chondrites reveal that chondrites possess internal nucleosynthetic isotope anomalies of varying degrees. A common feature is that the extent of isotope anomalies between different isotope pairs with a common denominator are mutually correlated across different sample fractions, forming straight lines in three isotope plots, which generally represent the binary mixing of a nearly pure s-process carrier phase and a terrestrial component. Unlike SIMS and RIMS techniques, analyses of chondrite constituents do not directly provide the endmember isotopic composition of an isotopically anomalous carrier phase (e.g., SiC). However, the slope of the correlation line in a three isotope plot allows calculating the ratio of two numerator isotopes of the putative s-process carrier phase. This further enables evaluating the physical conditions for the nucleosynthetic site of the s-process, as well as the discussion of unstable s-process branching. Brandon et al. (2005) determined the 186Os/188Os ratio of an s-process rich end-member component residing in chondrites to be ~0.48 using the slope of inter-correlation between ε186Osi

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and ε188Os obtained from the analyses of chondrite samples. This value is substantially lower than the theoretical solar s-process ratio of ~0.68 (Beer et al. 1997). Brandon et al. (2005) proposed that the lower s-process 186Os/188Os ratio was generated by increased branching at 186Re, due to an elevated stellar neutron density of 6 × 108 n/cm3 under the conditions of the main s-process in AGB stars, where a neutron is released by the 13C(α, n)16O reaction at T = ~1 × 108 K. Alternatively, it would be produced by the weak s-process that occurs in massive stars via the reaction of 13 C(α, n)16O at T = ~3 × 108 K, assuming a neutron density of 10 × 108 n/cm3. Such a high neutron density at the s-process nucleosynthetic site was challenged by Yokoyama et al. (2007) who pointed out that the neutron density estimated from the other important s-process paths (e.g., Nd–Pm–Sm, Er–Tm–Yb and Os–Ir–Pt) was inconsistent with that proposed by Brandon et al. (2005). Yokoyama et al. (2010) reevaluated s-process nucleosynthesis for Os using the ε186Osi, ε188Os, and ε190Os values of bulk chondrites and their acid residues, and obtained the correlation slopes of 4.84 ± 0.34 for ε186Osi vs ε188Os vs and 0.647 ± 0.027 for ε190Os vs. ε188Os (Fig. 24). In order to compare these values with those of the hypothesized mixing between the terrestrial and s-process components, they applied the following equation derived by Dauphas et al. (2004) which governs the mixing line in an εiOs–ε188Os diagram plotting chondrite data obtained via N-TIMS analysis with mass fractionation correction using a fixed, solar 192Os/189Os ratio: i ρiOs − ρ192 Os ⋅ µ Os εiOs =188 × ε188 Os 188 ρOs − ρ192 Os ⋅ µ Os

(3)

where  i Os  ρiOs =  189   Os  s − process

 i Os   189  − 1  Os solar

µiOs = ( i − 189 ) (192 − 189 ) By applying these equations, they obtained 6.47 and 0.64 for the mixing lines of ε186Os–ε188Os and ε190Os–ε188Os, respectively, using the parameters of a “stellar model” for s-process nucleosynthesis (Arlandini et al. 1999). This indicates that the 190Os/188Os ratio in the projected s-process component is in good agreement with the theoretical value, whereas the s-process 186Os/188Os ratio obtained from chondrite data is lower than the theoretical value. Yokoyama et al. (2007) concluded that the mismatch could be explained either by another s-process path to produce 187Re through the relatively long-lived isomer of 186 Re (T1/2 = 2.0 × 105 yr) (Hayakawa et al. 2005; Meyer and Wang 2007), or by errors in the input parameters to the nucleosynthetic calculations. For example, the Maxwellian-averaged neutron capture cross sections (MACS; sn) of 186Os, 187Os, and 188Os from Mosconi et al. (2006) yield lower s-process 186Os/188Os ratios than the earlier parameters used in Arlandini (1999) or Bao et al. (2000). Two follow-up studies (Reisberg et al. 2009; Yokoyama et al. 2010) reached the same conclusion that the hypothetical mixing between the terrestrial and s-process components reproduced well the 186Os/188Os ratio observed in chondritic acid leachates and residues when the new MACS values of Mosconi et al. (2006) were used. On the other hand, Humayun and Brandon (2007) proposed that correlated Os isotope anomalies in chondrites can be conversely used for tightly constraining the MACS value of sn(190Os) by combining the sn(188Os) value measured in the laboratory and the s-process 190 Os/188Os ratio projected by the chondrite data. This is useful because the sn(190Os) was not measured in the study of Mosconi et al. (2006), who gave reliable MACS of 186Os, 187Os, and 188 Os that matched the chondrite data as mentioned above. By using the Os isotope data from Brandon et al. (2005) and sn(188Os) = 291±15 mbarn (Mosconi et al. 2006), coupled with the

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s-process local approximation model (i.e., abundance × cross section = constant), the authors determined the MACS value of sn(190Os) = 249 ± 18 mbarn at a thermal energy of 30 keV. This is ~20% lower than the value of sn(190Os) = 295 ± 45 mbarn reported by Bao et al. (2000). Subsequently, Reisberg et al. (2009) followed this approach to determine the MACS value of sn(190Os) using their Os isotope data of acid leachates from the Murchison meteorites. Unlike Humayun and Brandon (2007), the authors did not assume the s-process local approximation but applied a full nuclear reaction network calculation, because 188Os is fed by several branching points in the s-process pathway, in particular 185W (Fig. 4). In addition, they included the effect of internal mass fractionation correction as presented in Equation (1). Consequently, the authors reported the MACS value of sn(190Os) = 200 ± 22 mbarn, which is lower than that reported by Bao et al. (2000) or that estimated by Humayun and Brandon (2007).

ORIGIN OF PLANETARY SCALE ISOTOPE ANOMALIES IN METEORITES As seen in previous sections, a variety of nucleosynthetic isotope anomalies have been documented in bulk meteorites and/or chondrite constituents (CAIs, acid leachates and residues) for siderophile elements such as Mo, Ru, W, and Os, yet isotopic anomalies for chalcophile elements such as Cd and Te are either barely resolved, or non-existent. Anomalies are have also been documented in meteorites for lithophile elements including Ti, Cr, Ni, Sr, Zr, Ba, Nd, Sm, as well as the siderophile element Ni (e.g., Schönbächler et al. 2003; Andreasen and Sharma 2007; Regelous et al. 2008; Trinquier et al. 2009; Qin et al. 2011a; Akram et al. 2013). Despite this wealth of data, however, it remains unclear to what extent nucleosynthetic isotope anomalies exist on the global scale of parent bodies of meteorites, and even planets. A useful approach for elucidating the origin of planetary scale isotope anomalies would be to directly plot and compare the isotopic data of different elements obtained from various meteorite groups. For instance, Warren (2011) found a fundamental dichotomy between carbonaceous chondrites and other meteorites when isotope anomalies for O, Ti, and Cr were plotted in diagrams such as ε54Cr–ε50Ti and ε54Cr–Δ17O. The author concluded that the isotopic dichotomy was due to the difference in formation location between two groups; non-carbonaceous meteorites accreted in the inner Solar System (sunward relative to Jupiter), whereas carbonaceous chondrites originally accreted in the outer Solar System. Two models have been proposed to account for the observed heterogeneous distribution of isotopically normal and anomalous components. Dauphas et al. (2010) and Qin et al. (2011b) discovered that the carrier phase of 54Cr anomaly in meteorites was nanoparticles (< 100 nm) of Cr-rich spinels with enhanced 54Cr/52Cr ratios, which were likely produced by a type II supernova. They argued that the planetary scale 54Cr anomaly was caused by late injection of a nearby supernova which scattered 54Cr rich nano-grains into the protoplanetary disk, followed by aerodynamic sorting of isotopically normal, coarser grains that resulted in variations in the isotopic abundance of 54Cr across planetary materials. This model can potentially explain the correlated isotope anomalies between 50Ti and 54Cr in bulk chondrites assuming that the carriers of these two neutron-rich isotopes were produced at the same stellar site and had similar grain sizes and chemical properties such that they could not be decoupled during incorporation in planetesimals. Alternatively, Trinquier et al. (2009) proposed that the correlation between 50 Ti and 54Cr in bulk chondrites was the result of selective destruction of thermally unstable, presolar grains in the solar nebula. As for siderophile elements, Dauphas et al. (2004) discovered that isotope anomalies of Mo (Dauphas et al. 2002a) and Ru (Chen et al. 2010) were well correlated in iron meteorites when bulk measurements were grouped by meteorite classes, forming a straight line on a plot

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of ε100Ru–ε92Mo, with the line passing through the origin of the plot which represents the terrestrial composition (Fig. 26). The Mo–Ru trend extends toward the direction of s-process Fig. 26 deficit, implying that isotopically anomalous Mo and Ru are hosted in the same s-process-rich carriers, which were distributed in the solar nebula in varying proportions of isotopically normal to anomalous phases. A similar correlation between isotope anomalies of Mo and Ru was also reported for bulk carbonaceous chondrites (Fischer-Gödde et al. 2013), although the slope of the trend in the ε100Ru–ε92Mo diagram deviates slightly from the theoretical s-process and terrestrial mixing line (Arlandini et al. 1999) (Fig. 26). However, the correlations are not conclusive for discriminating the two models mentioned above; late supernova injection coupled with grain size sorting, and nebular thermal processing. Burkhardt et al. (2012b) investigated Mo and W isotope compositions in bulk aliquots of meteorites, as well as in chondrite acid leachates and residues. Unlike previous studies, these isotope compositions were all measured by a single research group using the same sample fractions, eliminating the possibility of isotopic inconsistency caused by sample heterogeneity. As noted in the previous section, Mo and W isotope anomalies in chondritic acid leachates were broadly correlated, suggesting that isotopically anomalous Mo and W are likely hosted in the same carriers in chondrites (Fig. 27). However, as opposed to Mo isotopes, most meteorites do not show W isotope anomalies on the bulk scale. The discrepancy of external and internal isotope anomalies for Mo and W isotope suggests that initially homogeneous mixtures of isotopically normal and anomalous dust grains in the solar nebula was disturbed only for Mo, but not for W. The authors concluded that selective removal of volatile Mo oxides occurred due to destructive thermal processing within the solar nebula, resulting in Mo isotope heterogeneity, while refractory W was not affected. This model merits further consideration for explaining why bulk meteorites also do not possess Os isotope anomalies, whereas acid leachates and residues presented significant isotope anomalies for these elements. Reduced Os is similarly refractory to W, so it can perhaps be inferred that the isotopic compositions of Os in bulk meteorites have not been modified via thermal processing. Yokoyama et al. (2014) speculated that Te isotope compositions were not fractionated via thermal processing because their moderately volatile nature caused total evaporation of Te during heating, and that this did not create isotopic heterogeneity. In other words, to fractionate Mo isotopes via thermal processing, selective volatilization of isotopically

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Figure 27. Correlation of Mo and W isotope anomalies in meteorites. Data are from Burkhardt et al. (2012b).

anomalous Mo associated with physical separation of gas and remaining solid must have occurred. If true, isotope anomalies at the bulk meteorite scale can be observed only for elements with a 50% condensation temperature higher than Te (709 K) and lower than W (1789 K). Molybdenum and Ru match this condition, while Cd is more volatile than Te, so that no Cd isotope anomalies were observed in meteorites. Fig. 27

CONCLUDING REMARKS

In the last decade, an increasing amount of highly precise data for isotope anomalies of siderophile and chalcophile elements in a variety of meteorites have been produced by isotope measurements using the latest mass spectrometers. One of the most remarkable findings of these investigations is that isotope anomalies at the bulk meteorite scale do not prevail for all elements, but are limited to elements such as Mo and Ru. By contrast, isotope anomalies for chondrite constituents (CAIs, acid leachates, and residues) are more widespread for elements including Mo, Ru, W, and Os. In general, there are at least two isotopically anomalous carrier phases dominantly residing in chondrites; one is enriched in s-process isotopes, which most likely are represented by the mainstream presolar SiC grains, while the other phase is a complementary component to the s-process-rich material. Some unidentified phases may also be present in chondrites which are enriched or depleted in p-process nuclides, but these phases are evidently minor. The inter-correlation for the magnitude of isotope anomalies among different elements is useful for identifying the cause of planetary scale isotope anomalies. Correlations for siderophile and chalcophile elements have revealed that thermal processing in the nebula likely caused selective destruction of thermally labile presolar phases and preferential isotope fractionation of moderately refractory elements, such as Mo and Ru, via physical separation of gas and solid. However, other processes such as late injection of supernova associated with grain sorting, or parent body processes including aqueous alteration and thermal metamorphism, must be borne in mind as alternative causes for generating planetary scale isotope anomalies in meteorites.

ACKNOWLEDGMENTS We would like to extend special acknowledgment to the volume editors James M.D. Day and Jason Harvey for providing an opportunity for writing this chapter. We are indebted to Larry R. Nittler and Bradley Meyer for constructive review, which improved the quality of the paper significantly. This research was supported by Grants-in-Aid for Scientific Research from the Japan Society for the Promotion of Science (21740388 and 23340171), and from NASA (NNX13AF83G).

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Reviews in Mineralogy & Geochemistry Vol. 81 pp. 161-238, 2016 Copyright © Mineralogical Society of America

Highly Siderophile Elements in Earth, Mars, the Moon, and Asteroids James M.D. Day Geosciences Research Division Scripps Institution of Oceanography La Jolla, CA 92093-0244 USA [email protected]

Alan D. Brandon Department of Earth and Atmospheric Sciences University of Houston Houston, TX 77204-5007 USA [email protected]

Richard J. Walker Department of Geology University of Maryland College Park, MD 20742 USA [email protected]

INTRODUCTION The highly siderophile elements (HSE: Os, Ir, Ru, Rh, Pt, Pd, Re, Au) are key tracers of planetary accretion and differentiation processes due to their affinity for metal relative to silicate. Under low-pressure conditions the HSE are defined by having metal–silicate partition coefficients in excess of 104 (e.g., Kimura et al. 1974; Jones and Drake 1986; O’Neill et al. 1995; Borisov and Palme 1997; Mann et al. 2012). The HSE are geochemically distinct in that, with the exception of Au, they have elevated melting points relative to iron (1665 K), low vapour pressures, and are resistant to corrosion or oxidation. Under solar nebular conditions, Re, Os, Ir, Ru, Rh, and Pt, along with the moderately siderophile elements (MSE) Mo and W, condense as refractory-metal alloys. Palladium and Au are not as refractory and condense in solid solution with FeNi metal (Palme 2008). Assuming abundances of the HSE in materials that made up the bulk Earth were broadly similar to modern chondrite meteorites, mass balance calculations suggest that >98% of these elements reside in the metallic core (O’Neill and Palme 1998). In practical terms, the resultant low HSE abundance inventories in differentiated silicate crusts and mantles enables the use of these elements in order to effectively track metallic core formation and the subsequent additions of HSE-rich impactors to planets and asteroids (Fig. 1). In detail, the absolute and relative abundances of the HSE in planetary materials are also affected by mantle and crustal processes including melting, metasomatism, fractional crystallization, and crust-mantle remixing, as well as later impact processing, volatility of Re under oxidizing conditions, 1529-6466/16/0081-0004$10.00

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Figure 1. (a) Schematic of a hypothetical differentiated rocky planetary body, broken into a metal core (solid inner = black; liquid outer = gray) and silicate mantle and crust. (b) Baseline distributions of highly siderophile elements (not including monoisotopic Rh and Au) in different core, mantle and crust reservoirs of Earth, the Moon and Mars (adapted from Day 2013). The mantles of Earth (terrestrial mantle [TM] = black stippled line), the Moon (lunar mantle [LM] = black line), and Mars (martian mantle [MM] = gray dashed line) have broadly chondritic relative abundances of the HSE, with the martian and terrestrial mantles having similar absolute abundances (~0.007–0.008 × CI Chondrite), and the Moon’s mantle being > 40 times more depleted in the HSE (~0.0002 × CI Chondrite). These patterns do not match predicted silicate compositions after core formation (Mann et al. 2012), suggesting that late-accretion addition of HSE-rich impactors may ultimately have been required to obtain chondritic relative HSE abundances in planetary mantles. Terrestrial continental crust (CC = thin gray dashed line) shows similar crust–mantle partitioning characteristics to the lunar crust (LC = thin gray line), indicating strong HSE fractionation during crustal growth. Data for equilibrium metal matches the least evolved magmatic iron meteorite HSE abundances (e.g., IVB iron meteorite Warburton Range; Walker et al. 2008). Reservoir data given in the tables and numbers beneath the elements are melting temperatures of pure elemental metals in Kelvin (Emsley 1991).

and low-temperature secondary alteration (cf., Day 2013; Gannoun et al. 2016, this volume). In the absence of metal, the HSE are chalcophile, so these elements are also affected by processes involving growth and breakdown of sulfides. Work over the last several decades has led to a large available database for understanding processes affecting the HSE for planetary bodies. This chapter summarises this progress for rocky Solar System bodies, including the Earth, Moon, Mars and some asteroids, and examines the conceptual framework for interpreting these data. The first section outlines the motivation for measuring the HSE in planetary materials. The second section briefly considers methods for measuring and interpreting HSE abundance and Os isotopic data. The third section provides an outline of natural HSE abundance variations and Os isotope compositions in planetary materials. The fourth section outlines current interpretations of the available data and outstanding issues. The final sections offer some comparative planetology, implications for terrestrial planet formation, synthesis and future directions. This chapter does not consider nucleosynthetic variations, as these are the subject of a review by Yokoyama and Walker (2016, this volume), and does not provide a detailed consideration of experimental work, which is the subject of Brenan et al. (2016, this volume). While comparisons are made with terrestrial HSE compositions, these data are considered in detail elsewhere in this volume, or in Walker et al. (1997), Shirey and Walker (1998), Carlson (2005), Walker (2009), and Day (2013).

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MOTIVATION FOR STUDY AND BEHAVIOR OF THE HSE IN PLANETARY MATERIALS The HSE have a strong affinity for Fe-metal (or in the absence of metal, sulfide), rather than co-existing silicates or oxides, at low pressures. Experimental studies at relatively low pressures (105 Pascals), to pressures as high as 18 GPa, consistently show liquid metal/liquid silicate concentration ratios (D values) varying from 103 to 106 (e.g., Kimura et al. 1974; Jones and Drake 1986; Borisov et al. 1994; O’Neill et al. 1995; Holzheid et al. 2000; Ertel et al. 2001; Fortenfant et al. 2003; Yokoyama et al. 2009; Mann et al. 2012). The variations in D values partly reflect experimental charge issues (e.g., Brenan et al. 2016, this volume), but recent combined HSE experiments indicate significant interelement fractionation at a range of temperatures and pressures during metal–silicate equilibration (Mann et al. 2012). Consequently, the relative and absolute HSE abundances in planetary cores, mantles, and crusts might be expected to dominantly reflect metal–silicate equilibration at the conditions (e.g., pressure, temperature, oxygen fugacity, composition) relevant to the planetary body. Despite the fundamental control of core formation on HSE fractionation during planetary differentiation, terrestrial mantle peridotites and estimates of lunar and martian mantle compositions do not appear to solely record the effects of this process, and instead define mantles with broadly chondritic-relative abundances of the HSE (Fig. 1) (See section below, ‘What does chondritic or nearly/broadly chondritic actually mean?’). Chondritic-relative HSE abundances in terrestrial, martian and lunar mantles are supported by long-term Re/ Os and Pt/Os for the terrestrial mantle within ~± 5% and ~± 10% of chondrites, respectively (e.g., Brandon et al. 2006; Walker 2009), and Re/Os within ~± 10% for at least some portions of the martian and lunar mantles (Brandon et al. 2012; Day and Walker 2015). Furthermore, there is compelling evidence for similar crust-mantle partitioning behavior for Earth and the Moon (Day et al. 2010). These observations indicate that a process other than metal– silicate equilibration during core formation was involved in setting initial HSE abundances in planetary mantles, with the most likely candidate being late accretion of chondritic impactors after the major phases of core formation (Turekian and Clark 1969; Kimura et al. 1974; Chou 1978; Wanke 1981), although other classes of model, including incomplete core separation (Jones and Drake 1986), lowered metal–silicate D values (Ringwood 1977; Brett 1984; Murthy 1991), and variants of these models (e.g., Kramers 1998) have also been proposed. Combined with precise chronometry, it is possible to examine the timing of key planetary growth episodes and post-core formation (late) accretion addition using the HSE (Fig. 2; Day et al. 2012a). These initial parameters for establishing core–mantle–crust HSE abundances in planetary bodies are a key motivation for measuring the HSE in planetary materials, as they reveal the fundamentals of planetary accretion and differentiation. Within differentiated planetary mantles, the HSE exhibit contrasting behaviors during melting. The HSE comprise Re and Au, along with the platinum group elements (PGE), defined as Os, Ir, Ru, Rh (the so-called Ir-group PGE, or IPGE; Barnes et al. 1985) and Pt and Pd (the Pt-group PGE, or PPGE). The PPGE (melting temperature < 2000 °C), Re and Au are typically more incompatible during melting and crystallization relative to the IPGE (melting temperature >2000 °C; Barnes et al. 1985). For this reason, studies of the cosmochemical behavior of the HSE will typically list the HSE in order of melting temperature of the pure metal, or 50% condensation temperatures at pressures appropriate for the nebula (e.g., Fig. 1), whereas studies using the HSE to investigate mantle melting processes on Earth will typically order the HSE according to relative incompatibility during melting, which is generally considered to be: Re ≤ Au < Pd < Pt ≤ Rh < Ru ≤ Ir ≤ Os (Pearson et al. 2004; Becker et al. 2006; Fischer-Gödde et al. 2011). While most other planetary bodies appear to show similar relative incompatibilities, the oxidation state of their mantle plays a major control on melting behavior (Birck and Allègre 1994; Day and Walker 2015). The main controls on HSE fractionation behavior during partial

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Figure 2. Schematic diagram illustrating the distinction between core-formation accretion and post-core formation accretion, also known as late accretion (or, in some cases, a 'late veneer' for Earth). During the main phase of accretion, in the first 1–2 Ma of Solar System history, as the total mass of the body is predicted to increase rapidly, metal–silicate equilibration is expected to draw-down the HSE into a metallic core. After the main differentiation phase has completed and core formation ceases, late accretion commences, with a proportionally smaller predicted fraction of ‘source material’ being contributed into the silicate crust and mantles of planetary bodies. In the most general sense, the amount of late accretion added to a differentiated planet will be a function of cessation of metal–silicate equilibration (the main differentiation phase).

melting of the mantle are sulfide phases, platinum-group element minerals (PGM) (e.g., Jagoutz et al. 1979; Lorand et al. 2013; O'Driscoll and Gonzalez-Jimenez, 2016, this volume), and the presence or absence of residual metal. Experimental determination of sulfide–silicate partitioning of the HSE (with the exception of Re) are generally >104 (e.g., Peach et al. 1990, 1994) and the presence or absence of sulfide in the system plays a major role in HSE fractionation (e.g., Rehkämper et al. 1999). The HSE can also be used to track later differentiation and melting events in planetary bodies, after the major phases of core formation and accretion. Combined with two long-lived radiogenic decay schemes embedded within the HSE, the 187Re–187Os and 190 Pt–186Os systems, as well as the short-lived 107Pd–107Ag system, these elements offer powerful geochemical tools for interrogating planetary formation, accretion and differentiation processes.

METHODS APPLIED TO INVESTIGATING SIDEROPHILE ELEMENTS IN PLANETARY MATERIALS While not exhaustive, this section provides an overview of analytical techniques for the HSE, as well as modeling methods aimed at understanding fundamental processes, including melting processes in the presence of metallic liquids (core-formation processes), silicate partial-melting and estimation of planetary mantle HSE abundances. Some other ‘tools’ for examining planetary materials are given in later sections (e.g., estimating compositions of impactors from lunar impact melt breccia compositions). Analytical methods are also discussed in Meisel and Horan (2016, this volume).

HSE ABUNDANCES Despite important data on Au and/or Rh abundances obtained by standard addition (e.g., Walker et al. 2008; Fischer-Gödde et al. 2010, 2011), or neutron activation analyses (e.g., Wolf et al. 1979), these elements are mono-isotopic and are not reported in most studies that measure HSE abundances using isotope-dilution methodologies. This review primarily focuses on the coupled information derived from HSE abundances and Os isotopes, with an emphasis on Os,

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Ir, Ru, Pt, Pd, and Re abundances. Therefore, the focus is on preparation and analysis methods utilizing isotope dilution. Where available, Rh and Au data are discussed, but the reader should note that the analysis of these elements, along with the chalcogens (S, Te, Se; see Harvey and Day 2016, this volume) are currently rarely determined in tandem with Os isotopes, especially in low abundance planetary materials. This is due to the need for generally high abundances of Rh and Au for reliable and accurate internal–external standardization (comparison with the 193 Ir of the sample and external standard solutions; Meisel et al. 2003; Fischer-Gödde et al. 2010), or for low analytical blank work required for S, Te, and Se. A technique that allows for the combined determination of all HSE abundances, as well as other highly siderophile and chalcophile elements (e.g., Ni, Co, W, Mo), is laser-ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). LA-ICP-MS is a particularly useful technique for in situ analysis of individual sulfide and metal phases that have significant concentrations (>μg.g−1 level) of the HSE (e.g., Campbell and Humayun 1999), and this method of analysis has been extensively applied to the study of iron meteorites (e.g., Petaev and Jacobsen 2004; Walker et al. 2008), chondrites (e.g., Campbell et al. 2002; Humayun 2012), as well as some achondrite meteorite groups (e.g., van Acken et al. 2012; Day et al. 2012b; Yang et al. 2015). Because of the high affinity of the HSE in metal, LA-ICP-MS also allows for the in situ analysis of HSE abundances in metal or sulfide grains, even in rocks with relatively low bulk HSE abundances (cf., Alard et al. 2005). For example, Day et al. (2012a) used LA-ICP-MS analysis of metal grains in diogenite meteorites to show that the HSE were almost exclusively hosted in metal within brecciated clasts, with only ~0.001–0.01 mass % metal required to explain whole-

Figure 3. Example of the use of combined petrology and LA-ICP-MS analyses to investigate the location of the HSE within rock samples. The left panel (a) shows a back-scatter electron image of a portion of olivine diogenite MIL 07001, 26, showing the textural relationship of metal and sulfide in the sample. Top right panel (b) shows the field of metal grain HSE abundances measured for MIL 07001 versus whole-rock measurements of MIL 07001 and other diogenites (shown as lines), illustrating the strong control metal grains have on the whole-rock composition of diogenites. Thick solid line in (b) is the terrestrial primitive mantle composition (Becker et al. 2006). The siting of the metal and sulfide grains in brecciated diogenites is not in the breccia matrix, indicating setting of the HSE prior to crystallisation of the diogenites. (c) Relationship between Ni/Fe in metal grains from diogenite meteorites and their whole-rock HSE abundances (shown here in terms of Os concentration), emphasizing the siting of the HSE within metals in these particular meteorites. Figure adapted from Day et al. (2012a).

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rock HSE abundances (Fig. 3). Combined with detailed petrography and petrology showing that the metal grains were situated within crystalline portions of the diogenite meteorites, this study was able to establish early addition (within 2–3 Ma of Solar System formation) of the HSE to the diogenite parent body. LA-ICP-MS is therefore a powerful technique for understanding planetary HSE inventories, especially when used in combination with detailed petrographic, mineral chemistry and whole-rock studies. Methods for the preparation of whole rocks, mineral phases and other components (e.g., glasses, clasts) of planetary samples for HSE analysis include irradiation for neutron activation analysis and spectroscopic techniques (e.g., Goldberg et al. 1952; Lovering et al. 1957; Wasson 1967), nickel-sulfide fire assay (e.g., Ravizza and Pyle 1997), alkali fusion (Morgan and Walker 1989), selective leaching techniques (e.g., Walker et al. 2004) and dissolution in reducing media (e.g., HF–HBr; Birck et al. 1997). Neutron activation analysis is a particularly effective method for analyzing Ir and Au, which can be measured without dissolving the sample, or applying radiochemical separation procedures. Other HSE are less-effectively measured by this method, and the preferential analysis of Ir is partly responsible for the misconception of ‘Ir anomalies’ defining the Cretaceous–Palaeogene boundary (e.g., Alvarez et al. 1980), when, in fact, all the HSE are enriched in these layers (Goderis et al. 2013). The most popular current method of analysis is through digestion in concentrated HCl–HNO3 mixtures in Carius tubes at temperatures ≤ 270 °C, or using quartz vessels that are sealed and heated in a High-Pressure Asher (HPA) device at temperatures ≤ 320 °C and pressures ≤ 15 MPa (Shirey and Walker 1995; Meisel et al. 2003). The Carius tube and HPA digestion methods allow for the most complete digestion in combination with the lowest total analytical blanks (e.g., Reisberg and Meisel 2002) and allow equilibration of spike–sample mixtures while completely retaining volatile OsO4 species during oxidation of the sample. For example, analysis of extremely low abundances of the HSE in whole-rock lunar materials has been made possible through low total analytical blank HPA analysis and through the thorough pre-cleaning of Carius tubes (Day and Walker 2015a). Procedures for the chemical purification and analysis of the HSE are varied, but the most typical modern technique after Carius tube/HPA digestion is a solvent extraction (Cohen and Waters 1996) followed by a micro-distillation procedure (Birck et al. 1997) to separate and purify Os, and by an anion exchange column chemistry procedure to provide pure Ir, Ru, Pt, Pd, Re (± Rh, ± Au) separations (e.g., Rehkämper and Halliday 1997; Pearson and Woodland 2000; Reisberg and Meisel 2002; Fischer-Gödde et al. 2010). Concentrations of Ir, Ru, Pt, Pd, Re, Rh, and Au are then typically measured using inductively coupled plasma mass spectrometry techniques that allow the precise analysis of isotopic ratios of spike-sample mixtures (see also Meisel and Horan 2016, this volume). Studies have shown that an additional HF-digestion step is sometimes necessary to extract all the Re in some silicate rocks, as the Carius tube/HPA HCl–HNO3 acid mixture is ineffective at breaking down silica-bonds (Ishikawa et al. 2014). This method typically has to be applied after Os extraction due to the danger of residual fluorides attacking the sides of the Pyrex or quartz Carius tube or HPA vessel walls, or from the early loss of Os, as OsO4. An HF digestion-step has been reported to access up to 9–15% more Re from within silicate phases, compared with Carius tube/HPA digestion (Li et al. 2014), at least for some basaltic rocks. There is, however, a particular disadvantage to this method for ancient samples (i.e., most planetary materials), as measured Re/Os or Pt/Os obtained by HF silicate digestion after Os extraction will not reflect those in equilibrium with measured 187Os/188Os or 186Os/188Os, leading to the potential for young apparent ages and the loss of meaningful Re–Os or Pt–Os isotope chronology (Day et al. 2015).

The rhenium–osmium, platinum–osmium and palladium–silver isotope systems The ability of relative and absolute HSE abundances to record processes acting on rock materials are complemented by the existence of the long-lived 190Pt–186Os (190Pt → 186Os + α + Q; λ = 1.48 × 10−12 a−1; Walker et al. 1997) and 187Re–187Os

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(187Re–187Os + β- + ῡ; λ = 1.6668 × 10−11 a−1; Selby et al. 2007) chronometers. Both longlived radiogenically produced isotopes of Os are minor constituents of rocks with broadly chondritic Re/Os and Pt/Os (186Os = 1.6%; 187Os = 1.5%; Shirey and Walker 1998). In the case of the 187Re–187Os system, where 187Re is a major isotope (62.6%) of Re, and has a half-life of ~42 Ga, the range of natural materials spans several orders of magnitude and 187 Os/188Os can reasonably range from a Solar System initial ratio of 0.0952 ± 0.0002 to nearly pure 187Os, derived from samples essentially devoid of non-radiogenic Os and with high concentrations of Re (e.g., molybdenite; Luck and Allègre 1982). This characteristic means that the difference of 187Os/188Os between natural samples allows routine analysis of low Os abundance samples to percent precision or better, with the most widely used method of analysis being negative thermal ionisation mass spectrometry (N-TIMS; Völkening et al. 1991; Creaser et al. 1991) either by peak-jumping of small signals (~ 3 mV 192OsO3− or less) using a secondary electron multiplier, or by static or dynamic measurements of larger signals (typically > 0.5 V 192OsO3−) using Faraday collectors (see Shirey and Walker 1998; Reisberg and Meisel 2002; Carlson 2005). Publications pre-dating the exploitation of 186 Os/188Os variations often reported 187Os/186Os variations, assuming 186Os to be a stable isotope. Although radiogenic ingrowth corrections on 186Os are generally trivial, all studies now report 187Os/188Os. A correction factor of 0.12035 is commonly applied to convert 187 Os/186Os to 187Os/188Os (e.g., 1.0553 (187Os/186Os) × 0.12035 = 0.1270 (187Os/188Os)). By contrast with the 187Re–187Os decay system, 190Pt is a minor isotope of Pt (0.0129%) and has a much longer half-life (~470 Ga), so 186Os/188Os variations in nature are generally small. For example, in the terrestrial mantle variations are of the order of ~0.00015%, with an ‘average’ mantle value of 0.119837 ± 5 (2σ) (Brandon et al. 2006). The typically minor variations of 186Os/188Os in planetary materials require external analytical precision of better than 30 ppm. To obtain sufficient analytical precision, comparatively large quantities of Os are needed (typically 50–75 ng of Os) to generate sufficient signals on 186Os given the typical ionisation efficiency of Os by N-TIMS (~2–6%; Creaser et al. 1991). For example, to generate a stable 100 mV signal on 186Os (1.6% of chondritic Os), a ~2.6 V signal on 192 Os (the most abundant natural isotope of Os at 40.98%) is required. For this reason, and due to the precious nature of planetary materials, most studies utilizing high-precision measurement of Os have focused on meteorites characterized by high concentrations, such as iron meteorites (e.g., Cook et al. 2004) or chondrites (e.g., Brandon et al. 2005a). In addition to sample issues, some studies have argued for polyatomic interferences, leading to complications with high-precision 186Os/188Os measurements (e.g., Luguet et al., 2008). Inevitably, these forms of analytical challenge mean that there are far fewer high-precision 186 Os/188Os data currently available than there are for 187Os/188Os. Nonetheless, 186Os/188Os, in conjunction with 187Os/188Os and HSE abundances have the potential to provide robust constraints on planetary mantle differentiation processes. In addition to the Re–Os and Pt–Os isotope systems, the now-extinct isotope 107Pd decays to 107Ag with a half-life of 6.5 Ma. The Pd–Ag isotope system can be used to define the chronology of iron meteorite formation and the timing and mechanisms of early volatile depletion, since Ag is a moderately volatile siderophile element (50% condensation temperature = 996 K). As a result, some volatile-depleted iron meteorites have Pd/Ag >105, compared with a solar Pd/Ag ~3, leading to very high 107Ag/109Ag in some irons (Chen and Wasserburg 1996), and measureable differences in carbonaceous chondrites (Schönbachler et al. 2008). The resulting ages of iron meteorites defined by Pd–Ag chronology are ~9–20 Ma after Solar System formation and are longer than Hf–W isotope chronology (~1–3 Ma) (e.g., Kruijer et al. 2014). The varying ages are interpreted to reflect the chronology of different events: metal–silicate differentiation in the case of Hf–W, and the timing of volatile depletion in the case of Pd–Ag (Horan et al. 2012). It should be noted, however, that large uncertainties may exist due to neutron capture reactions on Ag in meteorites (Leya and Masarik 2013).

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Standardization in planetary studies In general, well-calibrated isotope dilution methodologies and efficient digestion and spike– sample equilibration procedures will lead to the determination of reliable HSE concentrations in most materials. Nonetheless, standard reference materials run at regular intervals within geochemical studies offer useful indicators of the efficiency and reliability of digestion procedures, chemistry and analytical methods. There are three issues that make the selection of standard materials challenging for planetary materials. First, the large range in HSE concentrations of planetary materials, from approximately μg.g−1 levels in chondrites and iron meteorites, to pg.g−1 levels in some lunar rocks, means that no single standard will be appropriate for regular analysis with this diversity of material types. Second, clear differences in bulk composition relative to terrestrial materials make matrix-matching difficult. For example, matrices (iron metal, carbonaceous chondrite, lunar anorthosite), or oxidation state (cf., terrestrial basalts are approximately at the quartz–magnetite–fayalite (QFM) buffer, versus iron–wüstite (IW) minus six to eight for enstatite achondrites) all pose challenges to selection of an appropriate reference standard. Third, fundamental differences in partitioning behavior result in terrestrial basalts having higher Re/Os than most planetary basalts, making their use as standards improper. The available masses of some chondrites and iron meteorites make them ideal reference materials for comparison with other iron meteorites, chondrites, or primitive achondrites. The Allende CV3 chondrite meteorite has been measured in numerous studies as a reference material for comparison with chondrites and primitive achondrite meteorites (Table 1). Iron meteorites such as Filomena, Coahuila and Hoba are also regularly used as natural standards for LA-ICP-MS analysis. Some studies of differentiated achondrites and Apollo samples have utilized terrestrial peridotites (e.g., GP13, UB-N, HARZ-01) or basalt (TDB-1). However, the higher Pt, Pd, and Re concentrations of most terrestrial basaltic standards (e.g., Meisel and Moser 2004a,b) make their use non-ideal in the study of ‘basaltic’ achondrites (e.g., lunar crust or mare basalt, angrites, eucrites, shergottite–nakhla–chassignite [SNC] meteorites), and TDB-1, in particular, is a very heterogeneous diabase sample (Ishikawa et al. 2014). To circumvent this issue, recent studies (Day et al. 2010; Riches et al. 2012; Day and Walker 2015) have analyzed samples in duplicate or triplicate, demonstrating reproducibility of analysis for these planetary materials, as well as enabling optimal spiking of samples.

Metal-sulfide–silicate modeling in chondritic systems The HSE are both strongly siderophile and chalcophile, and are sensitive tracers of the earliest stages of chondritic melting, when the first melts to be generated are Fe–Ni–S-rich (e.g., Mittlefehldt et al. 1996). Therefore, understanding the effects of metallic composition on partitioning behavior is important for any processes involving metallic liquid separation. Examples of such processes include asteroidal core formation (e.g., Jones and Drake 1982; Wasson 1999), or the separation of metallic core from silicate mantle during core formation (e.g., Righter and Drake 1997; Li and Agee 1996). Distributions of the HSE between a liquid metal phase and solid metal restite during partial melting in the Fe–Ni–S system are strongly dependent on the S-content of extracted melts, and on C and P that enter the metallic phase (e.g., Chabot and Jones 2003; Chabot et al. 2014). Parameterization of solid-metal–liquid-metal D-values has been established by several authors as a function of the metallic liquid composition, with the most commonly used iteration derived from Chabot and Jones (2003). Studies utilizing this parameterization have shown that these forms of models can be generally applied to studies of chondritic components (Horan et al. 2009), iron meteorites (Walker et al. 2008 McCoy et al. 2011), and some primitive achondrite groups (Rankenburg et al. 2008; Day et al. 2012b). An example of the method is shown in Figure 4, where a CI-chondrite (Orgueil) starting composition is used with Pd/Os ~1.6, and where it is assumed that all HSE were initially

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Table 1. Analyses of Os isotopes and HSE (in ng.g−1) in the Allende CV 3 carbonaceous chondrite and comparison with standards used for planetary materials research. Sample

Lithology

Allende

CV3 Chondrite

Compared with

Method

UB-N

GP13

CT

749

704

CT

767

716

Chondrites

CT

717

1016

CT

785

720

1118

Chondrites

CT

689

644

1007

Chondrites

HPA

795

705

1073

Chondrites

HPA

732

660

997

Chondrites

HPA

709

762

997

Chondrites

HPA

703

654

991

Chondrites

HPA

749

680

975

LEW 88763

CT

738

718

1031

LEW 88763

CT

666

657

995

Peridotites

CT

763

700

1140

13

737 ± 78

695 ± 68

1031 ± 110

CT

3

759 ± 27

691 ± 66

1012 ± 10

HPA

13

3.71 ± 0.53

3.37 ± 0.43

6.30 ± 0.58

CT

4

3.85 ± 0.32

3.58 ± 0.40

6.93 ± 0.47

Spinel-Bearing Serpentinite

LIMB

Harzburgite

TDB-1

Basalt

Ru (± 2SE)

Tagish Lake

Lunar glasses

HARZ 01

Ir (± 2SE)

Chondrites

Low Level Dilution (0.19%)

Spinel Lherzolite

Os (± 2SE)

Chondrites

Average Allende Allende

n

Peridotites

CT

4

3.51 ± 0.25

3.26 ± 0.25

6.51 ± 0.65

Peridotites

CT/ HPA

19

3.53 ± 0.5

3.16 ± 0.44

6.43 ± 0.76

Peridotites

CT

6

3.66 ± 0.3

3.24 ± 0.62

6.48 ± 0.58

Intrusive Rocks

CT

5

3.61 ± 0.12

3.70 ± 0.35

7.66 ± 0.87

Peridotites

CT

8

3.87 ± 0.17

3.56 ± 0.33

6.97 ± 0.23

HPA

4

4.06 ± 0.03

3.33 ± 0.09

6.25 ± 0.39

Diogenites

CT+ HPA

9

3.36 ± 0.20

3.32 ± 0.40

6.92 ± 0.41

Diogenites

CT

8

4.21 ± 1.30

3.38 ± 1.36

5.73 ± 2.19

HPA

7

0.12 ± 0.02

0.07 ± 0.02

0.20 ± 0.02

Achondrites

HPA

7

0.11 ± 0.02

0.06 ± 0.01

0.23

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Table 1. (cont'd) Analyses of Os isotopes and HSE (in ng.g−1) in the Allende CV 3 carbonaceous chondrite and comparison with standards used for planetary materials research. Sample

Lithology

Compared with

Method

Pt (± 2SE)

Pd (± 2SE)

Re (± 2SE)

Allende

CV3 Chondrite

Chondrites

CT

1336

662

60.1

Chondrites

CT

1345

682

61.8

Chondrites

CT

1364

678

Tagish Lake

CT

1421

682

63.5

Chondrites

CT

1321

652

62.0

Chondrites

HPA

1380

681

64.0

Chondrites

HPA

1334

684

58.7

Chondrites

HPA

1409

676

58.4

Chondrites

HPA

1321

692

58.6

Chondrites

HPA

1379

657

61.4

LEW 88763

CT

1503

638

60.3

LEW 88763

CT

1385

627

54.5

Peridotites

CT

1379

786

61.2

13

1375 ± 100

677 ± 76

60 ± 5

CT

3

1276 ± 65

698 ± 92

61 ± 4

HPA

13

7.42 ± 0.60

6.11 ± 0.36

0.206 ± 0.010

Average Allende Allende

UB-N

GP13

Low Level Dilution (0.19%) Spinel-Bearing Serpentinite

Spinel Lherzolite

Lunar glasses

LIMB

CT

4

7.47 ± 0.16

5.70 ± 0.12

0.213 ± 0.011

Peridotites

CT

4

7.00 ± 0.46

5.85 ± 0.53

0.205 ± 0.008

Peridotites

CT/ HPA

19

7.31 ± 0.94

5.85 ± 0.40

0.188 ± 0.048

Peridotites

CT

6

8.07 ± 2.34

6.17 ± 0.50

0.205 ± 0.014

Intrusive Rocks

CT

5

7.77 ± 0.70

5.71 ± 0.66

0.291 ± 0.013

Peridotites

HARZ 01 TDB-1

Harzburgite

n

CT

8

7.00 ± 0.52

5.64 ± 0.35

0.330 ± 0.010

HPA

4

6.69 ± 0.69

5.68 ± 0.27

0.312 ± 0.010

Diogenites

CT+ HPA

9

6.89 ± 1.05

5.46 ± 0.86

0.286 ± 0.013

Diogenites

CT

8

6.71 ± 2.69

5.12

0.062 ± 0.023

HPA

7

5.01 ± 0.36

24.4

0.79 ± 0.005

HPA

7

4.7 ± 1.2

22.3

1.01 ± 0.008

Basalt Achondrites

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Table 1. (cont'd) Analyses of Os isotopes and HSE (in ng.g−1) in the Allende CV 3 carbonaceous chondrite and comparison with standards used for planetary materials research. Sample

Lithology

Compared with

Method

Allende

CV3 Chondrite

Chondrites

CT

0.12615

Chondrites

CT

0.12614

Chondrites

CT

Tagish Lake

CT

UB-N

GP13

Low Level Dilution (0.19%) Spinel-Bearing Serpentinite

Spinel Lherzolite

TDB-1

Harzburgite

187

Os/188Os (± 2σ)

Ref. [1] [1] [1]

0.12596 ± 0.00001

[2]

CT

123

0.12624

[3]

Chondrites

HPA

140

0.12596

[3]

Chondrites

HPA

144

0.12590

[3]

Chondrites

HPA

152

0.12572

[3]

Chondrites

HPA

135

0.12614

[3]

Chondrites

HPA

0.12569

[3]

LEW 88763

CT

0.12586 ± 0.00007

[4]

LEW 88763

CT

0.12588 ± 0.00007

[4]

Peridotites

CT

0.12638

[5]

Lunar glasses

13 CT

3

HPA

13

139 ± 22

0.12600 ± 0.00042 [6]

-

[7]

LIMB

CT

4

0.1272 ± 0.0004

[8]

Peridotites

CT

4

0.1274 ± 0.0005

[5]

Peridotites

CT/ HPA

19

0.1272 ± 0.0006

[9]

Peridotites

CT

6

-

[10]

Intrusive Rocks

CT

5

0.1263 ± 0.0001

[11] [12]

Peridotites

HARZ 01

Au (± 2SE)

Chondrites

Average Allende Allende

n

CT

8

0.1262

HPA

4

-

[7]

Diogenites

CT+ HPA

9

0.1263 ± 0.0008

[13]

Diogenites

CT

8

0.1254 ± 0.0014

[13]

HPA

7

HPA

7

0.98 ± 0.13

[13]

Basalt Achondrites

[7]

References: [1] Walker et al. (2002); Horan et al. (2003) [2] Brandon et al. (2005a) [3] Fischer-Gödde et al. (2010) [4] Day et al. (2015b) [5] Becker et al. (2006) [6] Walker et al. (2004) [7] Meisel and Moser (2004a) [8] Puchtel et al. (2008) [9] Fischer-Gödde et al. 2011 [10] Luguet et al. (2007) [11] Day et al. (2008) [12] Pearson et al. (2004) [13] Dale et al. (2012). Definition of terms: HPA = high-pressure asher; CT = Carius tube; LIMB = lunar impact melt breccia.

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Figure 4. Plots of Pd/Os versus (a) Pt/Os and (b) Ir/Os for primitive achondrites and chondrites, with calculated compositions of metal residues. Shown are chondrite (Horan et al. 2003; Fischer-Gödde et al. 2010) ureilite (Rankenburg et al. 2008) and brachinite-like achondrite compositions (Day et al. 2012b). The melt calculations model the composition of residues that result from single episodes of batch melting. The starting HSE composition is the bulk composition of Orgueil with concentrations adjusted assuming that all of the HSE were originally in metal and that metal comprises 5% of the bulk. Curves show compositions of residues resulting from no sulfur, and 25% sulfur. Fractions of residue are labelled and are in increments of 5%. Solid-metal–liquid-metal D values were calculated using the parameterization of Chabot and Jones (2003). Carbonaceous chondrite [CC]; enstatite chondrite [EC]; ordinary chondrite [OC].

sited in metal. These types of model typically assume HSE partitioning only between solid metal and liquid metal, consistent with high D-values for the HSE (>104) for metal–silicate at the low pressures of asteroidal interiors (e.g., Righter 2003). It should be noted that the choice of starting composition is important, because the range of initial interelement fractionation observed in chondrites is quite significant (Pd/Os = 0.88–1.69; see below). Since HSE preferentially partition into metal, the HSE concentration present in the metal will depend on the percentage of metal in the parent body, which is also a reflection of its oxidation state. To estimate initial metal abundance in primitive achondrites, Rankenburg et al. (2008) compared total Fe and olivine composition in ureilites with the total Fe content of carbonaceous chondrites (e.g., Orgueil), which have relatively constant total Fe of ~24.5 ± 1.5 wt.% (Jarosewich 1990). As described in Walker et al. (2008) there is a hybrid method to that of the parameterization method of Chabot and Jones (2003) for estimating changing D values for the HSE, which involves observation of linear trends on logarithmic plots of the HSE (e.g., Cook et al. 2004). This technique is only really applicable to iron meteorites, as the initial D(Ir) is normally chosen as an ‘anchor’ because its solid metal-liquid metal partitioning is the best experimentally constrained of the HSE and as primitive achondrite and ‘anatectic’ chondrite datasets are often not as well defined as linear HSE trends for iron meteorites.

Partial melt modeling of planetary mantles Partial melting generally promotes HSE fractionation (Barnes et al. 1985; Rehkämper et al. 1999), as shown by the differences in relative and absolute HSE abundances for terrestrial high-degree partial melts (e.g., komatiites) versus lower-degree equivalents (e.g., alkali basalts; Day 2013). The HSE are strongly controlled by sulfide and HSE-rich alloys (formed by incongruent melting of sulfide) during partial melting of the terrestrial mantle

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(Barnes et al. 1985; Ballhaus et al. 2006). Numerous studies have modeled HSE behavior during terrestrial mantle melting (e.g., Fryer and Greenough 1992; Rehkämper et al. 1999; Bézos et al. 2005), showing that a range of physical and chemical factors control resultant HSE abundances in primary melts. In addition to the composition of the mantle source, and the strong control of sulfide during melting, the ‘shape’ and dynamic processes acting on the melting regime are important (e.g., O’Hara 1985). These observations have led to the concept of two different shaped primary melt regimes. One is batch-melting of a columnar (cylindrical) melting regime, assuming uniform melting of a mantle source and extraction of all the S-hosted HSE by magmas once sulfide is completely exhausted (20–25% partial melting in the terrestrial mantle, for example). This form of model reproduces the compositions of both terrestrial low-degree alkali basalts and high-degree partial melts, such as komatiites, reasonably well (Rehkämper et al. 1999), and is most consistent with an upwelling ‘mantle plume-like’ melting regime on Earth, or for most forms of basaltic melting on planetary bodies (Fig. 5). An alternative shaped melting regime that is applicable to adiabatic decompression melting is the triangular or cornerflow melting regime. This model assumes initial melting commences when mantle material crosses the solidus and where partial melting increases with vertical displacement, and assumes near-fractional melting, typically in 1% increments (Rehkämper et al. 1999). In this melting regime, mixing between deeper, S-saturated low-degree partial melts with low HSEconcentrations, and shallower, higher-degree partial melt with potentially S-undersaturated and HSE-rich compositions can occur. The triangular melting regime generates different outcomes for calculated HSE abundances for samples relative to the columnar melting regime, because of the hybridisation of melts. The HSE systematics are critically dependent on the presence or elimination of sulfides in the residue, so blended melts in a triangular melt regime will lead to more elevated HSE abundances at lower degrees of partial melting (10–20%) that better conform to the melting characteristics of some terrestrial tholeiitic lavas. In a mantle with 250 μg.g−1 S and 4.4 ng.g−1 Pd, this translates to generation of melts with 1000 μg.g−1 S in a triangular melting regime that form by 13% partial melting and contain 2 ng.g−1 Pd, versus > 20% partial melting required in a columnar melting regime to obtain similar Pd enrichments in the resultant magma. In these melting models, the key factor is whether partial melting was significant enough to exhaust residual sulfides in any part of the melting regime (Rehkämper et al. 1999; Mavrogenes and O’Neill 1999). Refined models developed for HSE abundances generated by partial melting to produce MORB (Bézos et al. 2005) and hotspot volcanics (Rehkämper et al. 1999) have allowed elucidation of distinct melt regimes. For most planetary examples, melting regimes are likely to conform to a columnar melting condition due to the absence of obvious plate-tectonic processes on Solar System bodies, other than Earth. Estimating the degree of partial melting for many asteroidal parent bodies is challenging because of uncertainties regarding source composition and S content in their sources. Estimates of the degree of partial melting and S content have been produced for lunar and martian melt sources, with martian melts having similar estimated degrees of partial melting and S content to terrestrial basalts considered to derive from columnar melting regimes (Fig. 5). The HSE contents of martian melts indicate similar absolute abundances of these elements in their source, with sulfide-melt partitioning and degree of S-undersaturation during partial melting similar to terrestrial basalts. By contrast, the lower HSE contents of lunar basalts, and their lower estimated degrees of partial melting either require much higher sulfide–melt partitioning, or lower initial HSE and S abundances in the lunar mantle. As discussed below, detailed work indicates the latter scenario as the cause of differences in melting models for lunar, compared with terrestrial or martian basalts.

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Figure 5. Melting models for terrestrial Pd concentrations as a function of partial melting (F) for different sulfide–melt partitioning (1000, 10,000, 100,000) in a columnar melting regime. Terrestrial volcanic rocks ranging from komatiites to tholeiites and alkali basalts are shown as dark gray boxes (from Day 2013). The field of estimated partial melting and Pd concentrations of martian shergottites and lunar mare basalts are from Brandon et al. (2012) and Day and Walker (2015), respectively, with melt contents estimated from shergottites and lunar mare basalts with a range of MgO contents (< 10 to >19 wt. % MgO). The lower solid line is sulfide-melt partitioning assuming a lower S concentration in the lu nar mantle (~75 mg.g−1) relative to the terrestrial or martian mantles (> 200 mg.g−1) and lower initial mantle source composition (Day and Walker 2015). Palladium is assumed to be perfectly incompatible in silicates and the terrestrial model assumes a primitive mantle Pd composition.

“Pristinity” of crustal and mantle samples Studies of lunar crustal samples have recognized the importance of identifying samples that experienced limited impactor contamination on the lunar surface. These studies have used Ir contents, petrography and other geochemical arguments to establish a ‘pristinity’ filter (e.g., Chao et al. 1976; Hertogen et al. 1977; Warren and Wasson 1977). The Ir abundance filter for ‘pristinity’ ( 105 enrichment), so even limited (< 0.1%) meteoritic addition will dominate the HSE inventory of impact-contaminated rocks. The HSE in particular, are sensitive tracers of impactor contamination in rocks with initially low HSE abundances. Osmium isotopic composition is especially diagnostic of meteoritic contamination because the limited range in Re/Os among chondrites—and some iron meteorites—leads to a restricted, well-defined range in presentday 187Os/188Os (e.g., Fig. 6). Additionally, chondritic materials (as well as some irons) are characterised by limited variations in the absolute abundances of these elements (1400 K (Table 3; the 50% condensation temperature (50%Tc) is a term used to describe the extent of condensation or volatility of minor and trace elements and can also be computed for major elements. At this temperature, half of an element is in the gas phase and the other half is sequestered into condensates (Lodders 2003)). Because of their low initial abundances, these elements are considered to have condensed within refractory metal alloys in a predictable fashion (Palme and Wlotzka 1976; Fegley and Palme 1985; Campbell et al. 2001). By contrast, two of the HSE (Pd, Au), as well as numerous moderately and slightly siderophile elements (Ag, As, Bi, Co, Cr, Cu, Ge, Ni, P, Pb, Sb, Sn, Te) have lower 50% Tc and are considered to have become sequestered in Fe–Ni alloy phases (pure Fe 50%Tc = 1357 K; Lodders 2003). Combined with the moderately volatile nature of Au at nebular conditions, these factors can lead to a fundamental difference in behavior between Au and Pd, and the rest of the HSE. Analysis of nano-meter sized metal alloys in spinel grains from Allende have HSE, W and Mo concentrations that are in excellent agreement with calculated concentrations from condensation in a cooling gas of solar composition, supporting the theoretical calculations (Eisenhour and Buseck 1992). Subsequently, exsolution, oxidation and sulfurisation led to the formation of complex assemblages in some early formed chondritic materials (Bischoff and Palme 1987). The most primitive and least modified accessible materials from the Solar System are represented by chondrite meteorites. Chondrites can be divided on the basis of their bulk

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Table 4. A simplified list of planetary materials and examples of HSE studies applied to them. Meteorite Type (asteroidal, unless stated)

Typical mode of origin

Sub-divisions

Examples of HSE and Os isotope studies

Earliest Time of formation

Primitive meteorites (witnessed limited/no igneous differentiation) Chondrite*

Ordinary (H, L, LL) [~85% of all meteorite finds]

"Cosmic sedimentation"

Early Solar System

Carbonaceous (CI, CM, CR, CB, CH, CV, CO, CV) [~5%]

"Cosmic sedimentation"

Early Solar System

Enstatite EH, EL) [~2%]

"Cosmic sedimentation"

Early Solar System

Ungrouped, Rumuruti or Kakangari-Type

"Cosmic sedimentation"

Early Solar System

Walker et al. (2002); Horan et al. (2003, 2009); Brandon et al. (2005a,b); Fischer-Godde et al. (2010); Archer et al. (2014)

Chondriteimpact melt rocks

-

Impacts of chondrite asteroids

Early to late

Norman and Mittlefehldt (2002)

Winonaites

-

Metamorphism of chondrites

Early Solar System

Iron meteorites

Primitive Irons (e.g., IAB, IIE, IIICD)

Partial melt differentiated core?

Early Solar System

Magmatic Irons (e.g., IAB, IC, IIAB, IIC, IID, IIE, IIIAB, IIICD, IIIE, IIIF, IVA, IVB)

Early asteroidal Fe–Ni core

Early Solar System

Walker et al. (2008); McCoy et al. (2011)

Pallasites

Differentiation: core–mantle material?

Early Solar System

Shen et al. (1998); Lee et al. (2006)

Mesosiderites

Differentiated metal-rich materials

Early Solar System

Shen et al. (1998)

Iron and stony-iron meteorites

Stony-iron meteorites

Partially melted achondrites AcapulcoiteLodranites

-

Partially melted rocks (~20%)

4563 ± 2 Ma

Ureilites

-

Melt residues

Early Solar System

Rankenburg et al. (2007; 2008)

Brachinites

-

Melt residue after 800 Ma), as for abyssal peridotites and most other mantle rocks. The suprachondritic Os compositions, as with those from the Oman ophiolite described earlier (Hanghøj et al. 2010) and many other ophiolites (see Figs. 14 and 15), require the addition of a radiogenic melt component (unless samples have experienced significant recent Re loss), likely during the formation of the Troodos around 90 Ma ago. The ultimate source of this radiogenic Os is not known, and could relate to seawater contamination prior to concentration in chromitites (because a radiogenic signature is also evident in the most Osrich chromitite samples) or to crustal contamination during emplacement, but the former at least is difficult to reconcile with the very low Os concentrations in seawater (Levasseur et al. 1998). Another possible mechanism, that would be applicable to both mid-ocean ridge and supra-subduction ophiolites, is the production of radiogenic melts due to preferential sampling of radiogenic interstitial sulfides (Alard et al. 2005; Harvey et al. 2011) or due to the presence of enriched domains in the mantle (cf. pyroxenites; Reisberg et al. 1991; Pearson and Nowell 2003). However, melting of enriched domains is not consistent with the refractory boninitic melt that typically produces HSE- and Cr-rich chromitites. Given the apparent global distinction in Os isotopes between ophiolites of convergent and mid-ocean ridge origin (Fig. 15), the most plausible explanation for a significant part of the radiogenic signature is a flux from the subducting slab, with Os mobilized in oxidized chlorine-rich fluids (Brandon et al. 1996; Becker et al. 2004). In this scenario, despite the extreme fractionation of Re from Os in oceanic crust, the low Os contents and relatively young age of subducted mafic crust would suggest that a sedimentary input may be required to provide sufficient radiogenic Os to impart that signature on the Os-rich mantle. The process(es) of dunite formation also induces significant HSE fractionation. Harzburgites, which could be simple residues of melting or, as Büchl et al. (2002) conclude, the product of melting during melt-percolation at low melt/rock ratios, have largely uniform IPGE patterns and concentrations that only range by roughly a factor of two (Fig. 16). Palladium and Re abundances do, however, vary over approximately an order of magnitude in harzburgites (Büchl et al. 2002). In contrast, a dunite rim and core, the product of high melt/ rock ratios, together with a websterite and a boninite all display high and remarkably uniform concentrations of Pt (6.5–12.2 ng/g), moderately variable Pd and Re, and two or more orders of magnitude variation in Os content. Qualitatively, it seems that dunites and websterite could be produced by some sort of reaction and mixing process between harzburgite and boninitic melt, retaining high Pt but removing/diluting Os; requiring Os to be mobilized. This is supported by modelling of HSE ratios (dominated by mixing of harzburgitic and magmatic sulfides) and REE in clinopyroxene during open system melting (Büchl et al. 2002). Shetland Ophiolite Complex (UK). Harzburgites from Unst, Shetland, have Os isotope compositions ranging from gOs of 2 to −6 (using an O-chondrite reference frame; 187 Re/188Os = 0.422, 187Os/188Os = 0.1283). Most Os isotope ratios are consistent with an ambient convecting mantle signature (see Os isotopic heterogeneity in the mantle in Discussion) but there is evidence of both melt depletion at ~ 1.2 Ga and also radiogenic Os addition for some samples (O’Driscoll et al. 2012). Dunites have a wider range of 187Os/188Os than harzburgites (gOs492Ma of −22 to 12), reflecting the effects of melt–rock reaction involved in their formation (O’Driscoll et al. 2012). Chromitites have the narrowest range of 187Os/188Os, from gOs +0 to +3.5. This relative homogeneity is perhaps surprising given the higher melt/rock ratios involved in producing chromitite, but this is

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Figure 14. 187Os/188Os–Al2O3 diagram for ophiolite ultramafic rocks (predominantly harzburgites, but also dunites), high-temperature convergent tectonites and sub-arc mantle xenoliths (see legend for symbols). Also shown for comparison are abyssal peridotites (diamonds), ocean island basalt mantle xenoliths (light grey circles), continent-ocean transitional tectonites (white squares) and sub-continental lithospheric xenoliths (mid-grey circles). There is considerable scatter in all datasets, partly reflecting variable ages of melt depletion, but also probably recent resetting of 187Os/188Os by seawater or melt interaction. The most Al-depleted ophiolitic samples (particularly those from convergent margin settings) and subduction zonerelated ultramafics have more radiogenic 187Os isotope compositions than peridotites from other settings. This presumably reflects a flux of radiogenic Os, or possibly a time-integrated addition of Re, related to the flux from the subducting slab, although greater melt–rock ratios in this environment may also play a part. Crustal contamination during emplacement is also possible. In part, the decoupling of 187Os/188Os from Al2O3 is due to the formation of dunitic rocks by melt–rock reaction, but many peridotites in convergent settings also possess more radiogenic Os than expected for a given Al content. Data sources for ophiolites: Snow et al. (2000); Kepezhinskas and Defant (2001); Büchl et al. (2004); Becker et al. (2006); Schulte et al. (2009); Hanghøj et al. (2010); Aldanmaz et al. (2012); O’Driscoll et al. (2012). High-T convergent margin tectonites: Reisberg et al. (1991); Roy-Barman et al. (1996); Becker et al. (2001, 2006); Pearson et al. (2004); Marchesi et al. (2014). Sub-arc xenoliths: Brandon et al. (1996); McInnes et al. (1999); Widom et al. (2003). Abyssal peridotites—see Figs. 1 and 2. Ocean island basalt mantle xenoliths—see Fig. 2. Continent/continent-ocean transition tectonite: Reisberg and Lorand (1995); Meisel et al. (1996); Roy-Barman et al. (1996); Rehkämper et al. (1999); Snow et al. (2000); Saal et al. (2001); Becker et al. (2006); Luguet et al. (2007); van Acken et al. (2008); Riches and Rogers (2011); Wang et al. (2013).

set against the extremely high Os concentrations, and low Re abundances, that allow for accurate estimation of the initial Os isotope composition. In part, the range for dunites (and harzburgites) may reflect difficulties in age correcting over 492 Ma (as this is dependent on measured Re and Os concentrations—with the potential for recent disturbance). Overall, however, a radiogenic Os flux is required to explain the supra-chondritic gOs values. As discussed for the Troodos Ophiolite, there are various possible sources of the radiogenic Os, but a flux from the downgoing slab may be the most plausible mechanism.

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Data for ophiolites of MORor uncertain origin; n = 88.

All samples

Convergent margin ophiolites All samples; n = 72.

Relative probability

Peridotites only; n = 67 (dashed)

Peridotites only; n = 41

-12

-10

-8

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-4

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0

γOs

2

4

6

8

10

12

14

Figure 15. Probability density plot of gOsinitial in ophiolitic ultramafic rocks (peridotites, dunites, and chromitites), grouped according to geological setting of formation. gOsi = (187Os/188Ossample initial/187Os/188Oschondrite initial −1) × 100. The absolute values in this plot should be treated with caution as these include a correction for ingrowth based on the measured 187Re/188Os, which, in some cases, may have been perturbed during/ since emplacement. In addition, this plot is based on a limited number of different ophiolites, with several ophiolites contributing a disproportionate number of data: Troodos (Cyprus), Samail (Oman), Shetland (UK), Taitao (Chile) and Jormua (Finland) ophiolites account for 129 of the 160 analyses. Given the different Os isotope records preserved by PGM grains from different ophiolites (Pearson et al. 2007), much of this difference could merely reflect large-scale mantle heterogeneity. Nonetheless, the overall offset between the two categories is two to four gamma units, which may represent a real difference generated by addition of radiogenic Os in the supra-subduction zone environment. Data sources given in Figure 14, except Becker et al. (2006) and Dijkstra et al. (2010).

Shetland Ophiolite samples display huge variations in HSE concentrations, with some chromitites containing up to ~ 100 µg/g Pt (Prichard and Lord 1996; O’Driscoll et al. 2012) while some dunites contain less than 100 pg/g Pt. The most HSE-rich chromitites (from Cliff) have Ir and Ru contents that are roughly two orders of magnitude higher than the range of chromitites analyzed from the Qalander, Luobusa, and Zambales ophiolites (Fig. 16). Moreover, these chromitites have unusual HSE patterns with PPGE/IPGE ratios greater than unity and Pd concentrations up to 156 µg/g (O’Driscoll et al. 2012), compared with typical IPGE-rich chromitites which have Pd and Pt contents approximately four orders of magnitude lower (Zhou et al. 1996, 2000, 2014; Ismail et al. 2014). The range of HSE abundances between chromitites from different localities is, in itself, huge. The two other localities analyzed have more typical HSE patterns, albeit in one case also enriched by one to two orders of magnitude. The degree of P-PGE enrichment has been linked to the thickness and sulfide content of the ultramafic dunite sequence and ultimately to the degree of melting, and, in the case of the extremely PPGE-enriched Cliff chromitites, also linked to hydrothermal redistribution from surrounding ultramafics (Prichard and Lord 1996). There are also large variations in the HSE concentrations and patterns of dunites, which show an overall depletion in Pt, relative to IPGE, and are enriched in Pd in many cases. Rhenium concentrations are low in almost all harzburgite, dunites, and chromitites, although enrichment in Re does also occur in some dunites.

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Chondrite-normalised conc.

0.1

- Harzburgites, dunites & lherzolite

Luobusa Ophiolite (Tibet) PM

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Harzburgite Lherzolite Dunite (For all panels)

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Shetland Ophiolite (UK)

- Harzburgites & dunites PM

0.01 PM

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Qalander (Iraq) & Zambales (Philippines) Ophiolites Qalander - Harzburgites & dunites Zambales

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Chondrite-normalised conc.

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PM

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Chondrite-normalised conc.

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Qalander Luobusa Zambales

Harold’s Grave 1

0.1 The Viels 0.01

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PM

Chromitites (& one dunite) from Shetland Ophiolite (UK)

Os

Ir

Ru

Rh

Pt

Pd

Re

Os

Ir

Ru

Rh

Pt

Figure 16. Chondrite-normalized concentration diagrams of the HSE in ophiolites of convergent margin (Troodos, Shetland, Zambales) or uncertain origin. PM estimate shown for comparison. See Fig. 3 for normalization values. White squares for Troodos denote Re–Os analyses of dunites. Qalander and Zambales chromitites: black lines—Cr-rich, grey lines—Al-rich; Luobusa chromitites: black—massive, grey—disseminated. It is not clear why there is a discrepancy in the Os data for Luobusa, across two studies. Given that Becker et al. (2006) used high-temperature acid digestion and isotope dilution, these Os data should be used in the first instance; the other HSE data is broadly comparable between the two studies. (References: Zhou et al. 1996, 2000; Büchl et al. 2002, 2004; O’Driscoll et al. 2012; Ismail et al. 2014).

Zambales Ophiolite (Philippines). The Zambales Ophiolite contains two distinct blocks, which differ in the composition of their chromitites. The Acoje Block contains chromitites with high-Cr spinel, while the Coto Block is characterized by more Al-rich spinel (Zhou et al. 2000). A comparative study of these two blocks found variations and similarities in the HSE budget of the two chromitite types. As in other studies (e.g., Ahmed et al. 2006; Ismail et al. 2014) high-Cr chromitites are found to be richer in HSE than those with high-Al spinel. In this case, however, the IPGE contents vary significantly (e.g., Ru = 8–38 ng/g for Coto, and 62–70 ng/g

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for Acoje), while Pt and Pd contents and ratios are similar in the two types (Fig. 16) (Zhou et al. 2000). Dunites are also found to vary, particularly in Pt content, with the Acoje Block having more Pt-rich dunites. These spinel compositions and HSE contents are linked to the parental magmas of the chromitites. The Cr-rich Acoje chromitites were likely generated by interaction with a refractory boninitic melt, while the Coto chromitites probably had a more tholeiitic source. Boninitic melts are typically sulfide undersaturated, and thus may form with, and retain, high HSE abundances, compared to tholeiitic melts which are commonly saturated in sulfide thus inducing its precipitation and a reduction in HSE content of the remaining melt (Zhou et al. 2014). Qalander Ophiolite (Iraq). The Qalander Ophiolite is a poorly preserved mélangetype complex, containing serpentinized dunites and harzburgites which surround two types of podiform chromitite; high-Al and high-Cr. The harzburgites and dunites analyzed have comparable HSE patterns overall (Ismail et al. 2014), broadly similar to PM estimates (Becker et al. 2006), except offset to higher concentrations (Fig. 16) particularly for Os (4–9 ng/g Ir, 10–17 ng/g Os). As with other chromitite occurrences, Cr-rich and Al-rich types have differing relative proportions of HSE, although they almost all possess high IPGE/PPGE ratios (see Zhou et al. 2014; cf. Shetland, above). Cr-rich chromitites are the most strongly enriched in IPGE, and have the highest IPGE/PPGE ratios. Al-rich chromitites have significantly higher PPGE concentrations, above those of peridotite, while the Cr-rich type has PPGE at the low end of the peridotite range. Egyptian ophiolites and podiform chromitites, Oman N massifs. The Os isotope composition of PGM from chromitites of the Proterozoic Eastern Desert ophiolite, Egypt and in the Phanerozoic Oman ophiolite were analyzed by Ahmed et al. (2006). It was found that PGM from different regions of each ophiolite have distinct 187Os/188Os ratios, from sub- to broadly chondritic in some regions, to significantly suprachondritic in others (0.1293 for the Proterozoic Eastern Desert ophiolite and up to 0.1459 for the Oman ophiolite). At the same time, there are also distinct compositions of the chromitites themselves, with (i) concordant lensoid forms with intermediate-Cr spinel, which are relatively PGE-poor, and (ii) discordant, dyke-like chromitites, with high Cr spinel, which are PGE-rich. The authors conclude that the variety of chromitites, and the Os–HSE signatures that they contain, reflects the variety of formation processes. The radiogenic chromitites of the Eastern Desert are thought to be affected by crustal contamination, whereas the radiogenic, Cr- and HSE-rich chromitites from Oman reflect high degree melting and an input from a subducting slab, most likely in a suprasubduction zone setting (Ahmed et al. 2006), although here we note that some workers prefer a MOR origin and obducted emplacement for the Oman ophiolite (see earlier section). Feather River ophiolite (California). A suite of serpentinized peridotites from the Feather River ophiolite has been compared with serpentinized abyssal peridotites and used as a means of establishing the chemical impacts of serpentinization at a range of water/rock ratios and depths in the mantle (Agranier et al. 2007). The serpentinites have elevated concentrations of seawater-derived fluid mobile elements, such as boron, although typically lower than abyssal peridotites. Feather River serpentinites do not have corresponding seawater-affected suprachondritic 187Os/188Os ratios (measured range: 0.1175–0.1279). Nonetheless, there is a probable covariation between Os abundance and Os isotope composition in Feather River rocks, which may reflect incorporation of seawater-derived radiogenic 187Os/188Os. Agranier et al. (2007) contend that the serpentinites formed at lower water/rock ratios (greater depth) than is typical for abyssal rocks, and are therefore more representative of bulk serpentinized lithosphere. In summary, melt percolation in the supra-subduction zone environment generates substantial lithological heterogeneity, which is accompanied by significant Os isotope and HSE variability, both between lithological groups (harzburgites, dunites, chromitites, pyroxenites) and within groups. There is compelling evidence for addition of melt-derived radiogenic 187Os to parts of the mantle sections of ophiolites (see above and Figs. 14 and

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15), most probably due to a degree of Os fluxing from the downgoing slab, although other possibilities exist. However, the precise mechanism for such a transfer is not yet clear. The process of melt–rock reaction during melt percolation results in a decoupling of Al2O3 and 187Os/188Os (Fig. 14), which for other suites is considered a fairly robust method for determining the approximate ages of depletion for suites of peridotites, where measured Re contents are often unreliable (Meisel et al. 2001; Lassiter et al. 2014).

Highly siderophile elements in the mantle sections of ophiolites of uncertain origin Luobusa ophiolite (Tibet). Chondrite-normalized HSE concentrations for harzburgites, dunites and chromitites from the Luobusa ophiolite are presented in Fig.16. The concentrations of Ir, Pt and Pd are broadly comparable between two different studies (Becker et al. 2006; Zhou et al. 1996), but the low Os/Ir ratios of the Ni-S fire assay data of Zhou et al. (1996) are not supported by the high temperature (345 °C) isotope dilution data of Becker et al. (2006), suggesting either different petrogenetic histories for the two sample sets or an unidentified analytical issue for Os in the Zhou et al. data. To err on the side of caution, we will assume the latter here and disregard the very low Os/Ir ratios in the harzburgites and chromitites. The harzburgites appear to represent residua after MORB extraction (Zhou et al. 1996). The HSE abundances are similar to the PM mantle estimate (Becker et al. 2006), and do not indicate significant melt depletion, except perhaps for Pt (although data for Re—the most incompatible HSE—is only available for two samples). The Cr-numbers of Cr-spinel in melt-reacted dunitic rocks are higher than those in the harzburgites, suggesting interaction of a boninitic melt with the residual peridotite, which also removed pyroxene (Zhou et al. 1996). As a result, melts became more boninitic and saturated in Cr-spinel, which precipitated to form chromitite pods within the dunite zones. The inferred boninitic melts suggest a subduction-related origin for this ophiolite. Chromitites have distinct, strongly fractionated HSE patterns with high IPGE/PPGE ratios (e.g., normalized Ir/Pt ratios ~ 100). The concentrations of IPGE in the chromitites are an order of magnitude or more greater than those of the harzburgites, while Pt abundances are approximately five times lower in the chromitites than the harzburgites, and are comparable to the dunites (Fig. 16). These concentrations and patterns are similar to other Cr-rich chromitites from the Qalander and Zambales ophiolites (Zhou et al. 2000; Ismail et al. 2014). Dunites have similar PPGE contents to the chromitites, but without the enrichment in IPGE, due, presumably, to a lack of PGE saturation, and consequent PGM formation, during dunite formation. Jormua ophiolite (Finland). Serpentinites, the oxides they contain, and podiform chromitites have all been analyzed for Re–Os abundances and Os isotopes (Tsuru et al. 2000). As with most abyssal peridotites that have undergone serpentinization, Os concentrations, although somewhat variable (1.5–11.7 ng/g) are broadly similar to those of the convecting upper mantle. Rhenium abundances are more variable; most samples are depleted in comparison with PM (Becker et al. 2006) but some experienced (probably recent) Re enrichment. Wholerock samples have experienced open-system behavior, with respect to Re–Os isotopes, but chromite to Cr-rich magnetite separates have extremely low Re/Os and largely homogenous initial 187Os/188Os values, with a mean initial gOs of approximately −5, suggesting closedsystem behavior. Other parts of the ophiolite contain chromitites with initial gOs between +1 and +3. Tsuru et al. (2000) concluded that the positive values may indicate the presence of MORB-type and subcontinental lithospheric mantle sources. Addition of radiogenic Os by melt percolation may be another mechanism to explain the Os isotope data. Outokumpu ophiolite (Finland). The Cr-rich nature of residual chromites and boninitelike volcanic rocks suggest a supra-subduction origin for this ophiolite, but an origin in a continental rift zone has also been proposed (Walker et al. 1996). The key conclusion of an Os isotope study (Walker et al. 1996), mainly of chromite, was that this mantle section displayed

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broadly chondritic 187Os/188Os ratios, and hence Re and Os abundances, which were used to support the ‘late veneer’ model (Chou 1978). In detail, however, there were variations from a ‘residual’ sub-chondritic laurite (Ru(Os,Ir)S2) to fluid addition with a composition of around 0.4 gOs. In this case, however, the radiogenic signature is thought to be derived either from seawater contamination or from a crustal input during emplacement, akin to that proposed for the Eastern Desert Ophiolite, Egypt (see previous section). Tethyan ophiolites (Turkey). Harzburgites and dunites from Tethyan ophiolites at Koycegiz, Marmaris, Tekirova, Adrasan and Lake Salda in Turkey have been analyzed by Aldanmaz et al. (2012). Both mid-ocean ridge and supra-subduction zone geochemical signatures have been identified in different parts of the ophiolites, and these have differing HSE systematics. The mid-ocean ridge harzburgites have broadly chondritic Os/Ir and supra-chondritic Pd/Ir and Rh/Ir, similar to PM estimates (Becker et al. 2006), although some PPGE/IPGE enrichment is ascribed to sulfide addition. They also have a sub-chondritic range of measured 187Os/188Os of 0.1223–0.1254, and have correspondingly depleted Re/Os ratios (Aldanmaz et al. 2012). In contrast, the peridotites of supra-subduction zone affinity have more variable HSE patterns and a wider range of 187Os/188Os from 0.1209 to 0.1318, which is −5.3 to 3.3 in gOs90Ma units, relative to O-chondrite evolution. The greater heterogeneity of supra-subduction zone peridotites, compared to those of mid-ocean ridge affinity, reflects a more complex evolution. Eastern Alps ophiolites (Austria). Peridotitic units of Eastern Alps ophiolites (the Reckner, Hochgrossen, Kraubath, Steinbach and Bernstein peridotites; including two chromitites) have been found to have remarkably uniform 187Os/188Os ratios (~ 0.1266–0.1281), clustering around the chondritic evolution curve (Meisel et al. 1997), with the exception of one locality (Dorfertal) which has an Os isotope composition consistent with a minimum age of Re depletion of ~ 1.6 Ga. The authors considered the uniformity of Os composition to be somewhat surprising given the uncertain age and affinity of the samples. One important finding of that study was the robustness of Os isotopes, given a high degree of serpentinization, compared with other geochemical data, and even petrographic and field methods. Mayari-Cristal ophiolite (Cuba). The key finding of a study of PGM in the MayariCristal ophiolite was the scale of Os isotope heterogeneity present within single hand specimens, thin sections and down to a scale of several millimeters that separated two PGM with contrasting Os isotope ratios (187Os/188Os: 0.1185 and 0.1232; Marchesi et al. 2011), which equate to Re depletion ages of 1370 and 720 Ma, respectively (O-chondrite reference). Given that the budget of Os for these PGM is thought to be sourced from several m3 of mantle, this has intriguing implications for mixing (or the lack thereof) of distinct percolating melts in the mantle (Marchesi et al. 2011).

DISCUSSION Influence of low-temperature alteration processes on the HSE in bulk rocks and minerals Here we briefly discuss the influence of low-temperature (non-magmatic) processes on the bulk rock, sulfide, and PGM composition of mantle tectonites. Ultrabasic rocks affected by oxidative weathering are usually not used for bulk rock chemical analyses to study hightemperature processes. Sulfides are at least partially oxidized by these processes, thus, it is expected that the abundances of chalcophile elements will be disturbed in non-systematic ways. Because areas of ultramafic rocks affected by oxidative weathering are easily identified by their brown color, stemming from ferric iron bearing secondary weathering products, such altered areas can be normally identified and removed.

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The influence of serpentinization on HSE abundances and 187Os/188Os. Serpentinization represents another common low temperature alteration process of ultrabasic rocks. Serpentinization reactions occur during the reaction of igneous and metamorphic ultrabasic rocks with seawater or freshwater under a range of geologic conditions and temperatures (e.g., Evans et al. 2013 and references therein). For instance, these processes occur today in oceanic mantle exposed on the seafloor and at greater depth where heated seawater moves within deep-reaching fractures. Similar processes occurred in ultramafic parts of ophiolites during their exhumation on or beneath past seafloors, during tectonic obduction or by reaction with fluids and meteoric water of variable origin during continental collision (Hirth and Guillot 2013). During serpentinization of peridotites, water reacts with olivine, pyroxenes, spinel (to a lesser extent) and sulfides that formed at high temperatures. Depending on temperature and progress of reaction, the new minerals formed include serpentine minerals (chrysotile, lizardite, at higher temperatures antigorite), magnetite and other secondary minerals such as brucite (see for example Bach et al. 2004). The influence of serpentinization on the abundances of HSE in mantle tectonites has not been studied in much detail. Early Re–Os studies of serpentinized peridotites (e.g., Snow and Reisberg 1995) have emphasized that serpentinization of peridotites in the oceanic lithosphere occurs under reducing conditions. Because of the low fO2 environment caused by the local production of hydrogen and methane (Evans et al. 2013), secondary sulfides (heazlewoodite, millerite, godlevskite), Fe–Ni alloy phases (awaruite) and native metals (Au, Cu) may form (Klein and Bach 2009) and thus, the HSE are able to retain a low valence. The extent to which the HSE are retained in these secondary phases compared to the original abundances in the unaltered bulk rocks and how much of the HSE may be lost into the fluids is poorly constrained. The similarities of abundances of Os, Ir, Ru, Rh, Pt, and Pd in fresh and variably serpentinized peridotites with similar lithophile element composition have been used to argue that serpentinization at reducing conditions results in only minor changes in the abundances of these elements in serpentinized ultramafic bulk rocks that are difficult to resolve from analytical or intrinsic variations in such rocks (e.g., Becker et al. 2006; van Acken et al. 2008; Fischer-Gödde et al. 2011; Marchesi et al. 2013; Foustoukos et al. 2015). This contention is supported by abundances of these elements in serpentinized komatiites, which often preserve correlations between PGE and Mg or Ni, which were unequivocally produced by igneous fractionation processes (e.g., Brügmann et al. 1987; Puchtel et al. 2004, 2005). The influence of serpentinization on Re and Au abundances is more difficult to predict, as no systematic studies exist and the applicability of experimental studies of Re behavior in specific hydrothermal fluids is difficult to evaluate (Xiong and Wood 1999; Pokrovski et al. 2014). Compared to Pd, Re is often depleted in serpentinized harzburgites, as expected for strongly depleted residues of partial melting; however, it may also be more enriched than Pd in normalized concentration diagrams (e.g., Figs. 3, 5, 9, 10, 16). It is difficult to judge if these abundances reflect secondary addition of Re from seawater (which has very low Re abundances) that has dissolved sulfides elsewhere, or, if re-enrichment of Re occurred before alteration (e.g., by precipitation of liquid sulfide from silicate melts, as may be plausible from observations of unaltered peridotites). Similar uncertainties arise in serpentinized lherzolites. Correlations of Re with indicators of melt extraction or refertilization such as Al, Ca or Mg/(Mg + Fe2+) in peridotites have been interpreted as evidence for limited mobilization of Re by low-temperature alteration processes (e.g., Becker et al. 2006). In mantle pyroxenites that were affected by variable degrees of serpentinization, Re seems to be unaffected by alteration because it is typically systematically more enriched than Pd and Pt. Such a behavior is expected from crystal fractionation products of basic melts (van Acken et al. 2010b). The behavior of gold during serpentinization of mantle peridotites has not been studied systematically either. Although Au, in some cases, follows Pd and Re in its geochemical behavior in unaltered

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peridotites (Fischer-Gödde et al. 2011), it shows scattered distributions in element variation diagrams that are not well understood. Because of the known mobility of Au in hydrothermal systems in basic and ultrabasic rocks (Pokrovski et al. 2014) and the enrichment of Au in some serpentinite-hosted sulfide deposits (e.g., the Lost City hydrothermal field, Mid Atlantic Ridge), it is to be expected that Au may be rather mobile during serpentinization. The question of whether or not the Os budget of serpentinized peridotites can be measurably affected by radiogenic 187Os from seawater has been discussed in several publications (e.g., Martin 1991; Roy-Barman and Allègre 1994; Snow and Reisberg 1995; Brandon et al. 2000; Standish et al. 2002; Alard et al. 2005; Harvey et al. 2006). Cenozoic seawater has highly variable and mostly very radiogenic 187Os/188Os ranging over 0.5–1 (Peucker-Ehrenbrink and Ravizza 2000), however, the concentration of Os in seawater is extremely low (about 3.8 fg/g Os, (Sharma et al. 1997). These low abundances are in stark contrast to the ng/g levels of Os in peridotites. Figure 17 illustrates the effects of simple peridotite–seawater mixing, assuming 187Os/188Os of 0.122 and 0.127 and 3.9 ng/g Os in unaltered peridotite and modern seawater with 187Os/188Os of 1 and 3.8 fg/g Os. Very high water–rock ratios of 103–104 are required in order to disturb the 187Os/188Os of peridotite bulk rocks at the % level or higher. Lower values of 187Os/188Os in seawater, such as 0.5, would not alter this conclusion. For comparison, water–rock ratios of significantly less than 100 have been calculated for rock units of the Oman ophiolite (McCulloch et al. 1981). Some workers have suspected that Mn hydroxide films in cracks and on surfaces may pose a problem because these phases tend to scavenge Os from seawater (Martin 1991; Roy-Barman and Allègre 1994). Although leaching studies of serpentinized peridotites have not yielded clear indications of contamination, it is preferable to remove such surfaces or avoid such rocks altogether. Many abyssal peridotites are strongly serpentinized, yet they are characterized by chondritic to subchondritic 187 Os/188Os, similar to unaltered or weakly serpentinized post-Archean peridotite xenoliths or other tectonites. Thus there appears to be no need to invoke late addition of radiogenic Os by serpentinization. Positive linear correlations of 187Os/188Os with Al2O3 contents in serpentinized peridotites provide the strongest argument against a significant influence of serpentinization on 187Os/188Os in such rocks (Reisberg and Lorand 1995). These correlations are a primary magmatic feature of mantle rocks (e.g., Handler et al. 1997; Peslier et al. 2000; Meisel et al. 2001; Gao et al. 2002). Suprachondritic 187Os/188Os occasionally occur in bulk rocks of strongly serpentinized abyssal peridotites (Standish et al. 2002) and from serpentinized harzburgites and dunites of ophiolite sections and peridotite massifs (e.g., Becker et al. 2001; Büchl et al. 2002;

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At high degrees of melting, partitioning of Os, Ir, Ru, Rh, and Pt may be controlled by the solubility of alloys of these elements in silicate melt (Pearson et al. 2004; Fonseca et al. 2011, 2012; Mungall and Brenan 2014; Brenan et al. 2016, this volume). The significance of this for the composition of harzburgites will be discussed later. Here, we specifically focus on processes during low and moderate degrees of melting in the deeper regions of the melting column where sulfide should be stable in the residue and sulfide–silicate partitioning has been proposed as the main control on the distribution of the HSE and other chalcophile elements (Barnes et al. 1985; Morgan 1986). However, it has been unclear if sulfide exists as a solid phase (mss), liquid sulfide, or both. Recent improvements in the accuracy and precision of liquid sulfide–silicate partition coefficients (Dsulf/sil) indicate values in the range of 105–106 and 104 for the PGE and Au, respectively (Li and Audétat 2013; Mungall and Brenan 2014; Brenan et al. 2016, this volume), whereas Re is much less chalcophile (Dsulf/sil ≈ 300–800, Fonseca et al. 2007; Brenan 2008). Assuming a simple fractional melting process (batch melting yields similar results as long as the elements are not highly incompatible), element concentrations in the residues can be calculated according to the mass balance equation Cr = Co b (1-F)((1/D )−1), with Cr = concentration of an element in the residue, Co = total concentration of an element in the bulk system (residue + melt), Db = bulk partition coefficient of an element between residue and melt, F = melt fraction. As long as sulfide is present in the mantle residue and it is equilibrated with silicates and silicate melt, the high Dsulf/sil require nearly constant concentrations of all PGE in peridotites (Fig. 18a), because bulk partition coefficients of the PGE in lherzolites are ≫ 1: At 0.02 wt% S in fertile lherzolite and 35 wt% S in monosulfide solid solution, DPGEb > 0.00057 × 105 + 0.9994 × 0.1 = 57, assuming DPdsil.min./sil.melt < 0.1 with other PGE likely having higher Dsil.min./sil.melt (Mungall and Brenan 2014). Gold should also be retained in lherzolites that have lost a significant fraction of melt (DAub ≥ 0.00057 ×5 ×103 = 3, assuming DAusil.min./sil.melt < 0.01 (Mungall and Brenan 2014), whereas Re should be moderately depleted at relevant fO2 in normal upper mantle (FMQ − 1), as its Db is always below 1 in cases where no garnet occurs in the residue (DReb ≤ 0.00057 ×800 + 0.9997 ×0.1 = 0.6, assuming DResil.min./sil.melt < 0.1 (no garnet), Mallmann and O’Neill 2007; Brenan 2008). The situation in mantle rocks, however, has been found to be more complicated; one indication being the difficulty in reproducing peridotite HSE patterns by sulfide–silicate equilibrium partitioning (Fig. 18a). In the following, we discuss evidence suggesting that many mantle peridotites are in chemical disequilibrium regarding chalcophile element partitioning at the scale of hand specimen to grain boundaries. An alternative partitioning scenario, such as mss–sulfide liquid–silicate liquid equilibrium, is also discussed below. Melt infiltration induces chemical disequilibrium of chalcophile elements in mantle peridotites. Studies of chalcophile element abundances in sulfides of different textural position, in mantle xenoliths and in peridotite tectonites, have shown that significant compositional differences may exist between sulfides that occur as inclusions in olivine (and sometimes pyroxenes and spinel) and sulfides present at grain boundaries. The former are rich in Ir-group PGE and depleted in Pd, Au, and Re, while the latter may or may not be depleted in IPGE and have higher Pd, Re, and Cu (Alard et al. 2000, 2002; Luguet et al. 2001, 2003, 2004). Although these different assemblages are sometimes complicated by internal separation into multi-phase assemblages (pentlandite, pyrrhotite, and other phases) that occurred late during slow cooling, it is clear from their different compositions that included and grain boundary sulfides were not chemically equilibrated during their formation. The sulfide assemblages on grain boundaries are sometimes associated with pyroxene– spinel assemblages that have been interpreted to have formed during melt infiltration and refertilization. From this observation, it follows that reactive melt infiltration likely led to sulfur saturation in these magmas and precipitation of the sulfides located on grain boundaries (e.g., Alard et al. 2000). The reaction of silicate melts and sulfide segregation

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processes are not only indicated by the different sulfide assemblages in the peridotites, but also by the HSE abundances in mineralogically zoned boundaries between pyroxenites and host peridotites and disequilibrium sulfide assemblages in mantle pyroxenites (see sections on mantle pyroxenites below). Some authors have proposed that sulfide melts may be mobile in mantle rocks, and thus may change the Re–Os and PGE systematics of mantle rocks (Gaetani and Grove 1999). The existing data on peridotites, however, do not support pervasive or widespread sulfide melt mobility, as linear correlations between 187Os/188Os, Re, and S abundances and lithophile elements such as Al, Ca, or Mg in peridotites would not be maintained over long periods of time in the mantle (Figs. 4, 7; e.g., Reisberg and Lorand 1995; Meisel et al. 2001; Becker et al. 2006; Wang and Becker 2013), although minor mobility is not precluded due to scatter in the datasets. The role of fluids as metasomatic agents in the redistribution of HSE and other chalcophile elements has been invoked in some cases (e.g., Lorand and Alard 2010). One possibility is that such fluids are the end products left after crystallization of mantle-derived melts or, if they are of external origin, may have been derived from crustal sources at lower temperatures during the exhumation history of mantle tectonites. Regardless of the origin of the fluids, what is not yet clear is the effect of these small-scale observations on the mass balance of bulk rocks. In summary, silicate melts are the main metasomatic agents that, by way of coupled precipitation of sulfide melt, pyroxenes and an Al phase, clearly produce significant modifications of HSE abundances and 187Os/188Os at magmatic temperatures in the mantle. Detailed surveys of the accessory mineral inventory of peridotites (e.g., Fig. 6) have revealed the occurrence of Pt–Ir alloys, Ru–Os-bearing sulfides and Os–Ir–Ru alloy phases (Luguet et al. 2007; Lorand et al. 2010; O’Driscoll and González-Jiménez 2016, this volume). These phases are expected to become stabilized by decreasing fS2 shortly before or during the exhaustion of liquid sulfide in harzburgite residues at moderate to high degrees of melting (e.g., Fonseca et al. 2012; Mungall and Brenan 2014; Brenan et al. 2016, this volume). Thus, their occurrence in harzburgites (e.g., at Lherz; Luguet et al. 2007) is not unexpected. However, such phases have also been detected in lherzolites from Lherz that formed by refertilization, albeit they occur in smaller proportions than in harzburgites (Lorand et al. 2010). If the alloy phases were indeed inherited from more depleted parent rocks, their presence in some lherzolites may also reflect chemical disequilibrium between these phases and the more abundant sulfide minerals that were precipitated as sulfide liquid from silicate melt. The impact of such inherited and presumably ‘residual’ alloy phases on bulk rock budgets of lherzolites that formed by refertilization appears rather limited. For instance, the bulk rock Os/Ir ratios of lherzolite tectonites is rather homogeneous and overlaps chondritic values (Fig. 19a, Pearson et al. 2004; Becker et al. 2006; Fischer-Gödde et al. 2011; Liu et al. 2009; Wang et al. 2013). Because of the different solubilities of Os and Ir metal in silicate melt (e.g., Mungall and Brenan 2014), chondritic Os/Ir are not a priori maintained in residual peridotites at higher degrees of melting (as witnessed by the larger scatter of this ratio in harzburgites). Pt/Ir and Pt/Os in lherzolites range from chondritic to mildly subchondritic. Only rarely do lherzolites display enrichments of Pt that are decoupled from Pd, Au, and Re (e.g., Fig. 5b, c) and might be ascribed to the excess presence of Pt minerals. In this context, it is noteworthy that ratios of Ir, Os, and Ru in mantle tectonites tend to be more scattered in harzburgites than in lherzolites (Fig. 19). The difference in homogeneity of the different rock types may either reflect digestion problems in the laboratory, i.e., the difficulty of complete dissolution of refractory platinum group metal alloys in harzburgites (Meisel and Horan 2016, this volume, and references therein), or it may be due to dissolution of refractory alloy phases in coexisting sulfur-undersaturated melt at high temperatures. Osmium isotopic disequilibrium within mantle peridotites. Evidence for small-scale chemical disequilibrium regarding chalcophile elements is provided by Re–Os data that suggest

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Al2O3 wt. % Figure 19. a) Os/Ir–Al2O3 and b) Ru/Ir–Al2O3 in peridotite tectonites. Representative lherzolites and harzburgites from continental extensional and transitional oceanic environments (Balmuccia: solid circles, Baldissero: open circles, Lherz: x, Turon de la Tecuere: +, Lanzo: solid diamond, Internal Ligurides: open diamond). Also shown are harzburgites (solid squares) and dunites (open squares) from the Wadi Tayin section of the Oman ophiolite, dunites from Balmuccia (solid circles within the Dunite fields, see also Figs. 5, 8b) and lherzolites from Ronda (open triangle) and Beni Bousera (solid triangle). For data sources of peridotites see Figure 5 and text. Chondritic range from Horan et al. (2003) and Fischer-Gödde et al. (2010). Primitive-mantle model from Becker et al. (2006). The data show relatively homogeneous ratios in lherzolites and larger variations in harzburgites and in replacive dunites (see text for details).

that grain- to hand specimen-scale Os isotopic disequilibrium is common in the mantle. Burton et al. (1999) found that different mineral separate fractions from mantle xenoliths showed differing 187Os/188Os that were not related by isochronous behavior. Leaching experiments of powders of refertilized mantle xenoliths and tectonites show that 187Os/188Os frozen in during the Archean or Proterozoic survived Phanerozoic refertilization, most likely because of the preservation of ancient chromite or olivine that contained inclusions of HSE carrier phases (Chesley et al. 1999; Becker et al. 2006; Wang et al. 2013). Alard et al. (2002, 2005) showed that the sulfide populations with different PGE compositions also display systematic differences in Re/Os and 187Os/188Os. In peridotite xenoliths and abyssal peridotites, sulfides on grain

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boundaries tend to have chondritic to suprachondritic Re/Os and 187Os/188Os, whereas sulfides in inclusions also display subchondritic values (Harvey et al. 2006, 2011; Warren and Shirey 2012). The heterogeneous 187Os/188Os in different bulk rocks of essentially all suites of peridotites, xenoliths or tectonites from different geodynamic environments (e.g., Figs. 1, 2, 4, 7, 9, 11; and Luguet and Reisberg 2016, this volume) also represents a manifestation of disequilibrium on the scale of hand specimen and outcrops. In principle, such variation may have been caused by differences in the age of partial melting and melt infiltration. However, evidence for grainscale initial Os isotopic heterogeneity at times of melt infiltration (in cases where the timing can be constrained) suggests that mixing of residues and melts with different 187Os/188Os during reactive melt infiltration did not result in full Os isotopic equilibrium. A good example are the ultramafic tectonites in the Pyrenees and in the Italian and Swiss Alps (Baldissero, Balmuccia, Lanzo, Totalp), where episodic melt infiltration into Proterozoic continental lithospheric mantle during Paleozoic and Mesozoic extension only partially re-equilibrated 187Os/188Os values. All these data and observations suggest that disequilibrium must have been maintained even at high temperatures in the upper mantle and in the presence of silicate melt. The widespread heterogeneity of initial 187Os/188Os at the grain boundary- to centimeter-scale in mantle rocks also suggests that sulfide liquids are efficiently trapped even during recrystallization processes. Alongside evidence from textures and lithophile elements (e.g., Rivalenti et al. 1995; Müntener et al. 2005; Le Roux et al. 2007; Mazzucchelli et al. 2009), the extent of re-equilibration is manifested in the scatter of HSE abundances displayed by different suites of peridotites, in the abundance of harzburgite rocks in outcrops and in the distribution of Re–Os model ages in these bodies. At Lherz, Lanzo and Baldissero Re depletion ages of peridotites display bimodal distributions of Proterozoic and Phanerozoic ages, with harzburgites or depleted lherzolites typically showing older model ages (i.e., lower measured 187Os/188Os) than lherzolites (Reisberg and Lorand 1995; Burnham et al. 1998; Becker et al. 2006; Fischer-Gödde et al. 2011; Wang et al. 2013). In contrast, at Balmuccia and Totalp, depleted lherzolites and harzburgites are rare and display Proterozoic Re depletion ages. Model ages of fertile lherzolites at these locales range from Phanerozoic to future ages (van Acken et al. 2008, 2010a; Wang et al. 2013). Of note is that the scatter of the concentrations of Os, Ir, and Ru in fertile peridotites at these localities is more limited than in other lherzolite bearing tectonites (compare Fig. 5b with 5a and 5c). Osmium isotopic heterogeneity is also prevalent in abyssal peridotites, which are commonly presumed to represent melting residues of MORB-type magmas. Harvey et al. (2006) have shown that sulfides in harzburgites from the 15° 20’ N fracture zone (Atlantic Ocean) preserve small-scale isochronous relationships that date back to the Paleo-Proterozoic. Such preservation of early- to mid-Proterozoic 187Os/188Os values in bulk rocks and sulfides has also been reported in other abyssal peridotites (Parkinson et al. 1998; Alard et al. 2005; Liu et al. 2008; Warren and Shirey 2012). Further evidence of small-scale disequilibrium is apparent in studies of platinumgroup minerals from ophiolites. Platinum-group minerals from the Mayari-Cristal Ophiolite, Cuba, have been found to have diverse 187Os/188Os ratios even on the scale of a single thin section (Marchesi et al. 2011). The most extreme example found was the presence of two PGM only a few millimeters apart, with 187Os/188Os ratios of 0.1185 and 0.1232 (Marchesi et al. 2011), which give TRD ages of 1370 and 720 Ma, respectively (ordinary chondrite reference evolution line; Walker et al. 2002a). The mechanism of formation for such PGM is not well known, but given that the budget of Os for these PGM is thought to be sourced from at least several m3 of mantle (total Os equivalent to ~ 1 m3 mantle), this would imply little if any mixing of percolating melts, or a lack of equilibration between mineral grains and subsequent percolating melts. The influence of disequilibrium between mantle and magmas on HSE distributions. The predicted behavior of the HSE can be compared with HSE patterns of peridotites. Relatively ‘constant’ concentrations have been noted for the Ir group PGE in many studies of lherzolite tectonites. However, Rh and Pt display a tendency towards higher concentrations

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in lherzolites (e.g., Fischer-Gödde et al. 2011). In some (but not all) suites of peridotites, Pd correlates with fertility indicators such as Al2O3 abundances (e.g., Becker et al. 2006). Some workers (e.g., Lorand et al. 1999; Pearson et al. 2004) have noted that the variable depletion of Pd in lherzolites is difficult to reconcile with partial melting and very high sulfide–silicate partition coefficients (> 104 to 105). The smooth depletion of Pd, Au, and Re relative to other HSE in lherzolites from Balmuccia and elsewhere (e.g., Fig. 18) is inconsistent with equilibrium partitioning and the liquid sulfide-liquid silicate partitioning data. It is also difficult to explain by other equilibrium partitioning processes involving sulfides, e.g., mss–liquid sulfide (see below). Furthermore, concentrations of Os, Ir, and Ru in peridotite tectonites of similar lithophile element composition display considerable scatter (e.g., Fig. 5), as do Os isotopic compositions. For lherzolites, at least, the different concentrations cannot entirely be an artifact of heterogeneous distribution of sulfide grains within sample powders or the rock (Meisel and Moser 2004; Meisel and Horan 2016, this volume). Instead, these concentration variations may reflect the compositional variability of sulfide grains in the rock; as indicated by variable Ir and Ru concentrations in peridotitic sulfides (e.g., Alard et al. 2000). As there is indisputable evidence for widespread, or even ubiquitous, chemical and isotopic disequilibrium of the HSE in peridotites, it is plausible that the distribution of chalcophile elements between peridotite and magma is partly controlled by the composition of sulfide liquids from infiltrating primitive magmas and partly by mixing processes between such liquids and sulfide liquids already present in the rocks (e.g., Lorand et al. 1999; Alard et al. 2000; Pearson et al. 2004; Lorand et al. 2010). In the melting model shown in Figure 18b apparent sulfide–silicate partition coefficients were used to match the patterns of peridotites from the Balmuccia peridotite massif. Apparent partition coefficients take into account the extent to which the HSE composition of peridotites displays the effects of mixing, and thus the influence of the original infiltrating melt compositions, rather than just sulfide melt–silicate melt equilibrium. It is clear that the fractionations inherited from the melt contribute to the lowering of Db, compared to the equilibrium case. The differences will be particularly notable for Pd and Au. As Pd in depleted lherzolites is commonly slightly depleted, the apparent bulk distribution coefficient for this element should be < 1 and apparent sulfide–silicate distribution coefficients in the model in Figure 18b would be about 1300; far lower than the 105–106 range for sulfide–silicate equilibrium (Mungall and Brenan 2014). For Pt and Rh apparent partition coefficients may also be lower. Gold abundances in depleted lherzolites are lower than in fertile lherzolites and this, coupled with the slight enrichment of Au in primitive basaltic magmas, suggests that Au also has an apparent bulk distribution coefficient < 1. Consequently, apparent sulfide–silicate distribution coefficients for Au are significantly lower (about 200 in the case of Fig. 18b) than equilibrium values (4000–10000; Mungall and Brenan 2014). Rhenium and other moderately chalcophile elements with equilibrium sulfide–silicate partition coefficients < 1500 are not sensitive enough to identify chemical disequilibrium, as the influence of the silicate mineral– silicate melt partition coefficients is substantial. Combined sulfide–silicate and silicate mineral– silicate melt partition coefficients of these elements yield bulk partition coefficients < 1, whether or not equilibrium is assumed. Figure 20 displays the variation of Re concentrations versus Pd concentrations in various suites of mantle tectonites (note that in more strongly serpentinized peridotites, such as from the Oman ophiolite, Re may also be affected by late-stage alteration). Both elements tend to correlate in harzburgites and in depleted lherzolites, however, in more fertile rocks, Re displays larger variations (0.07–0.4 ng/g) at relatively constant Pd (5–9 ng/g). The most likely explanation for this observation is that sulfide and other HSE carrier populations in harzburgites and depleted lherzolites reflect mixing and full disequilibrium, whereas preexisting phases in fertile lherzolites may have partially reacted and equilibrated with a larger

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Pd (ng/g) Figure 20. Concentrations of Re and Pd in peridotite tectonites and evolution of the composition of residues in different melting models. Symbols as in Figure 19, melting curves A to C calculated using parameters from Figure 18 and the text. A: equilibrium liquid sulfide–silicate partitioning (Fig. 18a), B: disequilibrium distribution, taking into account the effect of mixing of different types of sulfide with different partitioning histories (Fig. 18b), C: mss–liquid sulfide–silicate partitioning (Fig. 18c). Different Re/Pd ratios in lherzolites are indicated by dashed lines. None of the melting models yields a satisfactory match for the data distribution of different peridotite suites. In this diagram, ideal binary mixing processes without chemical reaction should result in linear correlations; e.g., mixing of ‘residual’ Re- and Pd-depleted sulfide liquid with Re–Pd-rich sulfide liquid precipitated from percolating magma. Most peridotites from Lanzo display such a trend along a Re/Pd of 0.05. Depleted lherzolites and harzburgites from Baldissero and Balmuccia also display a linear trend albeit at a lower Re/Pd, presumably because the infiltrating magma was more depleted in Re and other incompatible elements. In fertile lherzolites the data is scattered, likely because of the predominance of sulfides derived from infiltrating magma and partial chemical equilibration. Chemical equilibration tends to decouple variations of Re and Pd because of their very different partitioning behavior at low to moderate degrees of melting (Brenan et al. 2015, this volume). Symbol key: Balmuccia: solid circle (dunites at low Re and Pd concentrations), Baldissero: open circle, Lherz: x, Turon de la Tecuere: +, Lanzo: solid diamond, Internal Ligurides: open diamond, External Ligurides: gray diamond, Ronda: open triangle, Beni Bousera: solid triangle. Also shown are harzburgites (solid square) and dunites (open square) from the Wadi Tayin section of the Oman ophiolite. For data sources of peridotites see Figure 5 and text.

fraction of silicate melt and sulfide liquid. The data also suggest that HSE carriers in fertile peridotites of some suites (e.g., Balmuccia and Baldissero) must be more depleted in Re than other suites, which may be a property of the melts that precipitated sulfides during reactive infiltration. The curved trend defined by some data in Figure 20 may be related to the quantity of melt that reacted and precipitated sulfide liquid in the rock. The systematic behavior of Pd, Au, Re, and of other chalcophile elements such as S, Se, Te, Cu, and Ag in most peridotites and in MORB (Wang and Becker 2015b) indicates that the relative depletion and enrichments of these elements in peridotites and in MORB may be described by apparent bulk partition coefficients. Melt compositions calculated by this approach may yield similar concentrations of Pd, Au, and Re as in primitive MORB, although the latter almost certainly require a more complicated fractionation history (e.g., Langmuir et al. 1992; Rehkämper et al. 1999; Bezos et al. 2005; Mungall and Brenan 2014; Wang and Becker 2015c).

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An alternative model of HSE partitioning during mantle melting was presented by Bockrath et al. (2004) and Ballhaus et al. (2006). These authors proposed that residual mss may coexist with liquid sulfide over a significant pressure-temperature range in the mantle. Partitioning between these phases may control the HSE abundances in residues and silicate melts. However, because of uncertainties in the position of the sulfide liquidus in different experimental studies, the stability of mss in the asthenosphere or deeper lithosphere is debated (see Fonseca et al. 2012; Mungall and Brenan 2014). The relevance of mss–liquid sulfide partitioning in the upper mantle can be evaluated on the basis of existing partitioning data for chalcophile elements and the composition of mantle rocks, basalts and their sulfides. Melting models of bulk rock compositions of lherzolites that employ mss–liquid sulfide partition coefficients (Fig. 18c) display a poor match for Pt, Pd and Au. However, it must be acknowledged that bulk partition coefficients are strongly influenced by the silicate mineral– silicate melt partition coefficients. Only for olivine-silicate melt partitioning does sufficient data exist for Pt, Pd, and Au (see Eqns. (11)–(13) in Mungall and Brenan 2014, which yield low Dolivine/silicate melt for these elements at fO2 of 10-9–10−10 bar). Pyroxene-silicate melt partition coefficients for these elements are poorly constrained, and thus Db may be higher. As for sulfide liquid-silicate partition models, Re fits well because its Db is strongly controlled by the large mass fraction of silicates and the well-determined mineral–silicate melt partition coefficients. In principle, mss–liquid sulfide partitioning may account for the different patterns of Ir-group and Pt-group PGE in sulfide inclusions and sulfides on grain boundaries in peridotites (e.g., Ballhaus et al. 2006). However, the behavior of Re concentrations in sulfide inclusions versus grain boundary sulfides argues against this process. Equilibrium mss-–liquid sulfide partitioning would predict higher Re and Os concentrations in residual sulfides compared to coexisting sulfide liquids, because both elements are compatible in mss (DOsmss/sul liq = 3–7, DRemss/sul liq = 3, Brenan 2002; Ballhaus et al. 2006). Although sulfide inclusions in silicates of peridotites may have higher Ir and Os than sulfides on grain boundaries (e.g., Alard et al. 2000, 2002), Re is depleted in the former and enriched in the latter, commonly accompanied by correlated Re/Os (Alard et al. 2005). Recently, it has been proposed that some harzburgites contain sulfides with high Se/Te ratios similar to what is expected from mss–liquid sulfide partitioning (König et al. 2014, 2015). However, because of the low concentrations of these elements, the mass balance of such phases in strongly depleted peridotites is difficult to constrain, and they may also reflect precipitation of sulfide from somewhat more fractionated magma with high Se/Te and Re/Os (Wang and Becker 2015a). Work on Cu and Ag abundances in peridotites has shown that the relative behavior of these elements in bulk rock lherzolites is consistent with the systematics predicted by sulfide liquid-silicate partitioning but not with mss–liquid sulfide partitioning (Wang and Becker 2015b). The differing 187Os/188Os of the two sulfide populations suggests that sulfides precipitated on grain boundaries during melt infiltration did not equilibrate with included sulfides, which is a basic requirement for equilibrium mss–sulfide liquid–silicate melt partitioning models. Thus, as shown before in the discussion of sulfide liquid–silicate melt partitioning, none of the proposed partitioning processes that are potentially relevant during partial melting yields a satisfactory quantitative description of the HSE composition of many mantle peridotites. Sulfide melt–silicate melt partitioning seems to be the best match for the observed HSE pattern in lherzolite bulk rocks. However, at least for Pd, Au, Re, and S, their ratios in lherzolites may be mostly inherited from the melts that infiltrated depleted precursor rocks (e.g., harzburgites; Fig. 20). The origin of the HSE fractionation in the infiltrating melts and their sulfide liquids will be discussed below. HSE fractionation during the formation of mantle pyroxenites. Mantle pyroxenites are important because they represent products of magmatic fractionation in the mantle and thus yield information on the composition of relatively ‘primitive’ magmas (Bodinier and

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Godard 2003). Pyroxenites are cumulates that formed by reactive infiltration and fractional crystallization of primitive to more evolved basic magmas. Websterites (‘Cr diopside suite’) and orthopyroxenites sometimes display mineralogically zoned reaction domains with peridotites, which have formed due to melt infiltration into the surrounding peridotite (e.g., Bodinier et al. 1987, 2008; Becker et al. 2004). Quite often, clinopyroxenites (‘Al augite suite’) appear to have formed from more evolved compositions and the absence of reaction zones may indicate their formation at shallower levels (e.g., Sinigoi et al. 1983; Suen and Frey 1987). Only limited data are available for HSE abundances and Os isotopic compositions in mantle pyroxenites from tectonites, including pyroxenites from Ronda (Reisberg et al. 1991; Reisberg and Lorand 1995; Marchesi et al. 2014), Beni Bousera (Kumar et al. 1996, Pearson and Nowell 2004; Luguet et al. 2008b), Lower Austria (Becker et al. 2001, 2004), Troodos (Büchl et al. 2002), Totalp (van Acken et al. 2008, 2010b), Hori Bory (Ackerman et al. 2013) and Balmuccia (Wang and Becker 2015c). The HSE patterns of pyroxenites in mantle tectonites are broadly similar to data from sulfides in pyroxenite xenoliths. In general, the relative fractionation of the HSE is similar to that in basalts, but with higher concentrations of Os, Ir, Ru, Rh, Pt, and Pd than in MORB. Websterites and orthopyroxenites often display HSE patterns that are less strongly fractionated than clinopyroxenites (Fig. 21). Concentrations of S and Re in pyroxenites are similar or lower than in MORB, but often higher than in lherzolites. Abundances of other HSE in pyroxenites are similar or lower than in lherzolites (Fig. 21). Some pyroxenites display a depletion of Re relative to Pd, which may have been caused by multi-stage melting (Marchesi et al. 2014). The occurrence of cmscale Os isotopic heterogeneity between alternating pyroxenite–peridotite layers (Becker et al. 2001, 2004; Büchl et al. 2002; van Acken et al. 2008) is another indication of the difficulty of small-scale Os isotopic equilibration between silicate melt and existing sulfide populations. A study of a zoned clinopyroxenite–websterite–orthopyroxenite rock from Lower Austria that represents a former reaction zone between high-temperature silicate melt and peridotite has shown that Sr and Nd isotopic compositions were equilibrated across a 10-cm distance of the rock at the time of its formation (Becker et al. 2004). In contrast, both gOsi and Os concentrations display strong gradients over the same distance, indicating disequilibrium. HSE compositions of sulfides in single thin sections of Totalp pyroxenites vary from those with Ru/Ir, Pd/Ir, and Re/Ir similar to peridotitic sulfides, to those with high ratios of these elements, typical of melt compositions (van Acken et al. 2010b). The detailed processes that resulted in the close association of these different sulfide populations are not yet clear, but they suggest that disequilibrium among sulfides may be common in mantle pyroxenites as well as peridotites. A comparison of Re/Os and Pd/Ir in pyroxenites with data on ocean ridge basalts and gabbros from the lower oceanic crust indicates considerable overlap (Fig. 22). This observation suggests that significant fractionation of HSE ratios in magmas already occurs by precipitation of sulfide liquid during magmatic transport and reaction in the mantle (Wang and Becker 2015c). In contrast to Re/Os, which shows large variations in magmatic products over several orders of magnitude, the variation of Pd/Ir in the latter is much more limited and Pd and Ir show similar bulk partitioning behavior. Because of the segregation of sulfide liquid from magmas during magmatic transport in the mantle, the HSE compositions of basaltic magmas may preserve little direct information on HSE concentrations of deeper parts of the melting region. Figure 22a also shows that the data fields defined by most magmatic products, particularly the basalts, are offset from the bulk compositions of peridotites, but overlap with ratios in grain boundary sulfides from peridotites. A similar observation was made for variations of Se/Te (Wang and Becker 2015c). This observation may provide the best indication so far that most magmas that contribute to the oceanic crust did not fully equilibrate with the bulk rock of mantle peridotite residues.

Re-Pt-Os Isotopic and HSE Behavior in Mantle Tectonites 10.0000

Rock/PM

1.0000

Lherzolites

413

a

Websterites

0.1000 0.0100 Clinopyroxenites

0.0010

Balmuccia (Ivrea Zone) 0.0001

Os

Ir

Ru

Rh

Pt

Pd

Au

Re

100.00

Rock/PM

10.00

b

Totalp pyroxenites (Swiss Alps)

Websterites

Lherzolites 1.00

0.10 Clinopyroxenites 0.01

Os

Ir

Ru

Pt

Pd

Re

10.0

Rock/PM

Ronda websterites and hybrid peridotites

c

Websterites

Normal lherzolites 1.0 Hybrid lherzolites

0.1

Os

Ir

Ru

Pt

Pd

Re

Figure 21. Primitive mantle-normalized concentration diagrams of mantle pyroxenites from peridotite massifs. Websterites: gray lines, clinopyroxenites: black lines. a) Balmuccia (Wang and Becker 2015c) Balmuccia lherzolites from Wang et al. (2013). b) Totalp (van Acken et al. 2010b) Totalp lherzolites from van Acken et al. (2010a). c) Ronda (Marchesi et al. 2014): Hybrid lherzolites (dashed lines) were also affected by reactive infiltration of magma, but differ in composition from the pyroxenites and normal lherzolites. Typical Ronda lherzolites (dash-dotted lines) from Fischer-Gödde et al. (2011).

HSE fractionation during the formation of harzburgites and replacive dunites. Data on HSE and other chalcophile elements in harzburgites show that many of these rocks have high abundances of IPGE and lower abundances of Rh, Pt, and Pd (e.g., Pearson et al. 2004; Becker et al. 2006; Luguet et al. 2007). These IPGE–PPGE fractionations are generally consistent with fractionation of melting residues at moderate to high (15–30 %) degrees of partial melting (Mungall and Brenan 2014; Brenan et al. 2016, this volume, and references therein). The incongruent breakdown of liquid or solid sulfide occurs at advanced degrees of melting at low fS2 and may play an important role in the stabilization of Os–Ir–Ru and Pt–Ir alloy phases that

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BM web+op

BM cp

Other pyroxenites

BM peridotites

Oman gabbros

MORB

Cumulates

10 1 0.1

Harzburgite

Magmatic evolution

0.01

Melting

Lherzolite

Peridotites

Pd (ng/g) bulk rock

100

Fractionated melt Apparent ratio: (DPd-1)/(DIr-1)=0.95 (±0.11,2σ)

0.001 0.0001

0.001

0.01

0.1

Ir (ng/g) bulk rock

1

10

1000 Clinopyroxenites Interstitial sulfides in peridotites

100

Pd/Ir

10

Oman gabbros

Websterites

1

0.1

MORB

ck ro s lk tite Bu rido pe

HSE fractionation during magma transport in the mantle and in the oceanic crust

Sulfide inclusions in peridotites

0.01

0.001 0.01

0.1

1

10

100

1000

Re/Os Figure 22. a) Pd–Ir diagram of bulk rock concentrations in mantle peridotites, pyroxenites (BM = Balmuccia, Ivrea Zone), MORB and gabbros of the oceanic crust. Web: websterite, op: orthopyroxenite, cp: clinopyroxenite. The correlation suggests that, with the exception of a few gabbros and MORB, mantlederived magmatic rocks define a continuum between melt compositions and pyroxenites (‘cumulates’). Most magmatic products are offset from the peridotite data, indicative of disequilibrium between magmas and bulk peridotite. b) Pd/Ir–Re/Os diagram showing the limited range of fractionation of Pd/Ir in magmatic products compared to Re/Os. The Pd/Ir data are consistent with similar bulk distribution coefficients of these elements during magmatic processing in mantle and crust (see a). Most magmatic rocks shown in (b) define fields that overlap with or lie along the extension of grain boundary sulfides from peridotites, indicating a common origin of grain boundary sulfides and mantle-derived igneous rocks. Both diagrams are modified from Wang and Becker (2015c). Data sources: Oman gabbros, Peucker-Ehrenbrink et al. (2012); MORB: Hertogen et al. (1980), Rehkämper et al. (1999b), Bezos et al. (2005), Lissner et al. (2014); BM pyroxenites and peridotites, Wang et al. (2013, 2015c); other pyroxenites are from Totalp, van Acken et al. (2008, 2010b); Beni Bousera, Luguet et al. (2008); Ronda, Marchesi et al (2014); Horní Bory, Ackerman et al. (2013); Dramala massif, Pindos ophiolite, Sergeev et al. (2014); interstitial sulfides and sulfide inclusions in peridotites, Alard et al. (2005), Harvey et al. (2006).

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have been found in such rocks (Lorand et al. 1999, 2010; Luguet et al. 2007; Fonseca et al. 2012; Mungall and Brenan 2014; Brenan et al. 2016, this volume). With progressive melting in the absence of a Fe–Ni-rich sulfide phase, all Re, Au, and Pd should be dissolved in coexisting melts, provided that residues and melts were equilibrated. The abundances of Os, Ir, Ru, Rh, and Pt, and their fractionation in harzburgite residues (e.g., Fig. 20) should be controlled by the solubility of these elements in sulfur-bearing silicate melts and the stability of Os–Ir, Ru–Os, and Pt–Ir phases (Mungall and Brenan 2014). However, harzburgites may show variations in HSE abundances that are not entirely consistent with a simple melting history as envisioned before. Normalized abundances of Re and S in harzburgites are sometimes higher than normalized abundances of Pd (Figs. 5, 10). These patterns have been interpreted either in terms of precipitation of secondary sulfides from infiltrating melts with high Re/Os and fractionated HSE patterns (Chesley et al. 1999, Pearson et al. 2004, Becker et al. 2006; Wang and Becker 2015a). Alternatively, enrichments of Re and S compared to Pd and Pt (and of Se relative to Te) in some harzburgites have been interpreted to reflect the presence of mss of residual origin (König et al. 2014). The former explanation is consistent with magmatic re-enrichment processes of incompatible elements (e.g., light rare earth elements) in some of these rocks. Some harzburgites display lower abundances of IPGE than expected for depleted mantle peridotite, e.g., < 3 ng/g Ir, instead of 4–5 ng/g expected for residues of moderate to high degrees of melting (Figs. 5, 10). In order to understand this behavior, it is useful to recall that even at high temperatures most peridotites likely contain unequilibrated sulfide melt (maybe also mss), with a range of HSE concentrations. Complete dissolution of some of these sulfide droplets (but not others) into sulfur-undersaturated melt, without concurrent precipitation of IPGE alloy phases, will result in a net decrease of the abundances of all HSE. This process almost certainly plays an important role in the formation of some replacive dunites and associated harzburgite-–lherzolite–pyroxenite rock assemblages (Becker et al. 2001, 2004; Büchl et al. 2002, 2004; Hanghøj et al. 2010; Wang et al. 2013). For instance, the variable IPGE abundances and strong depletions of Pt, Pd, Re and other chalcophile elements in discordant dunite bodies in lherzolites at Balmuccia indicate that the magmas were undersaturated in sulfur, which caused the dissolution of sulfides from the lherzolitic protoliths of the dunites (Fig. 5, Wang et al. 2013). The harzburgites from Wadi Tayin (Oman ophiolite) display normal abundances of IPGE and tend to show primitive mantle-like or even slightly suprachondritic abundances of Pt, Pd and Re (Lorand et al. 2009; Hanghøj et al. 2010). Some of the harzburgites show selective enrichments of Pt that also have been noted from abyssal peridotites and other ophiolites (Fig. 10) and peridotite massifs (Fig. 5). The Pt enrichments may indicate the precipitation of Pt-enriched sulfide liquid from silicate melt that may have dissolved Pt from destabilized Pt–Ir alloys at high degrees of melting. Dunites from Wadi Tayin are similarly enriched in HSE, but show more fractionated Re/Os and PPGE/IPGE ratios. Because the dunites are thought to reflect pathways of olivine-saturated magmas, the enrichments of Pt, Pd, and Re in dunites and harzburgites likely reflect sulfide segregation from magmas enriched in these elements (Fig. 23). Although this process appears to have occurred pervasively, the initial 187Os/188Os (at around 90–95 Ma) in the mantle section at Wadi Tayin were not equilibrated (Fig. 23). The high abundances of Pt, Pd, and Re in otherwise incompatible element depleted mantle rocks suggest that sulfide saturation may play an important role in the uppermost mantle underneath fast-spreading ocean ridges. Dunites from the Troodos ophiolite also display ‘melt-like’ HSE compositions (Büchl et al. 2002). A common property of dunites is that their initial 187Os/188Os ratios extend to suprachondritic values (gOsi ranging from −3 to +17, e.g., Fig. 23 and Becker et al. 2001), suggesting that some of the parent magmas had suprachondritic Os isotopic compositions. However, as the case of the dunites from Balmuccia shows, not all dunites are characterized by an enrichment of Pt, Pd and Re and melt like HSE patterns.

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416 20 LT HT

γOs (90 Ma)

15 10

a

Hz Du

5 1

0 -5 -10

PM

2

Typical mantle peridotite

1

20

10

100

Pd (ng/g)

b

γOs (90 Ma)

15 10 5

2

0

Typical mantle peridotite

-5 -10

PM

1

1

10

Cu (µ µg/g)

100

Figure 23. The enrichment of chalcophile elements in harzburgites and dunites from Wadi Tayin (Oman ophiolite, Hanghøj et al. (2010). a) gOsi–Pd diagram shows that in most harzburgites and dunites Pd is enriched in comparison to typical mantle peridotites. b) gOsi–Cu diagram indicates that Cu in dunites loosely correlates with gOsi. In general, Cu is less enriched than Pd. Open symbols are low-temperature rocks, filled symbols high-temperature rocks (see Fig. 9). Arrow 1 indicates the expected depletion behavior due to melting, 2, redistribution of Pd due to dissolution and precipitation of sulfides and the dash-dotted arrow indicates correlated changes in gOsi, and Cu concentrations resulting from melts with suprachondritic Os isotopic composition. For Pd this correlation breaks down, presumably because of local sulfide segregation from coexisting magma.

PGE enrichments also occur in podiform chromitites, which are magmatic precipitates associated with dunites and harzburgites in ophiolites that formed in the proximity of convergent plate margins. Because chromitites may represent economically relevant sources of PGE, these hightemperature magmatic ore deposits will be discussed in Barnes and Ripley (2016, this volume).

Summary—Mantle melting and mantle–magma interaction—different sides of the same coin Models of partial melting of mantle tectonites must consider the natural open-system behavior relevant for melting column models, diapiric upwelling of partially molten mantle or conversion of lithospheric mantle to asthenosphere by melt infiltration (as was suggested to have occurred in the magmatic history of some mantle tectonites, e.g., Müntener et al. 2005). Thus, melt infiltration and melting should occur more or less simultaneously, provided that porosity and permeability permit melt infiltration. The composition of the residues will change

Re-Pt-Os Isotopic and HSE Behavior in Mantle Tectonites

417

with time until external processes cause upwelling and melting to stop and the mantle to cool. The HSE concentration and 187Os/188Os data on mantle tectonites with well-constrained ages (e.g., Oman ophiolite) show that the extent of sulfide–silicate equilibrium in these melting processes must be limited. Several different types of sulfide (presumably mostly liquids, but also mss and other solid phases at lower temperatures) may exist at high temperatures in peridotite (see also Lorand and Luguet 2016, this volume). Residual sulfides with subchondritic 187 Os/188Os occur as inclusions in silicates and are inherited from ancient melting processes. These sulfides may represent residual sulfide liquids or mss, or both. Sulfide liquids with chondritic to suprachondritic 187Os/188Os and higher Re/Os and Pd/Ir are precipitated from infiltrating silicate melt and mostly reflect the composition of these melts with variable reaction with peridotite. Hybrid sulfide liquids may form locally where magmas and peridotite react and magmas became oversaturated in sulfur. In addition, relic PGM phases such as Pt–Ir alloys inherited from depleted protoliths may survive these magmatic processes. An important aspect of melt infiltration in the lherzolite stability field is the co-precipitation of sulfides with pyroxene ± Al phase assemblages. Only such a process can explain correlations of Re, Re/Os, and sulfur concentrations with fertility indicators such as Al2O3. As it is likely that the same processes were also responsible for the correlations between 187Os/188Os and Al2O3 in many suites of mantle peridotites, the mass balance with inherited Re-depleted sulfides suggests that the infiltrating melts had suprachondritic 187Os/188Os (the origin of such melts will be discussed later). This notion is supported by Os isotopic measurements on grain boundary sulfides in peridotites and by initial Os isotopic compositions of most mantle pyroxenites (Alard et al. 2002, 2005; Harvey et al. 2010, 2011, 2016, this volume; Wang and Becker 2015c). Different modeling approaches, both complicated and simple may produce appropriate HSE compositions of basalts from model mantle compositions (e.g., Rehkämper et al. 1999; Bezos et al. 2005; Harvey et al. 2011; Mungall and Brenan 2014). As discussed here and elsewhere (e.g., Lorand et al. 1999; Pearson et al. 2004; Lorand and Alard 2010; Fischer-Gödde et al. 2011; König et al. 2014; Wang and Becker 2015a), models that employ equilibrium distribution of the HSE between mantle phases have difficulties in accounting for some of the detailed compositional variations of the compatible HSE in bulk peridotites. Studies of HSE in bulk rocks of mantle peridotites and pyroxenites and their trace phases indicate that in high temperature magmatic processes in the mantle, disequilibrium between different HSE host phases and silicates may be the rule (e.g., Burton et al. 1999; Alard et al. 2000, 2002, 2005). In spite of these complexities, a useful assessment of the bulk distribution behavior of the HSE is possible and their relative behavior is consistent with abundance data in komatiites and basalts. The data on bulk rocks and sulfides of mantle pyroxenites and sulfides from grain boundaries in peridotite tectonites and in xenoliths indicate that infiltrating melts show relative fractionation of the HSE and S similar to the fractionation pattern of basalts, with mantle normalized abundances of S ≈ Re > Au > Pd > Pt ≥ Rh > Ru > Ir ≥ Os. The HSE data on peridotites and pyroxenites suggest that the composition of infiltrating melts also affects the composition of peridotites (e.g., Figs. 5, 7 20). Notably, enrichments and depletions of Re in peridotites may be caused by precipitation of sulfides with suprachondritic Re/Os. If the abundances of Re, Au, Pd, Pt, and other chalcophile elements in mantle peridotites are predominantly controlled by sulfide segregation from primitive basic magma, the question arises, which partition process produced the relative fractionation among these elements in these magmas to begin with? The answer may lie in the increasing importance of alloy solubility in silicate melt during moderate to high degrees of melting in the shallow mantle, near or beyond the exhaustion of sulfide in the residues. At these conditions, the concentrations of the HSE in silicate melts may be controlled by residual PGE alloys, the different solubility of Pt, Rh, Ru, Ir, and Os and possibly silicate mineral-oxide-melt partitioning (Mungall and Brenan 2014; Brenan et al. 2016, this volume). Thus, basic melt infiltrating the asthenosphere and lithosphere at greater depth likely carries the HSE and 187Os/188Os signature of oceanic crust produced in previous Wilson cycles.

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This conclusion is consistent with suprachondritic initial 187Os/188Os of mantle pyroxenites and some peridotites that were affected by melt infiltration and coexisting harzburgites with subchondritic 187Os/188Os, which may represent ancient remnants of shallow oceanic mantle.

Os isotopic heterogeneity in the mantle The compatibility of Os during partial mantle melting, and the existence of two radioactive decay systems producing isotopes of Os, makes it an ideal element with which to investigate mantle heterogeneity (Hart and Ravizza 1996; Burton et al. 1999). The relative compatibility of Os and Re is primarily controlled by their differing preference for sulfide over melt (See section above: Behavior of HSE during partial melting). This produces strong fractionation of moderately incompatible Re from compatible Os during partial melting of the mantle, giving rise to very high Re/Os ratios in crust-forming melts (see Gannoun et al. 2016, this volume) and correspondingly low, sub-chondritic 187Os/188Os ratios in depleted mantle. In turn, crust recycled back into the mantle is potentially traceable due to its distinct Os isotope signature. Likewise, small degree melts within the mantle may also produce variations in Re/Os and thus, over time, in 187Os/188Os. Due to the chalcophile affinity of Os, Re–Os isotope variations can provide different, yet complementary, information to lithophile isotope systems, and can display behavior that is decoupled from lithophiles (e.g., Class et al. 2009). The 190Pt–186Os decay system, in contrast to the Re–Os system, does not typically produce resolvable differences in 186Os/188Os ratios in mantle rocks due to the much smaller decay constant compared to 187Re, and due to the lower degree of fractionation between parent and daughter. Only in specific cases of high-degree melting do Pt concentrations significantly exceed those of the mantle, such as in some volcanic arc settings (Dale et al. 2012b) and in komatiites (e.g., Puchtel and Humayun 2001; Fiorentini et al. 2011); but in the latter case Os in the melt approaches mantle concentrations and thus fractionation of Pt and Os remains limited. Recycled crust has only moderately high Pt/Os (Dale et al. 2009a; Peucker-Ehrenbrink et al. 2012) which is not sufficient to produce anomalous compositions given the subsidiary Os concentrations of crust, relative to mantle. Nevertheless, 186Os enrichments have been identified in some intraplate magmas (Brandon et al. 1998, 2003; Puchtel et al. 2005) and in a later section we briefly discuss whether mantle processes are a plausible mechanism by which to produce these enrichments. In this section, we focus on broad-scale mantle heterogeneity, whereas disequilibrium on a hand specimen scale, or smaller, is covered in the previous section Os isotopic disequilibrium. 187

Os/188Os mantle composition and heterogeneity. The bulk Os isotope composition of the silicate Earth was likely set by late accretion of material with a bulk primitive composition, after core formation had ceased (Kimura et al. 1974; Chou 1978). However, neither the 187Os/188Os composition (Meisel et al. 2001) nor the relative HSE abundances of PM estimates (Becker et al. 2006) match those of any known chondrite group. This difference has been reconciled by (i) late accretion of differentiated planetesimal core material and primitive chondritic material (Fischer-Gödde and Becker 2012), (ii) by a hybrid model for the enrichment of Earth’s HSE involving late accretion to a fractionated mantle signature (which may be a residue from metal-silicate segregation, cf. Righter et al. 2008; Walker 2009), or (iii) by mantle processes accounting for the combination of non-chondritic ratios involving Ru and Pd and chondritic ratios of other HSE in fertile lherzolites (e.g., Lorand et al. 2010). See Day et al. (2016 this volume) for further discussion. The processes of continental crust production and incomplete rehomogenization of recycled oceanic crust have likely both served to reduce the 187Os/188Os of the peridotitic mantle below that of the primitive mantle. Thus, heterogeneous distribution of 187Os in the mantle is due to the timing and degree of melt depletion and the presence of enriched domains, which may either be recycled crustal materials or domains within the mantle fertilizes by low-degree melts.

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Table 2. Summary of compiled 187Os/188Os data for mantle tectonites, by setting and sample type. Sample type

Mean Os/188Os

187

187

Mode Os/188Os

Low

High

Main range (% included)

n

Abyssal peridotites

0.1243

0.1261

0.1139

0.1382

0.024 (100%)

107

Continent/continent–ocean transition

0.1255

0.1262

0.1126

0.1372

0.025 (97%)

156

High-T convergent margin

0.1259

0.1237

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48

Ophiolites (all*)

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Arc xenoliths

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OIB xenoliths

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0.1248

0.1138

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0.026 (99%)

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Sub-continental xenoliths

0.126

0.1257-67

0.1094

0.1464

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* 2 Ga Finland ophiolite localities omitted due to long-term isolation from convecting mantle. For data sources see captions for Figs. 14 and 24.

A compilation of 187Os/188Os data for global peridotites (excluding pyroxenites), grouped according to the tectonic settings used in this chapter and in this volume, is shown in Figure 24, and a summary of the averages and ranges for each setting/sample type is shown in Table 2. Cratonic and circum-cratonic xenoliths, which won’t be discussed further here, are both typically strongly unradiogenic, reflecting their severe and early melt depletion and subsequent isolation from the convecting mantle (see Aulbach et al. 2016, this volume, and references therein). All major tectonite and xenolith groups (continental/continentocean transitional tectonites, high-T convergent tectonites, ophiolites, abyssal peridotites, oceanic mantle xenoliths, sub-continental lithosphere xenoliths and sub-arc xenoliths) have a considerable ‘peak’ in probability of 187Os/188Os between 0.125 and 0.128, indicating a degree of effective large-scale homogenization in the convecting mantle and younger lithosphere, albeit incomplete. Moreover, most groups have remarkably similar total ranges of 187Os/188Os (when excluding up to 3% of the most extreme data), between 0.026 and 0.029 units, with the exception of high-T convergent margin tectonites (n = 48) which have a range of 0.023, and sub-continental lithospheric mantle xenoliths, with a larger range of 0.037 (although in this latter case the primary data may be compromised by secondary processes such as weathering and reaction with host melts. Greater than 85% of samples from each tectonic setting fall within a narrower range of 187Os/188Os of around 0.015 units (the range of each group varies from 0.013 for all ophiolites, to 0.019 for continental/continent–ocean transition tectonites). In detail, however, each grouping displays a variable distribution of Os isotope composition, and the positions of the modal and mean 187Os/188Os compositions differ between many of the groupings. One caveat here is that the data plotted on Figure 24 are present-day measured 187 Os/188Os ratios, to reflect the current degree of overall mantle heterogeneity, and thus do not account for any isolation of portions of lithosphere sampled in this dataset. If these portions were exposed to gradual convective stirring then some of the ‘older’ depletion ages may have been remixed with more radiogenic ambient mantle. Not all components of the compilation, therefore, necessarily reflect the composition of the ‘convecting’ mantle. All tectonite groups have ranges that extend to sub-chondritic and supra-chondritic Os/188Os ratios, although some extend broadly equally in each sense, while others have a pronounced skew towards less or more radiogenic values. For instance, the ophiolite record has a modal 187Os/188Os of ~ 0.1255, with a broadly equal number of data extending in each sense down and up to values of 0.115 and 0.143, respectively (Fig. 24). At least half of the data fall between 0.1225 and 0.128. In contrast, the dataset for continental/continent-ocean transitional tectonites shows a modal 187Os/188Os of ~ 0.126, close to that of ophiolite ultramafics, but with 187

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Relative probability

Cratonic xenoliths (excl. China); n=127

CC OC EC

Continent/ Cont.-Ocean transition tectonite; n=156

CC OC EC

Relative probability

Ophiolites; n=142

PM

Abyssal peridotites; n=107

Circum-Craton xenoliths; n=70

SCLM xenoliths; n=228

CC OC EC

PM

OIB xenoliths; n=134

Relative probability 0.105

PM

High-T convergent tectonite; n=48

Arc xenoliths; n=37

0.110

0.115

0.120

0.125

187

0.130

Os/188Os

0.135

0.140

0.145

Figure 24. Probability density plots of present-day 187Os/188Os ratios in whole-rock samples grouped according to the tectonic settings discussed in this chapter: Ophiolites, abyssal peridotites, continent/continental–ocean transitional tectonite, high temperature convergent tectonite. Xenoliths from the subcontinental lithospheric mantle, oceanic lithosphere, cratonic lithosphere and circum-cratonic lithosphere are also shown (see Aulbach et al. 2015, this volume, and Luguet and Reisberg 2015, this volume, for a discussion of HSE in these xenolith groups). Ranges for primitive mantle (Meisel et al. 2001) and major chondrite groups also shown; CC – carbonaceous, OC – ordinary, EC – enstatite (Walker et al. 2002a). A universal uncertainty of 0.00125 was applied to each datum to avoid bias towards more precise analyses and to provide sufficient smoothing for the smaller datasets, where used. For data sources see Fig. 14, except cratonic xenoliths: Walker et al. (1989); Pearson et al. (1995a,b, 2004); Shirey and Walker (1995); Chesley et al. (1999); Meisel et al. (2001); Becker et al. (2006); Maier et al. (2012), and circum-craton xenoliths: Pearson et al. (2004); Luguet et al. (2009); Aulbach et al. (2014).

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a range extending down to 0.112 and up to 0.133, with a lower mean value than for ophiolites (Fig. 24). The abyssal peridotite samples of the convecting mantle show a remarkably similar probability profile to the continental/transitional tectonites, with a modal 187Os/188Os of ~ 0.126, and a range from 0.1125 to 0.140; possibly with similar subsidiary peaks at 0.1225 and perhaps even at 0.115 (although this most unradiogenic peak appears important for continental/ transitional tectonites, but likely is not significant for abyssal peridotites, given the sample size). The ‘tails’ to low and high 187Os/188Os reflect, respectively, ancient melt-depleted domains and enriched domains which have not fully re-homogenized with the rest of the convecting mantle through convecting stirring, melt percolation and infiltration. The distribution of the data is further mentioned below in the context of platinum-group mineral studies. Qualitatively, at least, re-enrichment of ophiolitic mantle is supported by the observation that convergent margin ophiolites appear to have more radiogenic 187Os than mid-ocean ridge ophiolites (Fig. 15), and by the absence of a skew to old depleted values in the overall ophiolite 187Os/188Os distribution (Fig. 24; cf. abyssal peridotite and ophiolite curves). The relatively radiogenic distribution of sub-arc xenoliths is also consistent with the process of re-enrichment in the subduction zone environment. The chromitite and PGM record of Os isotope mantle composition and heterogeneity. Here, we focus only on the Os isotope evidence from PGM, rather than the systematics of PGM formation and composition (see O’Driscoll and Gonzáles-Jiménez 2016, this volume, for a comprehensive review). The utility of chromitites, and the PGM that they typically contain, is that they are Os-rich, Re-poor and tend to be largely robust to subsequent alteration processes caused by metamorphism and/or fluid-rock interaction. The very low Re/Os ratios mean that their 187 Os/188Os isotope composition is almost ‘frozen in’ at the point of formation, or at worst require very small corrections for radiogenic ingrowth, even over periods of 3 Ga or greater (Malitch and Merkle 2004). For these reasons, they have been used to estimate the Os composition of the convecting mantle, to assess mantle heterogeneity and to identify potential major mantle melting events through Earth’s history. One caveat to this use, however, is that chromitite formation occurs in zones of high melt flow, and these melts may have imparted a radiogenic 187Os/188Os signature on the chromitite, thus rendering it no longer entirely representative of the ‘average’ upper mantle (e.g., O’Driscoll et al. 2012; see also the section on convergent ophiolites above). A global suite of ophiolitic chromites was used to provide an estimate of the average Os/188Os composition of the convecting mantle (Walker et al. 2002b). Linear regression of the isotope data relative to the age of the chromite provided an evolution curve with a present-day 187Os/188Os composition of 0.1281. Although the uncertainties overlap, this best estimate equates to approximately 5% less ingrowth of 187Os over the life of the Earth when compared to the PM (0.1296; Meisel et al. 2001). This is presumably due to continental crust extraction and the presence of recycled oceanic crust in the mantle, which has not (yet) been efficiently rehomogenized. A study of over 700 detrital PGM from the Josephine Ophiolite, California, found a Gaussian distribution of 187Os/188Os ratios from 0.119 to 0.130 (Meibom et al. 2002). This was interpreted to represent long-term heterogeneity (melt-enriched and -depleted endmembers) which has been partially erased and homogenized by metasomatism and melt–rock reaction processes. Further work on a range of global ophiolites, however, indicated a more complex distribution of Os isotope ratios in Earth’s mantle. Over 1000 detrital PGM from ophiolites in California, Urals, Tibet and Tasmania revealed a variety of 187Os/188Os distributions, from close to Gaussian to skewed towards old, unradiogenic values in the case of Urals, and a bimodal distribution for both Tibet and Tasmania (Pearson et al. 2007). It was proposed that the apparent ‘peaks’ in probability for certain 187Os/188Os ratios are consistent across different ophiolites and across other geological settings such as cratonic xenoliths, and that these peaks reflect global signatures produced by major global mantle melting episodes throughout Earth’s history which match the implied crustal growth record from zircon ages. The 187

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composition of the major peak in 187Os/188Os for PGM is 0.1276 (Pearson et al. 2007; adjusted to present-day in Dale et al. 2009b), although the mean composition is likely significantly lower because of the skewed distribution to less radiogenic values. Perhaps notably, when considering representative analyses of convecting mantle composition, this upper limit of 187 Os/188Os composition from PGM analysis is less radiogenic than the average of analyzed chromites (0.1281; Walker et al. 2002b), even though many of the PGM are also sourced from supra-subduction zone ophiolites and therefore may be subject to the same process of radiogenic Os addition. Also of note is the fact that ultramafics from most of the tectonic settings have ‘peak’ values that are slightly less radiogenic than the ‘peak’ value from PGM (see Fig. 24; 187Os/188Os ~ 0.1265, compared to 0.1276). In summary, although global compilations have inherent bias towards exposed and wellstudied areas, all the larger datasets (n > 100) for mantle settings that have not been isolated for long periods (cf. cratons), have very similar modal 187Os/188Os compositions of between 0.125 and 0.127, and mean compositions between 0.1243 and 0.1271. Such values equate to around 8 to 18% less ingrowth of 187Os over the life of the Earth than for PM evolution (cf. Meisel et al. 2001), presumably largely due to crustal extraction and long-term isolation – although the exact degree of mantle Re depletion is dependent on the timing of this extraction. These values are somewhat higher than the 5% estimated from chromitites (see above, cf. Walker et al. 2002b), but some of this discrepancy is due to the omission of pyroxenites and other enriched lithologies from this data compilation. The small variance in the isotopic ranges for each setting appears noteworthy in terms of gauging mantle mixing efficiency, but is beyond the scope of this review. 186

Os/188Os mantle composition and heterogeneity. Platinum-group minerals and chromitites have been used as recorders of the 186Os/188Os evolution of the mantle. Many PGM are IPGE-rich and have low Pt/Os and hence faithfully record the 186Os/188Os of the mantle at the time when those PGM formed. Brandon et al. (2006) used Os-rich PGM data, together with chondrite analyses, to constrain the terrestrial evolution of 186Os/188Os from an initial of ~ 0.1198269 ± 0.0000014 (2 s) at 4.567 Ga to a present-day value of 0.1198382 ± 0.0000028. The potential for large-scale heterogeneity generated by the 190Pt-186Os system is far smaller than that of the 187Re-187Os system, and in most cases is beyond what is distinguishable given current analytical capabilities. Nevertheless, anomalously radiogenic 186Os/188Os ratios have been found in some high-degree melts in intraplate settings in Hawaii, Gorgona Island and Kostomuksha, Russia (Brandon et al. 1998, 2003; Puchtel et al. 2005), coupled with only limited 187Os enrichment. Possible mechanisms to generate such signatures are discussed below. The range of Pt/Os ratios found in the supra-subduction zone environment indicates that there must be huge 186Os variations on a lithological and mineral scale, if those materials were isolated. Alaskan-Uralian complexes (see Johan 2002 for details ) also display a large range of Pt/Os ratios, but these are beyond the scope of this chapter. Chromitites from ophiolites typically possess very low Pt/Os ratios (~ 0.1, compared with 1.95 for the PM), but can sometimes have Pt/Os of > 10 (see ophiolite sections). Platinum-group minerals from within chromitites and other PGE-saturated ores can have even more extreme Pt/Os; laurites (Ru (Os, Ir)S2), may have ratios of < 0.01 (González-Jiménez et al. 2009) while Pt–Fe alloys can have Pt/Os of > 100,000 (Walker et al. 1997). Extremely high Pt/Os ratios, such as those of the Meratus Ophiolite, Borneo (up to 2000), evolve to much higher 186Os/188Os compositions than those of the bulk mantle, and because PGM are largely robust to subsequent processes, they may show isochronous behavior and can be used to date ophiolitic complexes (Coggon et al. 2011). These PGM, after ingrowth over as little as 200 Ma, have 186Os/188Os ratios that range from a slightly sub-PM value of 0.119801–0.120315. As a guide to the magnitude of this difference, it is at least 30 times greater than the difference between the bulk mantle and the highest 186Os/188Os mantle melt yet discovered (0.000015; Brandon et al. 1999). These data will be discussed further in the subsequent section on the production of HSE-Os signatures in mantle melts.

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A recent study of Eoarchaean chromitites from south-west Greenland found 186Os/188Os data proposed to reflect mantle melt depletion events in Earth’s earliest history, during the Hadean at approximately 4.1 Ga and possibly as old as 4.36 Ga (Coggon et al. 2013). In so doing, Coggon et al. (2013) also inferred that the late veneer must have occurred prior to this time, consistent with the message of an ‘early’ late veneer from studies of basaltic meteorites from different parent bodies (Dale et al. 2012a).

The role of recycled oceanic lithosphere in producing HSE and Os isotope signatures in magmas At least part of the compositional variability observed in mantle melts at Earth’s surface is derived from heterogeneity in the mantle. The biggest single process by which such heterogeneous chemistry is generated must be that of recycling of oceanic lithosphere through subduction (e.g., Hofmann and White 1982). In addition, there are other processes, such as melt percolation within the mantle and lithosphere (e.g., Halliday et al. 1995) that potentially play an important role in producing the variety of magma compositions that we observe at Earth’s surface. Many instances of melt percolation may ultimately be sourced from enriched recycled material, but this is not a requirement in producing variations in fertility in the mantle. Here, we focus on the composition of recycled ultramafic and mafic lithosphere within the mantle, and its impact within the source regions of oceanic magmas. Oceanic alteration. Prior to subduction, the oceanic lithosphere gains variable amounts of water and trace elements during seawater interaction or hydrothermal alteration, resulting in the formation of serpentine minerals, at the expense of olivine. This alteration may, in more extreme cases, be accompanied by elevated 187Os/188Os and the loss of Os relative to the other IPGE (see abyssal peridotite section), but typically, abyssal peridotites retain mantle-like HSE proportions and 187Os/188Os ratios. Regardless of the precise HSE signature, serpentinization permits water transport deep into subduction zones and beyond into the deep mantle. Together with the hydrous mafic crust, this provides fluxes of fluids from the downgoing slab into the mantle wedge at a range of depths, as well as retention of water beyond the supra-subduction setting. The potential for the slab to transport water beyond the zone of sub-arc melting is likely to be important for promoting small-degree hydrous melting in the mantle, which may have an impact on HSE through refertilization processes. The impact of subduction zone processes on HSE in convergent margin magmas and recycled oceanic lithosphere. Fluxes into the mantle wedge produce two effects which have a bearing on HSE behavior and Os isotope composition. First, as discussed above, radiogenic Os may, in certain cases (Brandon et al. 1996; Becker et al. 2004), be transferred from the slab into the mantle wedge and then transferred by melts into arc crust and supra-subduction oceanic crust, sampled by ophiolites. Second, fluid addition will promote hydrous melting, allowing otherwise refractory mantle domains to partially melt and permitting melting of the mantle at temperatures below those of the normal geothermal regime. The evidence for a radiogenic Os flux to arc magma sources is equivocal, due to the difficulty in knowing the precursor 187Os/188Os of the mantle source and other potential sources of radiogenic Os such as arc crust. Nevertheless, the ophiolite record provides a firmer basis for this contention. An additional HSE flux is the loss of Re from metabasic rocks during dehydration (~ 50–60%; Becker 2000; Dale et al. 2007), and likely enrichment of Re in the mantle wedge (Sun et al. 2003a,b). This flux could contribute, over time, to radiogenic 187Os in the mantle wedge and also has implications for the composition of recycled crust which are outlined below. Other HSE may also be mobilized (McInnes et al. 1999; Kepezhinskas et al. 2002; Dale et al. 2009a), but whether the magnitude of flux is sufficient to produce a measurable effect in supra-subduction zone magmas is doubtful, given the relatively high concentrations of these elements in the mantle.

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Melting of refractory domains increases the likelihood of sulfide exhaustion, which, under most circumstances, would reduce the compatibility of all HSE, resulting in less fractionated HSE patterns such as those seen in picrites and komatiites (e.g., Puchtel and Humayun 2000). In the Tonga Arc, however, the relative proportions of the HSE are amongst the most fractionated for mantle melts (Dale et al. 2012b), with extreme Pt/Os approaching 15. This fractionation may be caused by increased HSE-rich phase stability during lower temperature hydrous melting (e.g., laurite stable up to 1275 °C; Brenan and Andrews 2001) and/or the promotion of chromitite formation by interaction between hydrous melts and refractory mantle (Dale et al. 2012b). Chromitite formation during melt–rock reaction in the mantle is expected to fractionate HSE significantly, sequestering IPGE in PGM and producing a melt with high (Re + Au + PPGE)/IPGE (see Ophiolite sections). The role of recycled lithosphere in producing HSE-Os signatures in convecting mantle melts. Many previous attempts have been made to model the effects of recycling oceanic lithosphere, particularly the mafic crustal portion (e.g., Roy-Barman et al. 1996; Brandon et al. 1999, 2007; Becker 2000; Dale et al. 2009b; Day et al. 2009). While we recognize the importance of quantitatively assessing whether a particular process is possible or likely, given the numerous previous attempts and the dependency on the parameters chosen, here we direct the reader to those previous studies and we instead choose to focus on the record of pyroxenites in the mantle, as direct recorders of enriched, hybridized lithologies. Of course, it is important to bear in mind that the sampled pyroxenite database is still relatively small (62 samples with HSE and/or Os isotope data collated in Fig. 25) and thus it is difficult to relate this to the mantle as a whole. That said, the processes identified are broadly applicable. Both eclogitic and pyroxenitic enriched lithologies are present in the mantle. Eclogites represent unequivocal crustal materials, sampled as xenoliths in intraplate volcanic settings, which retain much of their crustal geochemical signature, albeit modified by subduction processing. The term ‘pyroxenite’ covers a complex array of lithologies and petrogenetic pathways that are beyond the scope of this chapter (see Lambart et al. 2013). In simple terms, pyroxenites are variably hybridized lithologies produced during reaction of peridotite with silica-saturated melt derived from an enriched lithology such as eclogite (or possibly also derived from small-degree melting of peridotite). Reaction with a silica-undersaturated, olivine-saturated melt would instead produce dunite, so depending on the exact mode of formation of particular dunites (some dunites might be cumulates), they may also carry an enriched Os signature, as seen in the ‘convergent margin ophiolite’ section. Unlike eclogites, pyroxenites form a significant part of mantle tectonites, constituting between 1 and 9% of the Beni Bousera mantle tectonite massif (Pearson and Nowell 2004). These pyroxenites at Beni Bousera have been identified as having a recycled crust origin, on the basis of lithophile and stable isotopes. They typically have radiogenic 187Os/188Os ratios, even in samples that are Os-rich (> 2 ng/g). Pyroxenites and peridotites from the Totalp ultramafic massif, Swiss Alps, preserve a record of refertilization of peridotites by both melt percolation from the pyroxenites and from mechanical stretching and thinning of websterite layers (van Acken et al. 2008). The pyroxenites are strongly enriched in 187Os (187Os/188Os: 0.122–0.866; main range: 0.13–0.16) and in Re, whereas peridotites have a broadly chondritic average gOs value. It is noted, therefore, that refertilization does not completely homogenize Os isotopes, at least not on a small scale, but isotopic differences are reduced due to reaction of pyroxenite melt with peridotite. A compilation of ultramafic mantle samples, in terms of Pt/Os and Re/Os ratios, is presented in Figure 25. Pyroxenites form a distinct group at elevated Re/Os and Pt/Os ratio, relative to peridotites. The degree of this enrichment of Re is, in itself, consistent with a partially pyroxenite source for some mantle melts with radiogenic Os over a period of ingrowth of 1 Ga or more. Actual measured 187Os/188Os for global pyroxenites, excluding

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Common range of PGM and other ores

Required for observed 186 Os-187Os enrichment

(Typically Pt/Os 8. However, many rocks with elevated Pt/Os also possess elevated Re/Os which evolves to much higher 187Os/188Os ratios than observed in intraplate magmas with enriched 186Os. Therefore, rocks with Pt/Os, Re/Os and Pt/Re all greater than the PM are of particular interest for the generation of coupled enrichments of 186Os and 187 Os, but such rocks are a very minor proportion of the current mantle database (Fig. 25).

This difficulty in generating radiogenic 186Os, without also producing enrichments in Os beyond those observed, led Brandon et al. (1998), after Walker et al. (1995), to propose a role for transfer of an outer core Os signature into the plume source of some high-degree melts in intraplate settings. Twenty years later, this remains a possible scenario, despite the alternative mechanisms proposed that are outlined here. The core-mantle interaction model does, however, require an early onset of inner core solidification (by 2.5 Ga, and earlier for 2.8 Ga Kostomuksha komatiites; Puchtel et al. 2005) in order to allow sufficient time for ingrowth to produce enrichments in 186Os and 187Os in the predicted high (Pt–Re)/Os outer core. A more complete discussion of the core-mantle interaction debate can be found in Brandon and Walker (2005) and Lassiter (2006). 187

Since the emergence of the core-mantle interaction theory, several other possible sources of radiogenic 186Os have been proposed (e.g., Baker and Jensen 2004; Luguet et al. 2008b), though no proposed mechanism is completely convincing. The modification of pyroxenites, refertilization of peridotites and accompanying sulfide removal and/or metasomatism is the most likely alternative to core-mantle interaction (Luguet et al. 2008; Marchesi et al. 2014), but suitable Pt/Os and Re/Os ratios in the current mantle database are the exception, rather than the rule (Fig. 25). One further, more complex, possibility is that signatures may be combined from separate mantle components each with either high Pt/Os or high Re/Os, but not both. As outlined in a previous section, extreme Pt/Os fractionation exists on a variety of scales in Earth’s mantle, particularly during the formation of PGM. What is not yet clear is the fate of such PGM during mantle convection and whether there is sufficient separation and sampling of particular PGM compositions to produce specific signatures in mantle melts. In summary, processes exist in Earth’s mantle that can account for the enrichments observed in intraplate magmas, but currently they appear to be rare.

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The relationship between abyssal peridotites and MORB: an osmium isotope perspective One major debate in the field of HSE chemistry, and a key issue for mantle geology as a whole, is the extent to which abyssal peridotites represent the mantle residues of partial melting at oceanic spreading centres. Osmium isotopes have been a key part of this debate, but the evidence is complex. Early analyses identified a large range of 187Os/188Os compositions in abyssal peridotites, ranging from sub-chondritic to significantly supra-chondritic (see abyssal peridotite section above). The elevated signatures were largely attributed to seawater interaction. After taking into account this process, the remaining abyssal peridotite data appeared to be far less radiogenic than data for mid-ocean ridge basalts, thus casting doubt on a genetic link between abyssal peridotites and MORB. Since that time, two important findings have been made which reduce this discrepancy.

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Firstly, was the discovery of interstitial sulfides of magmatic origin possessing radiogenic, supra-chondritic 187Os/188Os ratios (Alard et al. 2005), together with non-chondritic PGE ratios (Alard et al. 2000). A preferential contribution from these interstitial sulfides to a partial melt, compared with the contribution from ancient, unradiogenic sulfides enclosed within silicates, could account for the more radiogenic signatures of MORB and other partial melts of the oceanic mantle, compared with those recorded in bulk-rock abyssal peridotites. Secondly, but of at least equal importance, was the finding that the Os isotope compositions of MORB (see Gannoun et al. 2016, this volume) were less radiogenic than previous thought. In particular, the range of 187Os/188Os ratios in MORB glasses was found to be considerably less (0.126–0.148) than previous findings (e.g., Schiano et al. 1997), with a lower mean of 0.133 ± 0.009, in part due to an analytical artefact in the original data (Gannoun et al. 2007). This mean value, while reduced, remains in excess of typical values for abyssal peridotites (187Os/188Os: 0.118–0.130). However, it was also found that the constituent phases of basalts had variable 187 Os/188Os due to (i) ingrowth over poorly-constrained periods of time since emplacement (Gannoun et al. 2004), and (ii) the timing of crystallization of different phases with respect to the evolution of the melt and its interaction with seawater-modified crust (Gannoun et al. 2007). Most notably, the latter manifests itself in significantly less radiogenic Os isotope compositions in early-formed relatively Os-rich sulfides compared with their (Os-poor) host glasses. In some cases there is a difference of ~ 0.015 in the 187Os/188Os of glasses and corresponding sulfides (e.g., glasses: 0.1383 and 0.1479; sulfides: 0.1249 and 0.1308, respectively), with the sulfides falling in the range 0.1236–0.1310, largely equivalent to the range seen in abyssal peridotites. Moreover, a negative covariation of 187Os/188Os and Os content in MORB sulfides might indicate that MORB sulfides are also affected by interaction with a radiogenic contaminant, casting doubt on the significance of more radiogenic data for Os-poor sulfides. Although sulfides included within silicates in abyssal peridotites (and other mantle tectonites) are known to possess even lower 187Os/188Os than bulk-rock samples (~ 0.114; Harvey et al. 2006)—and are therefore also lower than estimates of primitive MORB—such shielded sulfides likely contribute little to moderate degree partial melts relevant for MORB genesis. Therefore, in conclusion, not only has the ‘gap’ in composition between abyssal peridotites and MORB been largely bridged by radiogenic interstitial sulfides, but it seems likely that the gap is minor or non-existent when the most primitive parts of the MORB system are analyzed.

Interpretation of Re–Os model ages Model ages, whereby the isotope ratio of a sample is compared to the evolution of a reference frame such as average chondrite compositions, have been extensively used in geochemistry to give melt depletion ages in systems where recent mobility of elements has obscured any isochronous isotope systematics. The Re–Os system has been of particular use in this regard, due to the contrasting behavior of Re and Os which can result in, for high degree melts, effective Re removal from the source, while Os remains present in high enough abundances (several ng/g) to provide a degree of robustness against alteration and contamination. For Os, the measured 187Os/188Os ratio of a sample (or, for xenoliths, the ratio calculated at the time of the host eruption) is compared to the evolution curve of the mantle (commonly either a chondrite reference or the primitive mantle estimate). For Re depletion ages (TRD) it is assumed that the residue is completely depleted in Re after partial melting and, thus, there is no further ingrowth of 187Os. The advantage of this method is that it provides a relatively robust guide to the long-term evolution of the sample, due to the generally conservative behavior of Os, without the difficulties induced by recent Re addition or loss. In reality, however, only in high degree melting events is complete Re removal attained and in many cases the TRD age merely provides a minimum age. An alternative type of model age uses the measured Re/Os ratio to calculate the time when the 187Os/188Os of the sample intersected that of the reference frame (TMA or TRe–Os). In theory, this can provide a more accurate age, but it suffers the same sensitivity to Re mobility as do attempts to identify Re–Os isochron relationships.

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Numerous caveats and potential pitfalls of model age determinations have now been recognized and the reliability and interpretation of Re–Os model ages in peridotites was the subject of a comprehensive review by Rudnick and Walker (2009). Here, we summarize the main issues surrounding such model ages, in the context of the processes and tectonic settings discussed in this chapter. Perhaps the most obvious issue encountered has already been mentioned above—that of the degree of depletion of Re. Rudnick and Walker (2009) demonstrate that mantle melting at 3.5 Ga to form a basaltic melt would result in vastly different age estimates from TMA and TRD methods: the TMA age for the residue would be 3.5 Ga, because the Re/Os ratio of the residue is used to back-calculate the isotope evolution of the sample, whereas the assumption of complete Re depletion in the case of a TRD age would produce an age of just over 1 Ga. Clearly at this level of depletion TRD ages are not useful and they only become more valuable when Re removal is close to complete (probably a boninitic or komatiitic melt depletion event). Alternatives to isochron ages and Re depletion ages have been used to gain age information for sample suites where, respectively, Re mobility is suspected or Re removal was not complete. An element of similar compatibility to Re, but less mobile, such as Al2O3, can be used as a proxy for Re on an isochron diagram (Reisberg and Lorand 1995; see earlier). Although there is sometimes much scatter on such plots, they appear to be broadly robust. For large datasets of > 50 samples, but preferably more, probability density function plots provide a means to identify common apparent depletion ages, which lends weight to an argument for those ages having age significance. For instance, a range of 187Os/188Os ratios could be produced by variable degrees of depletion or by the same degree of depletion at different times. The identification of peaks on probability plots might indicate discrete times of melt depletion (perhaps partially obscured by variable depletion, preservation issues and/or inheritance) rather than a more continuous spectrum of compositions which might be expected from a suite of variably depleted samples. There is significant inherent uncertainty with any TRD age, because they are based on a model evolution curve. There are two aspects to this issue: (i) it is known that Earth’s mantle has broadly chondritic proportions of the HSE, but it is not known which chondrite group – if indeed any in the global collection – supplied Earth’s HSE or whether there was any fractionation of HSE during core formation. Models to account for the apparently suprachondritic Ru/Ir and Pd/Ir ratios of the PM (Becker et al. 2006; Walker 2009; Fischer-Gödde et al. 2011) may also have implications for the Re–Os isotope evolution of the PM. The choice of type of chondrite or PM estimate to use for the model evolution can result in an age variation of nearly 200 Ma for a 187Os/188Os of ~ 0.124, decreasing with increased age to an uncertainty of ~ 100 Ma at around 2 Ga (187Os/188Os = 0.114). (ii) As with lithophile isotope systems (e.g., Sm–Nd) a choice has to be made whether to use a primitive or depleted mantle reference frame. This can make an even more significant difference to the age given that the estimated 187Os/188Os of the primitive mantle is 0.1296, whereas an ‘average’ depleted mantle composition might be somewhere between 0.1245 and 0.128, depending on whether the average for abyssal peridotites or a combination of chromitites, PGM and high-degree mantle melts is used (Walker et al. 2002b; Pearson et al. 2007; Dale et al. 2009b). This also illustrates the problem of inheritance, which relates to the large degree of Os isotope heterogeneity observed in the convecting mantle and is amongst the most important considerations. This effectively means that for small datasets without additional information there is little way of knowing whether an apparent old age reflects a significant ancient melt depletion event in the context of its tectonic setting, or whether the measured 187Os/188Os is a composite of that event superimposed on an already depleted (or enriched) Os signature. For this reason, larger datasets obviously produce more robust age estimates and plots displaying probability can be used to identify ‘significant’ common ages or ‘peaks’ (Pearson et al. 2007; Rudge 2008).

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So far, we have made no mention of potential petrological pitfalls for model ages. These encompass serpentinization, sulfide breakdown, refertilization and melt–rock reaction (Rudnick and Walker 2009). Serpentinization, as discussed in an earlier section, does not typically affect Os isotope systematics except in extreme cases, which can easily be avoided when selecting samples with which to gain age information. Sulfide breakdown is known to occur in mantle xenoliths, due to interaction with the host melt. This commonly results in Os loss which could potentially impact upon the model age if 187Os/188Os is variable between different host phases, and which also leaves the sample more susceptible to contamination and alteration. Depending on the tectonic setting, some processes may or may not impact on model ages. For instance, melt–rock reaction in the convecting mantle is commonly associated with melting, and is therefore effectively zero age with respect to melting and should not normally affect the model age recorded for that melting event. Such melt–rock reaction also usually produces discordant samples on an 187Os/188Os–Al2O3 diagram, and can thus be identified and avoided for the purposes of dating. Conversely, processes of melt percolation and reaction in the continental lithosphere may occur long after the melt depletion episode of interest and this has the potential to obscure the true age (Rudnick and Walker 2009). These issues mean that samples with the lowest 187Os/188Os give the most reliable ages, but they too may still have experienced radiogenic Os input. The extent to which this process affects ages depends on the amount of addition of sulfide, and the Os isotope composition and concentrations of those sulfides. Such sulfides are typically poorer in Os than enclosed sulfides so significant additions of sulfide may be required to significantly affect the age. Although the processes of metasomatism and refertilization can have a significant effect on model ages, sometimes leading to recent TRD ages or “future” TMA ages, in some cases these processes can be traced using HSE behavior. For example, it has been recognized, in the cratonic setting, that the oldest TRD ages for a suite of samples are associated with the lowest Pd/Ir ratios, reflecting the most pristine and severe melt depletion signatures (Pearson et al. 2004). Recently, the Se/Te ratio has also been combined with Pd/Ir, in order to further understand the effects of metasomatic sulfide addition on model ages and place limits on the levels of addition that can occur before the model age may no longer be reliable (Luguet et al. 2015). In summary, there are numerous potential pitfalls and limitations for Re–Os model age determinations but, in the absence of isochron dating, the system remains amongst the most useful for providing the ages of melt depletion of the mantle.

ACKNOWLEDGMENTS We thank Chris Ballhaus, Al Brandon, James Brenan, Kevin Burton, Rick Carlson, James Day, Mario Fischer-Gödde, Mouhcine Gannoun, Timo Gawronski, Jason Harvey, Akira Ishikawa, Yogita Kadlag, John Lassiter, Ambre Luguet, Jean-Pierre Lorand, Claudio Marchesi, Graham Pearson, Igor Puchtel, Dave Rubie, Steve Shirey, David van Acken, Richard Walker and Zaicong Wang for valuable insight and discussions over the years. Thanks to Jason Harvey, Chuan-Zhou Liu, Wendy Nelson and Jessica Warren for helpful reviews of the manuscript.

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J Petrol 37:3–21 Zhou MF, Sun M, Keays RR, Kerrich RW (1998) Controls on platinum-group elemental distributions of podiform chromitites: A case study of high-Cr and high-Al chromitites from Chinese orogenic belts. Geochim Cosmochim Acta 62:677–688 Zhou M-F, Robinson PT, Su B-X, Gao J-F, Li J-W, Yang J-S, Malpas J (2014) Compositions of chromite, associated minerals, and parental magmas of podiform chromite deposits: The role of slab contamination of asthenospheric melts in suprasubduction zone environments. Gondwana Res 26, 262–283 Zhou MF, Yumul GP, Malpas J, Sun M (2000) Comparative study of platinum-group elements in the Coto and Acoje blocks of the Zambales Ophiolite Complex, Philippines. Isl Arc 9:556–564

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Reviews in Mineralogy & Geochemistry Vol. 81 pp. 441-488, 2016 Copyright © Mineralogical Society of America

Chalcophile and Siderophile Elements in Mantle Rocks: Trace Elements Controlled By Trace Minerals Jean-Pierre Lorand Laboratoire de Planétologie et Géodynamique de Nantes UMR 6112, Université de Nantes Faculté des Sciences, BP 92208 2 rue de la Houssinière, 44322 Nantes Cedex France

Ambre Luguet Rheinische Steinmann Institut für Geologie, Mineralogie und Paläontologie Poppelsdorfer Schloss Friedrich Wilhelms Universität Bonn 53115 Bonn Germany INTRODUCTION Since V.M. Goldschmidt’s pioneering work, chalcophile elements have been identified as showing the greatest affinity for sulfur. Goldschmidt (1954) attempted to chart the distribution of these elements between the silicate (lithophiles), metal (siderophiles) and sulfide (chalcophiles) portions of meteorites by using sulfidation curves of metal 2M + S2 ⇌ 2 MS. Using a similar approach, Arculus and Delano (1981) suggested the following decreasing order of chalcophilic behavior: Ga > Cu > Mo >Fe > Ni > W > Co > Sn > Pb > Ag > Pt > Ir > Os > Sb > Ge > Re. Clearly such classifications are not suitable for discussing mantle chalcophiles. Siderophile and chalcophile elements have intermediate electronegativities and tend to form covalent or metallic bonds that are predominant in sulfide structures. Most elements that are siderophile are usually also somewhat chalcophile and vice versa. For example, highly siderophile elements (HSE) such as platinumgroup elements (PGEs: Os, Ir, Ru, Rh, Pt, Pd), Re and Au are strongly concentrated in the sulfide phases, compared to nominally chalcophile elements (e.g., Pb, Ga, Ni) in terms of mass balance. Highly siderophile elements are assumed to be controlled by sulfide phases in the source of most mantle rocks and mantle-derived melts examined so far, because the uppermost mantle is not saturated with respect to Fe–Ni metal (Rohrbach et al. 2007). For this reason, the broad definition of chalcophile elements in the mantle should include all of the elements that are collected into sulfides, i.e., including highly siderophile elements (HSE), i.e., the platinum-group elements (PGE), Re, Au, Ag and the chalcogenides Se and Te. One way of sorting chalcophiles is by considering their sulfide melt/silicate melt partitioning behavior (D sulfide melt/ silicate melt = the weight fraction of metal in sulfide melt/ the weight fraction of metal in silicate melt). Empirically and experimentally determined D sulfide melt/ silicate melt increase from 15 to 80 for Co, 300 to 1000 for Ni, 900 to 1,400 for Cu, Se, to 104 for Te, Re and Ag, and > 105 for the PGE (e.g., Peach et al. 1990; Fleet et al. 1991, 1999; Sattari et al. 2002; Brenan et al. 2016, this volume). For the sake of this paper, chalcophile, and siderophile elements will be referred to as low-D (D sulfide melt/ silicate melt < 103; i.e., Ni, Co, S, Se), middle-D (D sulfide melt/ silicate melt > 103; i.e., Cu, Te, Au, Ag, Re, Te), and high-D elements (D sulfide melt/ silicate melt > 104; , PGEs-Os, Ir, Ru, Rh, Pt, Pd). 1529-6466/16/0081-0008$5.00

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Chalcophile and siderophile elements belong to a very specific class of trace elements in mantle rocks because they are controlled by a wide range of minor or trace minerals that are not necessarily familiar to mantle petrologists. Due to petrogenetic processing of the upper mantle, major (Fe), minor (Ni, Co), trace (Cu, S) or ultra-trace elements (Se, Te, Au, PGE, Re) may be incorporated in minor sulfides (0.01 to 0.1 vol.%) as both metal cation- or anion-replacing elements. They may also form their own phases at the sub-micron to nanometer scale depending on ambient conditions. Thus, in contrast to lithophile elements, chalcophile and siderophile elements may occur in very different host minerals as a function of sulfur-saturation, redox conditions, pressure, fugacity of sulfur and melt compositions. The aim of the present chapter is to explore in detail the relationships between whole-rock contents of siderophile–chalcophile elements and sulfur and the abundance, identity and distribution of phases that contain these elements at all scales, from the bulk-rock scale to nanometric minerals. Characteristic suites of phase assemblages and their chemical fingerprints will be reviewed in the light of currently debated petrogenetic processes for the upper mantle: adiabatic partial melting, channelized melt flow, open system melting, metasomatism and magmatic percolation of the lithospheric mantle. Before reviewing whole-rock concentrations, it is necessary to recall the fundamentals of mantle rocks and their provenance, which phases are present in the convecting mantle before any petrogenetic processing and their fate during subsolidus cooling of mantle samples.

BACKGROUND Rocks from the Earth’s mantle occur in a wide range of geologic settings (1) as peridotite massifs, now referred to as orogenic peridotites; (2) as spinel-(rarely garnet) -peridotite xenoliths from alkali basalts and abyssal peridotites. Orogenic peridotite massifs and basalt-hosted peridotite are probes of the shallow (< 100 km) post-Archean, off-cratonic sub-continental mantle lithosphere (SCLM). These peridotites are moderately depleted lherzolites (mostly from the spinel lherzolite facies) along with subordinate amounts of refractory harzburgites and/or dunites (see Bodinier and Godard 2014; Pearson et al. 2014). They are commonly associated with pyroxenites (Cr-diopside, Al-rich augite suites), minor hornblendites and disaggregated xenocrysts of these rocks. Garnet peridotite xenoliths carried by kimberlites provide a window on deeper levels of the SCLM beneath Archean cratons and platforms. Orogenic peridotite massifs, tectonically emplaced by deep transcurrent faults in Phanerozoïc mountain belts, also sample fossil suboceanic lithosphere (Bodinier and Godard 2003). The mantle sections in ophiolites and abyssal peridotites, dredged or drilled from the ocean floor, have sampled the oceanic mantle beneath fossil or present spreading centers. More detailed informations on cratonic peridotites, peridotites derived from oceanic settings, and SCLM can be found in Aulbach et al. (2016, this volume), Becker and Dale (2016, this volume), and Luguet and Reisberg (2016, this volume), respectively. Each of those mantle samples brought to the Earth’s surface has experienced a cycle of petrogenetic processes starting with partial melting (resulting from adiabatic decompression of upwelling mantle and/or fluid input), followed by storage in the rigid lithosphere within which a wide variety of magmatic alteration are experienced. These may include melt intrusions producing pyroxenites, porous flow percolation at variable melt/rock ratios leading to re-enrichment/depletion of the rigid lithosphere in basalt-forming major elements  (i.e., Al2O3, FeO, CaO), or fluid flow that enriches mainly trace elements, i.e., mantle metasomatism (Pearson et al. 2014), and often ending with hydrothermal alteration in the example of oceanic peridotites. Even abyssal peridotites that have long been considered to be residual solids of single-stage near-fractional adiabiatic melting episodes of the mantle beneath spreading ridges bear evidence for this post-partial melting petrogenetic processing (Seyler et al. 2007; Bodinier and Godard 2014).

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This chapter is primarily based on integrated studies that have analyzed bulk rocks and their constituent minerals, using in-situ analyses such as scanning electron microscope (SEM) and laser-ablation-inductively coupled plasma mass spectrometry (LA-ICPMS), coupled with optical microscopy. To date, there are few studies that have made such an effort to unravel which minerals control the siderophile/chalcophile element budget of the upper mantle. This paucity can be explained by the challenges of sample preparation required for fine-scale preservation of trace phases, by difficulty of obtaining representative sampling and in situ analytical limitations. Most, if not all integrated studies, are from a few basalt-hosted xenolith localities (Massif Central, France; Bohemian Massif, Czech Republic, Kilbourne Hole, SW USA, Hawaii and Eastern Asia, China, Korea), orogenic peridotites (Western Europe), and abyssal peridotites from the Mid-Atlantic Ridge. These rocks encompass fertile mantle samples where sulfides play a significant rôle in controlling siderophile and chalcophile element budgets and where the effects of crustal alteration are minimized. Elsewhere, literature data reports only whole-rock abundances of PGE with either S or Cu data, occasionally Se and, more uncommonly, Te (see Lorand et al. 2013 and reference therein; Wang et al. 2013a). Inferences on mineral carriers can be made only by examining relationships between whole-rock abundances of highly siderophiles and chalcophiles and their constituent minerals. Many previous studies used Os isotopes for age determination and geodynamical inferences regarding mantle samples (See Aulbach et al. 2016, this volume; Becker and Dale 2016, this volume; Harvey et al. 2016, this volume; Luguet and Reisberg 2016; this volume). The reader is referred to these chapters for further information regarding geochronological aspects of the HSE.

Sulfides in the upper mantle and mantle rocks Nomenclature regarding the naming of sulfides as it will be employed in the text below is specified here. Mantle sulfides are Cu–Fe–Ni sulfides or Base Metal sulfides (BMS), in line with terminology adopted for describing magmatic sulfide ore bodies that concentrate such phases to levels of economic interest (e.g., Naldrett 2005). All minerals of the Cu–Fe–Ni–S system, with the exception of native sulfur, have been reported from mantle rocks but only a handful are assumed to be stable in the lithospheric upper mantle (Fig. 1). Considering only phases stable above 900 °C (a reasonable temperature for the continental lithospheric mantle in high heat-flow areas, Pearson et al. 2014), BMS are, by decreasing order of melting point, Monosulfide Solid Solution (Mss—Fe1–x,Ni1-xS), T = 1190 °C at 1 bar), Bornite Solid Solution (Bn; Cu5FeS4 to Cu2S), T > 1070 °C at 1 bar) and Intermediate Solid Solution (Iss; CuFeS2-x–CuFe2S3-x; x = 0–1; T = 880 °C) of the Cu–Fe–S system (Craig and Kullerud 1969; Raghavan 2004; Fleet 2006). In specific areas of the mantle characterized by low heat flow (e.g., cratons) other sulfides, like the high-T polymorph of heazlewoodite (Ni3+xS and its Co analog, with a melting point of 862 °C; e.g., Aullbach et al. 2004) can theoretically form in the lithospheric mantle, as does pyrite (T > 810 °C for P = 0.4 GPa; Fleet 2006 and references therein). However, pyrite is not a mantle mineral, because it requires high-S activity conditions that are usually not reached in the mantle (e.g., Lorand and Alard 2011, and references therein). The major phases seen today in mantle rocks are pentlandite, chalcopyrite and pyrrhotitelike phases, along with occasional bornite occurrences. These minerals are usually referred to as primary, i.e., produced by closed-system subsolidus exsolution/solid-state reactions from the sulfide solid solutions that are stable in the mantle. Experimental studies on Ni–PGE ore deposits show that such reactions operate over a large temperature range, from 900 °C to less than 200 °C (Nalldrett 2005; Fleet 2006). The series of subsolidus re-equilibration reactions that lead to primary mantle sulfides were discussed in full in several papers (e.g., De Waal and Calk 1975; Lorand and Conquéré 1983; Dromgoole and Pasteris 1987; Szabó and Bodnar 1995; Guo et al. 1999; Lorand and Grégoire 2006 and references therein). These papers defined Mss phases in mantle rocks as homogeneous pyrrhotite-like phases containing more than ca. 5 wt.% Ni and free of pentlandite exsolution, at least at the highest magnification possible for optical microscocopy. Considerations

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Figure 1. BMS phases usually identified in mantle-derived rocks and phase relationships in the Cu– Fe–Ni–S system at 125 °C (Fleet 2006 and references therein). Abbreviations of mineral names: Mss: Monosulfide solid solution; Pn = pentlandite; Cp = chalcopyrite; Bo = bornite; Py = pyrite; Vs = vaesite; Pd = polydymite; Mi = millerite; Gd = godlevskyite; Hz = heazlewoodite; Aw = awaruite; Fe = native iron; Mw = mackinawite; Tr = troilite; Hpo = hexagonal pyrrhotite; Mpo = monoclinic pyrrhotite; G = greigite; Cv = covellite; Cc = chalcocite; Dg = digenite; Cu = native copper. Mss1 and Mss 2 are the two compositional ranges of Mss below 300 °C (Fleet 2006 and references therein).

of metal/sulfur ratios (M/S) that spread over the whole range of pyrrhotite compositions, instead of clustering around the three end-members (FeS, Fe9S10, Fe7S8) also helped in characterizing mantle-derived Mss (e.g., Lorand and Conquéré 1983; Guo et al. 1999). Monosulfide solid solution inclusions have been reported in diamonds, eclogite, pyroxenite, and peridotite xenoliths from alkali lavas. Peridotite-hosted Mss spans a large compositional range from 10 to 30 wt.% Ni, whereas eclogitic and pyroxenite-hosted Mss are less Ni-rich (< 10 wt.% Ni). At upper crustal temperatures (300 °C), the primary field of the Mss is interrupted by a solvus that splits it into distinct fields corresponding to Ni-poor (Mss1) and Ni-rich (Mss2) compositions (Fig. 1). These Mss phases have been reported to occur as sigmoidal intergrowths enclosed within silicates in xenocrysts, diamonds and xenoliths with worldwide provenance (Fig. 2A,C). One may surmise that the Mss was preserved by rapid cooling of these mantle samples. Once such inclusions are fractured, Mss are degraded into pyrrhotite-pentlandite intergrowths textures (Fig. 1). Not only the ascent rate, but also pore fluids may contribute to subsolidus breakdown of Mss into pentlandite by enhancing Ni–Fe diffusion and readjustment inside the pyrrhotite lattice (Lorand 1985, 1987b, 1989a). Independent experimental work on pentlandite exsolution has shown that other parameters (i.e., the initial metal deficiency, accumulation of crystalline defects inside the

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Figure 2. Photomicrographs of representative mantle-derived BMS. A: (Type-e) rounded Mss inclusions in olivine from a cratonic garnet peridotite xenolith (SEM-backscattered-electron image (BSE) image showing finely intergrown Ni-rich (greyish) and Ni-poor (darker) Mss phases; Lorand and Grégoire 2006) B: plane polariser reflected light microscopy: intergranular Cu–Ni-rich BMS from off-craton peridotite xenoliths (Lorand and Alard 2001); note the curved shape of this grain and the core-to-rim distribution of BMS: relict Mss containing thin pentlandite platelets along cleavage planes in the core is surrounded by an inner discontinous aggregate of pentlandite (light grey) and outer patches of Cu sulfides (isocubanite). This is the microtexture predicted by the solidii of the Cu–Fe–S phase diagram of Figure 16; C: euhedrally-shaped Mss inclusion in olivine from a cratonic garnet peridotite xenolith (plane polariser reflected light microscopy); twophase Mss are seen in the BSE image of the grain. D: highly altered Mss in a Massif Central spinel peridotite xenolith (France); only pentlandite (white) survived alteration (plane polariser reflected light microscopy).

Mss structure and its degree of Ni supersaturation) may interact in determining shape and the distribution of exsolved pentlandite (Etschmann et al. 2004). The fugacity of sulfur in the lithospheric mantle is oxygen fugacity-dependent through the sulfidation reaction of olivine into orthopyroxene and Mss. 2 Fe2SiO4 + S2 = 2 FeS +Fe2Si2O6 + O2 The calibration of this reaction by Eggler and Lorand (1993) as a sulfide barometer showed that Mss-dominated sulfide assemblages span a range of ƒS2 from 1–2 log units above to 1–2 log units below the quartz–fayalite–magnetite-–pyrrhotite reference buffer (FMQ–Po). Calculated ƒO2 is consistent with the ƒO2 range reported for the shallowest subcontinental mantle lithosphere (FMQ to FMQ – 2; Woodland et al. 1992). Other BMS assemblages reported in mantle rocks are representative of redox conditions outside this range. Figure 1 provides a list of secondary minerals resulting from exsolution, alteration, oxidation, sulfidation/desulfidation of primary sulfides, each characterizing a sequence of crustal alteration of primary (i.e., mantle-derived) sulfide assemblages. Secondary crustal processes affecting mantle rocks include metamorphism, and serpentinization for tectonically emplaced samples (especially for oceanic peridotite that have reacted with seawater-derived hydrothermal fluids on the ocean floor) and sulfide breakdown prior

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to, during, or after entrainment and eruption of the host basalt for xenoliths. Two major trends are clear: an overall desulfidation trend, closely related to serpentinization, and an overall oxidation trend related to weathering. Serpentinization affects mainly oceanic mantle peridotites at the ocean floor or during reactions during on-going subduction (ophiolites, abyssal peridotites; e.g., Alt et al. 1998), but also orogenic peridotites and some cratonic xenoliths that have been stored at T > 150 °C inside the upper crust. Serpentinization usually produces S-poor assemblages (troilite, mackinawite, Fe-rich pentlandite, awaruite, native copper, native iron ; Lorand 1985, 1987, 1988; Abrajano and Pasteris 1989; Alt and Shanks 1998; Luguet et al. 2003; Seyler et al. 2007). More Nirich silicate lithologies (dunites, harzburgites) lead to godlevskiite, heazlewoodite, Ni-pentlandite and Ni-rich awaruite (e.g., Lorand 1987a,b; Ackerman et al. 2009; Marchesi et al. 2013). Nirich sulfides are spatially associated with grain boundary serpentine. Detailed studies of mineral compositions can be found in the literature for Betico Rifean orogenic peridotites (Beni Bousera, Morocco and Ronda, Spain; Lorand 1985) and Bay-Of-Islands and Oman ophiolites (Lorand 1987a, 1988). Chemographic analyses of reducing conditions produced by serpentinization were performed on Zambales ophiolitic peridotites (Philippina; Abrajano and Pasteris 1989; see also Frost 1985). These studies established that the most suitable conditions for generating reducing conditions stabilizing native iron during retrogressive serpentinization are low water–rock ratios, olivine-rich lithologies, and restitic olivine preserved in the serpentinized assemblage. By contrast, supergene weathering becomes active when mantle rocks are exposed to oxygenated and aerated atmosphere or to groundwaters (or to sea-water alteration ; Luguet et al. 2003). Weathering generates metal-deficient sulfides (smythite Fe3S4 , greigite, FeNi9S11, violarite, FeNi2S4, polydymite, Ni3S4), and Ni-rich porous pyrite. Weathering is not uncommon during exposure of orogenic peridotites (e.g., Lorand 1989a,b) and basalt-hosted xenoliths (Lorand 1990). Peridotite xenoliths emplaced in continental basalts show advanced alteration of their BMS in the form Fe oxyhydroxides (Fig. 2D). The alteration sequence for Cu sulfides includes secondary bornite replacing chalcopyrite and covellite. Hydrothermal alteration can also modify mantle-derived BMS assemblage via sulfate/sulfide addition, generating euhedral pyrite and/or chalcopyrite, (Luguet et al. 2004; Alt and Shanks 1998) or replacing pyrrhotite or Mss by Ni-rich pyrite and smythite (Desborough and Czamanske 1973) in Robert Victor eclogites, South Africa (see also Gréault et al. 2013). Lherzolitic bodies contaminated by evaporitic sulfides show a large amount of Co-rich, high-temperature pyrite intergrown with pentlandite (Lorand and Alard 2011).

Abundance and phase control on chalcophile and siderophile elements in the fertile upper mantle Low D elements. Sulfur—Available modal estimates (by point counting) indicate that BMS modal proportions broadly increase from 0–0.1 vol.% in lherzolitic peridotites up to 1 vol.% in pyroxenites or highly metasomatized peridotites (e.g., Dromgoole and Pasteris 1987; Wang et al. 2009; Sen et al. 2010) . How much S is present in the present-day upper mantle is not easily determined by terrestrial accretion models because of the strong partitioning of S to the metallic core (Dreibus and Wanke 1996; Li and Agee 2001). Whether the S in the present day mantle is a restitic element from a high-pressure core-formation event or an input from late accretionary materials (the Late Veneer Hypothesis) is still debated (Labidi et al. 2013). Since McDonough and Sun (1995), the preferred (and now widely accepted) value for the hypothetical Primitive Upper Mantle (PUM—the geochemical reservoir representing the continental crust and upper mantle before differentiation of the continental crust; Bodinier and Godard 2014; Pearson et al. 2014 and references therein) is 250 µg.g−1 (Fig. 3). Orogenic lherzolites from Western Europe (e.g., Pyrénées, Lorand 1989c, 1991; Baldissero Balmuccia, Garuti et al. 1984; Wang et al. 2013a) and a handful of basalt-hosted xenolith suites (e.g., Damaping-Hannuoba; Trans-China orogenic belt; North China craton; Liu et al. 2010; the Kilbourne Hole maar (South Western USA; Morgan 1986) approach the S abundance of PUM, in addition to a few samples from the UM suite (Basalt Study Volcanic Project project; see details in Morgan 1986; Becker et al. 2006)

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and others from Montferrier, southern France (Alard et al. 2011) and eastern Australia (e.g., Mt Quincan; Handler and Bennett; 1999; Handler et al. 2005). Gakkel Ridge abyssal peridotites (recovered from an amagmatic ultra-slow spreading ridge) possess similar abundances of S (Liu et al. 2009). This value is slightly lower, but less well constrained than the present-day Depleted MORB-source Mantle (DMM, the convecting upper mantle that was depleted in incompatible trace elements by extraction of the crust; Bodinier and Godard 2014; Pearson et al. 2014 and references therein). Most oceanic peridotites, whether orogenic peridotites (e.g., Lanzo), abyssal peridotites from North Atlantic, or ophiolitic plagioclase lherzolites, suggest abundances closer to 150–200 µg.g−1 for the DMM (Lorand 1991; Lorand et al. 1993; Luguet et al. 2004) as suggested by Salters and Stracke (2004) and Nielsen et al. (2014). It is widely accepted that S in the convecting upper mantle is totally accounted for by BMS, i.e., BMS-free samples are almost devoid of S. Whether this element can enter silicates at mantle pressure and temperature and is exsolved during cooling and decompression (as suggested by Ionov et al. 1992) is still a matter of speculation, although recent partitioning study of Calegaro et al. (2014) shows that S can be sparingly soluble in cpx. Base Metal Sulfides occur as both enclosed (Type-e) and interstitial (Type-i) grains. Sulfur in orogenic peridotites and S-rich xenoliths is governed by abundant Type-i BMS. Samples of DMM show very similar Type-i BMS assemblages to continental orogenic peridotites (Lorand et al. 1993; Luguet et al. 2003, 2004), although these assemblages are more altered by serpentinization and hydrothermal alteration. Type-i BMS grains are ovoids, rounded to ellipsoidal droplets, 50 × 100 μm on average. Larger grains (up to 500 μm across) show polyhedral shapes with convex inward or convex outward grain boundaries (Fig. 4A,B). Intergranular sulfides show evidence of low dihedral angles suggesting

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wetting behavior at silicate triple junction (Fig. 4B). Blocky pentlandite is by far the most abundant phase and is compositionally homogeneous (80 ± 10 vol.%), followed by chalcopyrite, whereas pyrrhotite is subordinate (Garuti et al. 1984; Lorand 1989a; Lorand et al. 2008b; Liu et al. 2010). Liu et al. (2009) reported Cu-rich (up to 12 wt.% Cu) pentlandite + FeS + Ferich pentlandite associated with orthopyroxene–clinopyroxene–spinel (opx–cpx–sp) clusters in Gakkel ridge fertile abyssal peridotites. At mantle temperatures (e.g., 1000 °C), this BMS assemblage corresponds to a metal-rich Mss (atomic metal/sulfur ratio close to 1) coexisting with a sulfide melt enriched in Ni and Cu (Fig. 5). Binary plots vs. S provide information on the chalcophilic behavior of low-D elements. There is no correlation between low-D elements (Fe and Ni) with S (not shown). On a bulkrock scale, Ni and Fe exhibit lithophile behavior and are concentrated into mafic minerals. The amount of Fe and Ni stored inside BMS can be estimated from a simple mass balance combining the average BMS modal composition of the PUM-type samples (80 ± 10% Pn 20 ± 10% Cp), their putatite BMS modal content (0.07 ± 0.02 wt.% assuming 200–250 µg.g−1 S in the fertile upper mantle and 34 wt.% S in 100 wt.% BMS) and the Fe and Ni concentrations in those BMS ( 25–45 wt.%; Fig. 5). This amount ranges from ca. 10% for Ni (220 ± 80 vs. 1960 µg.g−1 for the PUM) to almost nothing for Fe (220 ± 80 µg.g−1 vs. 80,000 µg.g−1 Fe for the PUM; McDonough and Sun 1995; Bodinier and Godard 2014; Pearson et al. 2014).

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Figure 4. BSE images of intergranular (Type-i) BMS in lherzolites displaying a Primitive Upper Mantle (PUM) systematic of high-D and middle-D chalcophile and siderophile elements. A: polyhedral pentlandite grain displaying very low dihedral angles with matrix silicates B: polyhedral pentlandite–chalcopyrite intergrowth illustrating the high wetting capacity of Cu–Ni-rich sulfide melts in the mantle (Lherz; Pyrénées France). C: Type-i pentlandite adjacent to Al-spinel D: Enlargment of C showing a moncheite (Pt–Te (Bi)) lamella (bright grey) inside a cleavage plane of pentlandite (Lherz; Pyrénées France).

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Type-e sulfides, MC granular lherzolites (Lorand and Conquéré, 1983)

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Selenium has long been considered to be a non-metal proxy of S, entering mantle BMS as an anion replacing sulfur (e.g., Garuti et al. 1984; Morgan 1986). Empirically determined D sulfide melt/ silicate melt estimates are very similar for S and Se (3 × 102; Peach et al. 1990; Patten et al. 2013 and references therein). High-precision whole-rock analyses are now available, using either external calibration standards and ICPMS or thiol cotton fiber (TCF) separation, isotope

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dilution and ICPMS (König et al. 2012, 2014; Wang et al. 2013; Fig. 6). Selenium positively correlates with S, and the S/Se of PUM-type lherzolitic samples is close to chondritic, while the correlation intercepts the intersection of the x- and y- axis at close to 0. This Se vs. S correlation is scattered by weathering (basalt-hosted xenoliths) or hydrothermal alteration (oceanic peridotites). A positive correlation between whole-rock Se contents and BMS modal abundances (r = 0.75) was also reported by Lorand et al. (2003a). This strong affinity of Se for BMS is supported by in-situ analyses. Older proton-induced X-ray emission (PIXE) probe analyses detected 100–150 µg.g−1 Se in Quilin mantle xenoliths (Southeastern China; Guo et al. 1999; see also Hattori et al. 2002). For basalt-hosted non-cratonic lherzolite xenoliths from the Massif Central (France), Lorand and Alard (2001) reported 108 ± 21 µg.g−1 Se for 8 Mss inclusions analyzed by laser ablation microprobe (LAM)-ICPMS, and 160–230 µg.g−1 for isocubanite-rich inclusions. In Ligurian ophiolitic lherzolites (Italy), pentlandite shows rather constant Se concentrations (110 ± 10 µg.g−1) and S/Se = 3000 that match the whole-rock ratio (Luguet et al. 2004). Lorand and Alard (2010) discussed the distribution of Se in Pyrenean

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orogenic lherzolites from systematic LAM-ICPMS analyses of BMS and high-precision whole-rock analyses in more detail. The average Se content of Lherz pentlandites ranges between 127 and 160 µg.g−1 with occasional higher values for chalcopyrite (up to 268 µg.g−1); each sample was shown to be fairly homogeneous in terms of mass balance calculations that clearly identify BMS as the only Se carrier in the fertile mantle (see also Lorand et al. 2008b). No mineral containing Se as major element was detected in these rocks. Middle D elements. Copper is indisputably a chalcophile element (D values up to 2,000; Patten et al. 2013; Mungall and Brenan 2014 and references therein). It is intuitively treated as such in all the recent literature on mantle rocks, although no accurate mass balance has yet been published. Orogenic lherzolites and other PUM-type related S-rich xenoliths display a reasonably good positive between Cu and S (Fig. 7). Like the S vs. Se plot, the data in the Cu vs. S plot are scattered by weathering-related S loss for basalt-hosted xenoliths that spread toward much higher Cu/S, and by the precipitation of hydrothemal sulfides in highly serpentinized samples (oceanic and cratonic peridotites). Originally estimated at 25 µg.g−1 (McDonough and Sun (1995), the Cu value for the PUM was recently challenged by Liu et et al. (2014, 2015) and Fellow and Canil (2012), who suggested a slightly higher figure of 30 µg.g−1 from the Cu contents of basalts (MORB, OIB, and arc basalts). Orogenic peridotites and PUM-type related S-rich xenoliths are distinctly enriched in Cu, regardless of which PUM estimate is chosen (Cu/S = 0.15 ± 0.05; Fig. 7). The partitioning of Cu into silicates has recently been addressed in an experimental study performed under sulfur-free conditions involving synthetic compositions (Liu et al. 2014). The 60

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authors of that study suggested that Cu is not equally distributed in all non-sulfide phases but rather concentrated in the modally dominant phases olivine and orthopyroxene. The Cu abundances in upper mantle spinels or garnets are thought to be less than 1 µg.g−1. However, two lines of evidence suggest that Cu is quantitatively hosted inside BMS in mantle peridotites: the correlation in Figure 7 intercepts the intersection of the x- and y-axis at close to 0: the most depleted harzburgites do not contain more than 5 µg.g−1. Cu. If recalculated in terms of BMS composition, a whole-rock Cu/S of ca. 0.15 corresponds to a BMS modal composition of ca. 15% Cp + 85% Pn, asssuming that pentlandite does not accommodate Cu. These modal proportions are in excellent agreement with those estimated from fertile orogenic peridotites (Lorand 1989).

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Gold in mantle rocks is also treated as a chalcophile element, showing a marked preference for metal-rich Cu-sulfide melts compared to monosulfides, in agreement with the very similar electronic structure of both elements (DMss-sulf. melt < 0.2, e.g., Li et al. 1996; Mungall et al. 2005). Gold has been analysed infrequently as it is mono-isotopic and not suitable for analysis using isotope dilution. Orogenic peridotites and PUM-like S-rich xenoliths produce a broadly positive correlation between Au and S (Fig. 8). This is strongly scattered by the well known mobility of Au in supercritical fluids involved in metamorphism, hydrothermal alteration, serpentinization and sea-floor alteration of mantle peridotites. Strong gold enrichments were reported for serpentinized abyssal peridotites (Luguet et al. 2003, 2004; Marchesi et al. 2013) and mantle peridotite bodies contaminated by near-surface H2O–CO2 fluids (Lorand 8 Cratonic xenoliths et al. 1989, 1999). Despite evidence Off-craton xenoliths 7 for chalcophile behavior, mantle Orogenic lherzolites lherzolites highlight a large deficit (as Ophiolites 6 large as that of Pt (see next section); Abyssal peridotites crustal contamination, Luguet et al. 2001) in the whole-rock 5 metasomatism budget of Au, which is not accounted 4 for by BMS. Time-resolved spectra collected during LAM-ICPMS analysis 3 of BMS reveal Au concentration 2 spikes corresponding to sub-μm-sized nuggets intercepted by the laser beam 1 (Fig. 9). Such spikes provide evidence 0 for discrete micro- to nanometric 0 50 100 150 200 250 300 350 Fig. 8gold particles inside BMS (cf. Luguet 2 et al. 2001, 2004). Investigations sea-floor of orogenic and ophiolitic mantle hydrothermalism 1.6 peridotites by SEM also identified micrometric native Au, Au tellurides and various Au–Cu and Au–Ag alloys 1.2 PUM as microphases attached to (or remote from) BMS (Ohnenstetter 1992; 0.8 Lorand et al. 2010). This also explains the poorly reproducible whole-rock Au contents determined from small-sized 0.4 subsample peridotite aliquants (cf. Fischer-Gödde et al. 2011; Marchesi et 0 0 50 100 150 200 250 300 350 al. 2013; Wang et al. 2013). S (µg.g-1)

Figure 8. Plot of Au vs. S (whole-rock data). Note the overall Au enrichment of oceanic peridotites and some orogenic lherzolites (source of data: Lorand et al. 2013).

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because Au, like Pd, is soluble in isocubanite, but not in low-temperature chalcopyrite (Luguet et al. 2001; Lorand and Alard 2001). Moreover, gold which occurs in low valence state (0, +1) does not enter the octahedral site of pentlandite, unlike Ag which may form argentian pentlandite. Ferraris and Lorand (2015) recently examined the origin of that missing (“invisible”) gold fraction with a Transmission Electron Microscope (TEM). They identified nanometric inclusions of novodneprite (AuPb3) and anyuiite [Au(Pb,Sb)2] together with nanometric clusters of Au particles enclosed in olivine from a Lherz lherzolite. These authors proposed that the novodneprite and anyuiite inclusions are the result of subsolidus recrystallization of the Pyrenean lherzolites with olivine grain growth that accidentally trapped intergranular component.

Tellurium is a large-size semi-metal forming soft ligands with transition metals (Ni, Pt, Pd, Au). The most recently published whole-rock analyses define a positive Se–Te correlation (Fig. 10). However, the experiments of Helmy et al. (2010, 2013) and observations in nature (Patten et al. 2013) demonstrate a much greater preference of Te for metal-rich BMS melts 14

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(metal/sulfur atomic ratio > 1) compared to covalent monosulfides. Earlier proton probe (PIXE) analyses were insufficiently sensitive to provide information on the partitioning behavior of Te; Guo et al. (1999) reported unrealistically high Te contents (16–226 µg.g−1) whereas all BMS analysed by Hattori et al. (2002) show Te content below the detection limit values of 20 µg.g−1. Using LAM-ICPMS in-situ analyses Lorand and Alard (2010) provided evidence that major BMS do not balance the whole-rock Te concentrations of fertile mantle peridotites. The missing Te fraction of the whole-rock budget (30–90 %) was attributed to randomly distributed μmsized discrete Pt–Pd tellurides. That point will be addressed further in the next section.

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High-D elements: Significant effort has recently been made to define PGE partition coefficients between sulfide melts and silicate melt through both experimental and empirical determination (Brenan et al. 2016, this volume). In parallel, progress has been made for evaluating the composition of the fertile upper mantle (DMM) and their extrapolation to the hypothetical PUM reservoir (Meisel et al. 2001; Becker et al. 2006; Fischer-Gödde et al. 2011; Day et al. 2016, this volume). Orogenic lherzolites, along with the few basalt-hosted xenoliths and oceanic peridotites that plot within the same S vs. Se, Cu vs. S and S vs. Al2O3 arrays (Figs 6, 10, 17, 18) are the cornerstone for PUM estimates. Platinum group elements occur at very constant ng.g−1 to sub ng.g−1 concentration levels in these rocks (Lorand et al. 2013 and references therein; Becker and Dale 2016, this volume; Luguet and Reisberg 2016, this volume). Normalized to CI-chondrite concentration, the estimated PUM 0.1 composition is characterized by broadly chondritic relative abundances with slight positive anomalies of light (Ru, Rh, Pd) vs. heavy (Os, Ir, Pt) PGE (Fig. 11). This 0.01 composition is currently believed to be a planetary scale signature inherited from post-cratonic harzburgites late-accreting materials postdating the 0.001 core-mantle separation event (“The Late Veneer Theory” e.g., Chou 1978; Morgan Primitive Upper Mantle 1986; Becker et al. 2006; Lorand et al. (Becker et al., 2006) 0.1 2008a; Day et al. 2016, this volume). 0.01

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Binary plots of Os vs. Ir, Ru vs. Ir and Ir vs. Pt produce reasonably good positive correlations, suggesting that all PGE are hosted in a single mineral phase (Fig. 12). The strongly chalcophile behavior of high-D elements in the terrestrial mantle is supported by the concentrations of Ir, Ru, Rh, Pt and Pd measured in separated mantle BMS (Morgan and Baedeker 1983). Pattou et al. (1996) analyzed those five PGE in a Lherz fertile lherzolite and reported total PGE content of 11.2 µg.g−1, i.e., a three order-of-magnitude higher concentration compared to the wholerock analysis. Moreoever, if normalized to the composition of CI-chondrites, the patterns of the separated sulfide fraction and the whole-rock display parallel shapes (Fig. 13). Our knowledge of the location of

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PGE mineral carriers has been greatly improved by in-situ analyses using either PIXE or LAM-ICPMS, and the amount of in-situ analyses has increased considerably over the past ten years. However, the PGE budget in PUM-type mantle lherzolites (what is really present inside BMS) was determined only in a limited number of suitable occurrences. Data are available for one ophiolitic lherzolite (Luguet et al. 2004), one orogenic sample (FON B 93; a fertile Pyrenean lherzolite studied by a wide range of analytical techniques; Lorand et al. 2008b), a Lherz fertile lherzolite (84−1, Alard et al. 2000; Lorand et al. 1999) and two abyssal peridotites from the Kane Fracture zone (Mid-Atlantic ridge; Luguet et al. 2001). All these mass balance calculations were based on in-situ LAM-ICPMS analyses of individual “primary” Type-i BMS (pentlandite, chalcopyrite). The good agreement between measured and calculated concentrations provided further evidence that BMS are the major repository for Os, Ir, Ru, Rh and Pd (Fig. 13). This was not the case for Pt for which the calculated whole-rocks always show a large (50–90%) deficit for which it is difficult to account (Fig. 13). Unlike bulk analyses of separated minerals, in-situ analyses can distinguish between the mixed signals of PGE that reside in solid solution in the BMS and discrete platinum-group minerals (PGM) hosted in the sulfide minerals. Time-resolved LAM-ICPMS spectra collected during analyses produce huge concentration spikes for these minerals, in a similar manner to Au in Fig. 9 (Luguet et al. 2001; Lorand et al. 2010). Platinum alloys were reported by early studies using optical microscopy (Garuti et al. 1984; Lorand et al. 1999). In situ analyses determined positive correlations between Pt and Te, indicating the presence of Pt tellurides (Luguet et al. 2004; Fig. 14). For example, X-ray synchrotron analyses detected Pt–Ir–Os alloys in Horoman orogenic lherzolites (Kogiso et al. 2008). Detailed SEM studies performed on Pyrenean lherzolites identified a far more diverse assemblage of Pt-rich microminerals (alloys, Pt–Ir–Os

FON B 93 measured bulk sulfide composition

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Figure 13. Examples of mass balance estimates of the contribution of BMS to the PGE budget of orogenic lherzolites (FON B−93 Fontête Rouge, Lorand et al. 2008b; 84−1, Lherz, Pyrénées, Alard et al. 2000) and abyssal lherzolite (KN 3−4; Kane Fracture Zone, 23°N, North Atlantic ridge; Luguet et al. 2001). The hypothetical bulk-composition of mantle-derived base-metal sulfide (BMS) was calculated assuming that the whole-rock PGE budget is entirely accounted for by a BMS phase with 34 wt% S corresponding to 90% pentlandite and 10% chalcopyrite (see text). Error bars on the calculated compositions (bulk sulfide for FON B 93 and 84−1; bulk rock for KN 3−4) were estimated as 1 sigma standard deviation on the average composition. The measured Os, Ir, Ru, Rh and Pd Fig. 13 are within the range of calculated values; values this suggests that BMS are the major repository for those five elements. Platinum (as well as Au, not shown) is systematically outside the range for all three samples. Both elements are controlled by other, submicrometric discrete minerals (see text). Note that the bulk analysis of a separated sulfide fraction (Pattou et al. 1996) produced a smoother CI-chondrite normalized PGE patterns devoid of a strong negative Pt anomaly compared to the Laser Ablation-ICPMS analyses of Alard et al. (2000). This suggests that most of the Pt budget is intimately associated with BMS phases, although not residing inside their crystalline lattice.

Mantle Rocks: Trace Elements Controlled By Trace Minerals 60

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Figure 14. Binary plot of Pt vs Te concentrations (in-situ LAM-ICPMS analyses of pentlandite-chalcopyrite) for Pyrenean lherzolites (A, Lorand et al. 2008b, 2010) and Liguride ophiolitic lherzolites (B, Luguet et al. 2004). Both diagrams reveal strong enrichments in Pt and Te corresponding to Pt–Te-rich microphases (Pt/Te ratio of moncheite, PtTe2) or Pt-rich alloys (see text).

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and Pt–Fe alloys, sulfides, tellurides, arsenides/ sulfarsenides, intermetallic compounds; Lorand et al. 2008b, 2010; e.g., Fig. 4d). Such microphases generate a well-known nugget effect resulting in poorly reproducible wholerock Pt concentrations. This nugget effect is the rule rather than an exception, as shown by the high relative standard deviations (RSD up to 20%) reported for Pt analyses compared to the other PGE, at least in all BMS-rich mantle lherzolites, whether orogenic (Becker et al. 2006; Lorand et al. 2008b; Fischer-Gödde et al. 2011) ophiolitic (e.g., Batanova et al. 2008) or abyssal peridotites (Liu et al. 2009). Apart from refractory Pt–Ir–Os alloys, most of the Pt-rich minerals so far described in mantle peridotites formed from BMS via fractional crystallization, and/or subsolidus exsolution/oxidation/reduction (Lorand et al. 2008b). The Pt-depletion of pentlandite is a wellknown compositional feature of PGE orebodies (Fig. 15) and is related to the incompatibility of Pt for this BMS (Barnes et al. 2006; Godel and Barnes 2008). The shallow upper mantle is not expected to be saturated with respect to any PGM. As shown by Figure 15, most of the PGE concentrations so far measured in mantle-derived sulfides from fertile lherzolites (5–25 µg.g−1 for Os and Ir, 5–30 µg.g−1 for Ru and 5–50 µg.g−1 for Pd; e.g., Burton et al. 1999, Alard et al. 2000; Luguet et al. 2001, 2004;

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Ru

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Figure 15. Comparison between high-temperature solubility of platinum-group elements (PGE) in Monosulfide solid solutions (Mss) and Pyrrhotite solid solution (Poss) and LAM-ICPMS in-situ analyses of PGE in Type-i BMS (pentlandite) (Pyrenean peridotites; Lorand et al. 2008b, 2010). In-situ analyses of pentlandite and chalcopyrite from magmatic PGE ore bodies are shown for comparison (modified after Lorand et al. 2013). Note that mantle-derived pentlandite show very similar CI-chondrite normalized PGE pattern as pentlandite from PGE ore bodies, although on average poorer in Rh and Pd.

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Lorand et al. 2008b, 2010) are several orders of magnitude lower than experimentally determined solubility of PGE in the temperature range of the shallow upper mantle (hundreds to thousands of µg.g−1 of each PGE). If the Mss is stable in the uppermost convecting mantle, as suggested by experimental work (Bockrath et al. 2004a), it could accommodate wt.% levels of Os, Ir, Ru, and Rh at 900–1100 °C (Li et al. 1996; Mungall et al. 2005; Ballhaus et al. 2006). Mazlan et al. (2002) and Mackovicky and Karup Møller (2000) reported 5 wt.% Pt in Fe0.9S at 1100 °C in presence of Pt–Fe alloys. Fonseca et al. (2011) found 140–960 µg.g−1 Os and 4600–7800 µg.g−1 Ir in molten FeS at 1200 °C. Even at saturation with respect to Os–Ir, they reported solubility for Os, Ir, Pt, and Pd (55–240 µg.g−1 Os, 370−3200 µg.g−1 Ir, 8,100–15,700 µg.g−1 Ru, 6,100–62, 000 µg.g−1 Pt, depending how high the ƒS2 is) which are still large compared to those measured in mantle-derived BMS from fertile mantle samples. Experimentally determined solubilities in subsolidus BMS (pentlandite, pyrrhotite) increase with ƒS2 from 1.8 (FeS) to 11.8 wt.% (Fe7S8) for Pd, 0.05–2.2 wt.% for Pt, 0.1–3.3 wt.% for Ru and 6.5–44 wt.% for Rh (Mackovicky et al. 1986; Ballhaus and Ulmer 1995; Fig. 15). It can reasonably be assumed that this huge solubility of PGE into the BMS supresses any significant incorporation of high-D elements into silicates and Al–Cr-spinel. Each attempt at very precise analyses of separates of these minerals led to the conclusion that the major silicates are negligible (≪ 5%) contributors to the whole-rock PGE budget. This is also true for Al-spinel, especially studied for Os (0.036 ng.g−1; Hart Ravizza 1996; < 0.174 ng.g−1, Pearson et al. 2004; Lorand et al. 2008b, < 0.1 ng.g−1; Fig. 13). Moreover, even the cleanest mineral separates were recognized to be prone to contamination by PGE-rich micro-inclusion that are not seen under binocular microscope due to their very small average grain size in the sub-μm range (Handler and Bennett 1999). Evidence of PGE compatibility in silicates either comes from mantle rocks that have been modified by petrogenetic processes that removed BMS, e.g., partial melting residues (see next section), from liquidus phases of mantle-derived melts such as OIB, komatiites and related rocks (Burton et al. 2002; Day 2013) or from experimental works (Brenan et al. 2016, this volume).

1600°C s

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Figure 16. P–T diagram showing the liquidus and solidus curve of Mss (Bockrath et al. 2004a) compared 16 to various melting curve forFig.troilite. Note that the Bockrath et al. curves suggest that the two-phase volume Mss+ Cu-Ni-rich sulfide melt expands at increasing depth. Cu–Fe–Ni–S system at 1000 °C after Craig and Kullerud (1969) and Rahgavan (2004). Lherzolite anhydrous solidus from Pearson et al. (2003).

Partial melting of the mantle: a BMS-removing and PGM producing petrogenetic process.

Base metal sulfides have low melting temperatures compared to other upper mantle minerals. Experimentally determined liquidus curves show a minimum positive dependence on pressure, although some disagreement remains between the different authors about the exact position of the curves (Fig. 16). Figure 16 suggests a partially molten to homogeneous sulfide melt (“matte”) to occur in the uppermost mantle. Consistent with the low melting temperature of BMS, S has long been demonstrated to be a moderately incompatible element in mantle processes, including partial melting. It is enriched by a factor of 5−10 in mantle-derived primary melt (e.g., MORB) compared to the preferred mantle values. Partial melting is thus sulfur-consuming and the upper mantle is usually assumed to be completely

Mantle Rocks: Trace Elements Controlled By Trace Minerals Cratonic xenoliths

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stripped of BMS once the amount of S extracted is no longer sufficient to match the sulfur capacity at sulfur saturation (SCSS) of the partial melt. Mavrogenes and O’Neill (1999) studied SCSS in basaltic melts containing 6–14 wt.% FeO at pressures ranging over 5–90 kbar and temperatures of 1400–1800 °C. Their data indicate that SCSS decreases exponentially from a few thousands of µg.g−1 to a few hundred µg.g−1 with increasing pressure, and increases only slightly with temperature. Fig. 17

S (µg.g-1)

There are now more than 1,000 whole-rock S analyses available in the literature (taking into account only 300 high-precision analyses). The most250 refractory, Al-poor rocks are devoid PUM of S, in agreement with the complete 200 extraction of BMS at the highest degrees of partial melting (Fig. 17). In practice, 150 the mass fraction of BMS left after extraction of a given melt increment 100 S/BMS is usually calculated assuming a removal fertile mantle source containing 50 250 ± 50 µg.g−1 S, an initial BMS phase with 35 wt.% S in 100 wt.% sulfides, a 0 0 1 2 3 4 sulfur solubility of 1,000 µg.g−1 in the Al2O3 (wt.%) partial melt and a batch/equilibrium Figure 17. Plots of S abundances vs bulk rock Al2O3 melting process controlling dissolution contents taken as fertility index). See Figure 7 for the of the BMS into the partial melt. Under hydrothermal trend defined by cratonic and abyssal those “standard” conditions, BMS peridotites. This trend overlaps another S enrichment disappear at F = 25% (e.g., Lorand pattern at constant Al2O3 which was ascribed to late magmatic BMS in abyssal and ophiolitic peridotites 1989a; Keays 1995; Fischer-Gödde et (Luguet et al. 2003; Marchesi et al. 2013; Lorand et al. al. 2011). Only orogenic lherzolites 2009). By contrast, most of post-cratonic basalt-hosted and related S-rich basalt-hosted mantle spinel peridotite xenoliths have lost S, irrespective of xenoliths, in addition to some ophiolitic their fertility index. The curves indicates the amount of rocks and abyssal peridotites plot on S left after extraction of a given melt increment. It was calculated assuming a sulphur solubility of 1000 µg.g−1 a robust trend with respect to meltin the partial melt in the case of spinel peridotites and depletion indicators (Fig. 17). This 700 µg.g−1 for garnet peridotites (Mavrogenes and trend is widely interpreted as a partial O’Neill 1999). Primitive Upper Mantle (PUM) values melting trend, indicating similar bulk after Lorand (1990), McDonough and Sun (1995) and solid–liquid partition coefficients for S Palme and O’Neill (2003). See Lorand et al. (2013) and references therein for further detail on models. and Al2O3 (0.2−0.3; e.g., Lorand 1991). The S concentrations lying above and below this trend reflect some inhomogeneity in the BMS distribution at a hand sample scale, coupled with the strong pressure effect on SCSS that modifies the point of complete BMS exhaustion (see below). Cutting at a high angle to the melting trend are cratonic peridotites that are strongly S-enriched by hydrothermal BMS, ancient metasomatism, and interaction with host kimberlites (Griffin et al. 2004; Pearson et al. 2004; Lorand and Grégoire 2006; Aulbach et al. 2016, this volume). By contrast, the majority of basalt-hosted xenoliths are

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strongly depleted. Their S concentrations scatter parallel to the x-axis and do not correlate with fertility indices. Such a distribution points to a secondary, post-melting mobility for S. As far as “PUM-like” peridotites are concerned, Cu abundances mirror that of S in defining a positive correlation with melt depletion indicators (Fig. 18). The Al2O3-poor refractory samples that escaped hydrothermal alteration or metasomatism are almost Cu-free (< 3 µg.g−1), in agreement with total removal of BMS by partial melting. Most basalt-borne xenoliths are Cu-depleted compared to the Cu vs. S trend of Fig. 18, irrespective of their degree of fertility. The PGE have long been recognized to fractionate from each other with the extent correlating with increasing melting point of the pure metal (e.g., Barnes et al. 1988). The observed compatibility of the PGE in mantle-rocks, in layered intrusions and the relative enrichments of these elements in basalts and komatiites has led to an overall agreement that bulk peridotite-melt partition coefficients follow the sequence: Pd = Pt ≤ Rh = Ir ≤ Ru ≤ Os (Barnes et al. 1988; Lorand et al. 2008a; 2013; Walker 2009; Day 2013 and references therein;

50

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Al2O3 (wt.%) Figure 18. Plots of Cu abundances vs fertility index (whole-rock data). Hydrothermal sulfide enrichment trend as in Figures 7 and 17. The various Cu enrichment trends documented in ophiolitic (Schulte et al. 2008, Aldanmaz et al. 2012; Uyssal et al. 2012) and abyssal peridotites (Marchesi et al. 2013) correspond to late magmatic Cu-sulfide Fig. 18 precipitation controlled by clinopyroxene + spinel + BMS precipitation (see Fig. 28). Note that orogenic peridotites follow a similar Cu enrichment trend compared with PUM estimates (Fellow and Canil 2012; Liu et al. 2014). By contrast, most basalt-borne xenoliths are Cu-depleted irrespective of their fertility. Two curves were calculated for Cu; one (short dashed lines) is based on DCu (sulfide melt⁄silicate melt of 103 (Handler et al. 1999; Mungall and Brenan 2014). The other cuve (long dashed line) assumes that Cu is controlled by monosulfide solid solution (Mss)-sulfide liquid (sl) partitioning during melt extraction (Fischer-Gödde et al. 2011). Experimentally determined Mss-liquid sulfide partition coefficients (DCuMss/sulfide melt of 0.1) were chosen from the literature (Mungall et al. 2005; Ballhaus et al. 2006) and were scaled according to the mass fraction of sulfides. The model should produce concentration plateaus as long as the Cu-Ni-rich sulfide melt is trapped in the solid residue. Physical removing of the Cu–Ni-rich sulfide melt is expected to lower the concentrations of Cu to near 0 over a narrow range of melt extraction percentage. Labels on long-dashed line indicate melting percentages.

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see also Fig. 11). Only Pd shows covariation trends with S, yet those trends are considerably scattered by the post-melting alteration of S abundances and late-magmatic/metasomatic disturbances of Pd abundances in oceanic and continental peridotites (Fig. 19). Iridium and Pt concentrations are insensitive to melt depletion indicators (Fig. 20). Geochemical modelling of PGE abundances in peridotitic residues assumes that the partitioning behavior of PGE is controlled by sulfide melt–silicate melt partition coefficients and congruent melting of BMS (e.g., Morgan 1986; Lorand et al. 1999; Luguet et al. 2003; Handler and Bennett 1999; Pearson et al. 2004). Most recent D sulfide melt/ silicate melt estimates range from 106 for Pd to 3 x 106 for Pt and 4 x 105 for Ru (Mungall and Brenan 2014); hence, they are unable to account for the decoupling between Pd and the other PGE at increasing melting degree. These simple models predict that PGE should behave compatibly as long as BMS survive in solid residues of melting. These elements should start behaving incompatibly only once BMS are totally eliminated (e.g., Keays 1995 and reference therein).

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Figure 19. Plots of Pd vs S (bulk-rock data). PUM estimates after Becker et al. (2006) and Fischer-Gödde et al. (2011). Palladium abundances were modelled with mass-balance equations for fractional melting of the sulfide phase, assuming all of the Pd residing in the sulfides, with DPd (sulfide melt⁄silicate melt) of 10,000 (long-dashed curve) and 1,000 (full line) and a solubility of S in the partial melt of 900 and 1000 µg.g−1 (short-dashed lines). For details on modelling, see Lorand et al. (2009, 2013 and reference therein). Note that many ophiolitic and abyssal peridotites (Rehkämper et al. 1999; Luguet et al. 2003; Lorand et al. 2009; Marchesi et al. 2013) do not plot on the trend defined by the continental mantle peridotites that are currently believed to be melting residues from a PUM-like mantle source. The Pd vs. S trends of oceanic peridotites correspond to mixing trend between a strongly Pd-depleted harzburgitic mantle and Pd-rich late-magmatic/metasomatic BMS. More extreme Pd enrichment at very low S content is seen in continental peridotites (Ackerman et al. 2009) which may track discrete Pd-rich platinum-group minerals (PGM) coupled with the overall S depletion of these rocks (see text).

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A different school of thought assumes that mantle BMS melts incongruently, Off-craton xenoliths which is based on the high temperature, 10 Orogenic lherzolites high-pressure phase diagram of the Fe– PUM Ni–Cu–S system (Fig. 16). A two-phase 8 volume surrounds the stability field of Mss and expands at increasing pressure 6 (Bockrath et al. 2004a). Thus, BMS should melt incongruently, liberating first a 4 Cu–Ni-rich sulfide melt at low temperature that accommodates Pt and Pd in square2 planar sites while leaving residual Mss that hydrothermal sulfides have octahedral sites for sequestering Re, 0 Fig. 20 Os, Ir, Ru, and Rh (Bockrath et al. 2004a; 8 Ophiolites Ballhaus et al. 2006). Incongruent-melting Abyssal peridotites 7 models of BMS are broadly successful in accounting for the behavior of Mss6 compatible elements (Os, Ir, Ru, Rh), as 5 long as BMS is residual Mss, not extracted PUM from the residual peridotites. The increasing 4 Ir contents at decreasing S contents of 3 Figure 20 is reproduced by choosing very high sulfide-melt–silicate-melt partition 2 coefficients (> 105) in line with most 1 recent experiments. The overall coherent behavior of Pd, Au, and Cu (Figs. 7, 0 0 50 100 150 200 250 300 350 19–20) is consistent with the evolution line S (µg.g-1) for residual Mss after gradual extraction of the Cu–Ni sulfide melt from solid Figure 20. Plot of Pt vs S and Ir vs S (bulk-rock data). Solid lines delineate the field for orogenic melting residues (e.g., Ballhaus et al. 2006; peridotites (excluding a few outliers). Note the Fisher-Gödde et al. (2011). Residual Mss overall Pt depletion of cratonic peridotites compared inclusions (now decomposed by subsolidus to post-cratonic peridotites and the Ir-depletion of re-equilibration into a Ni-rich and a Ni-poor basalt-borne lherzolites. Mss ± pentlandite) have been documented in many off-craton lherzolite xenoliths (e.g., Lorand and Conquéré 1983; Dromgoole and Pasteris 1987; Szabó and Bodnar 1995; Guo et al. 1999; Alard et al. 2000; Lorand and Alard 2001; Wang et al. 2009 and references therein). One objection against incongruent melting models is the poorly determined liquidus–solidus relationships of the Cu–Fe–Ni–S system at mantle P–T conditions; in other words, whether Mss can be stable once silicate partial melting occurs is not universally accepted (Fig. 16). The liquidus curve of Mss certainly precludes its survival as a solid for hypersolidus temperatures reached in extreme mantle melting degrees (yet there is disagreement between the different experimental data sets; Fig. 16). Enclosed Mss blebs in mantle peridotite xenoliths show a rounded geometry typical of sulfide liquid with a minimum surface area and non-wetting capacity (Hattori et al; 2002; Ballhaus and Ellis 2006). Molten Mss were physically excluded from the peridotite-melt system and, thus, preserved at some transient melting degrees by dynamic or static recrystallization of olivine. Cratonic xenoliths

Ir (ng.g-1)

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Another major objection against incongruent melting model of mantle BMS is the inability of this model to explain the more severe Pd depletion relative to broadly constant Pt concentrations that is observed for instance between off-cratonic harzburgites and Archean cratonic harzburgites (Figs. 11, 20). In a natural peridotite, at low to moderate degrees of partial melt extraction (F = 10%), Pt is preferentially retained in the residue relative to Pd. However, all available

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laboratory partitioning experiments reported almost identical or very similar D Mss/sulfide melt (0.1−0.2) for Pt and Pd, indicating similar incompatible behavior for both elements in the Mss structure (Li et al. 1996; Barnes et al. 2001; Mungall et al. 2005; Ballhaus et al. 2006). The conundrum of this Pt–Pd decoupling was solved when Pt–Ir–Os alloys were detected by X-ray synchrotron analyses in refractory peridotites (Horoman harzburgite, Japan; Kogiso et al. 2008) and by detailed SEM studies (Luguet et al. 2007; Lorand et al. 2010). These alloys are interpreted as desulfidation products from Mss (Fonseca et al. 2012, and reference therein). Peregoedova et al. (2004) conducted desulfidation experiments on Mss which produced metal-rich Mss, Pt–Ir alloys and a Cu–Pd-rich sulfide melt at 1000°C for sulfur fugacity conditions corresponding to the Fe–FeS buffer and 10 µg.g−1 bulk Pt content in the BMS assemblage. Partition coefficients for Pt decrease with the S content of the Mss by a factor of 20 (D Mss/sulfide melt = 0.2−0.01) from S-oversaturated (Po-saturation) to S-undersaturated (Fe metal-saturation; Mungall et al. 2005; Balhaus et al. 2006). Thus, any process reducing the ƒS2 will drive PGE alloy nucleation. Moreover, due to the extreme D values for PGE (Mungall and Brenan 2014), residual BMS will accumulate very high PGE tenors for increasing melting percentages, which make it a near certainty that they will be saturated with respect to Pt–Ir alloys for quite low melting degrees (10%).

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1000/T (°K) Fig. 22 the sulfidation curves of Figure 22. Log ƒS2–T diagram locating Ru, Pt, Ir, and Os at 1.5 GPa and the domain (hatched area) in which RuS2 is stable along with Pt-rich alloys while Mss is unstable. The olivine–opx–Mss curve is from Eggler and Lorand (1993) (assuming log ƒO2 = FMQ − 1.5 log unit for the DMM). Redrawn after Melcher (2000) and Luguet et al. (2007); the sulfidation curves for Os and Ir are from Barin (1995) extrapolated to 1.5 GPa using molar volume data of Robie et al. (1995). Approximate stability field for the phase CuPt2S4. The sulfur condensation curve was omitted for clarity.

Both experiments and calculated stabilities of PGM in the log ƒS2–T diagram of Figure 22 indicate that Pt-rich alloys start to form at higher sulfur fugacity than do Ru-rich alloys/sulfides. Thus, Pt–Ir–Os alloys can coexist with Mss in as much as the solubility of Pt in residual Mss is low compared to the solubility of Ru (Fonseca et al. 2012). Laurite and Mss can coexist at higher degree of melting, over a very narrow window in the log ƒS2–T space, before the Mss is totally substituted by laurite, and then by Ru–Os–Ir alloys with a further decrease in the fugacity of sulfur. Detailed integrated studies of orogenic peridotites have identified laurite + Pt–Ir–Os alloys in BMS-free residual harzburgites (Luguet et al. 2007). This assemblage, located in the intergranular pores, accounts for more than 90% of the wholerock PGE budget. It is likely more abundant than indicated by SEM data, as suggested by the poor reproducibility of whole-rock Pt concentrations reported for refractory peridotites (RSD up to 25%; Lorand et al. 1999, 2004; Luguet et al. 2007; Kogiso et al. 2008; Liu et al. 2009).

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Figure 21. In-situ analyses of BMS assemblages from peridotite xenoliths compared with solubility of platinum-group elements (PGE) in Monosulfide solid solutions (Mss) and Pyrrhotite solid solution (Poss)(see Lorand et al. 2013 and references therein). Osmium–Ir– Ru-rich (> × 10 CI-chondrites) compositions correspond to Mss-dominated Type-e BMS inclusions; note the high Os and Ir contents in cratonic Mss, which agrees with the strongly chalcophile behavior of this PGE population; opposite “basalt-like” patterns (Pd/Ir ratios > 10) correspond to Cu-rich Type-i assemblages (isocubanite + chalcopyrite). Intermediate patterns characterized by nearly flat segment between Os and Rh and a strong negative Ptanomaly are Type-i pentlandite similar to orogenic lherzolites patterns. CI-chondrite values after Fischer-Gödde et al. (2011).

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Some speculation can be made about which minerals (refractory alloys or sulfides) control the budget of Os, Ir, and Fig. Ru 21 in BMS-free refractory peridotites at different depths in the mantle, based on the pressure dependence of the solubility of S in basaltic melts. Alloys of Os–Ir–Ru are expected to be stable in the low-pressure environment characterizing the melting conditions of MORB because BMS are eliminated at only 12% melting. Luguet et al. (2003) were the first to identify Ru–Os–Ir alloys in abyssal harzburgites. Platinum-depleted Os–Ir–Ru alloys were also reported from ophiolitic chromitites that are low-pressure crystallization products from S-undersaturated melts (O’Driscoll and Gonzáles-Jiménez 2016, this volume). Osmium– Ir–Ru alloys can also be direct precipitates from boninitic melts, i.e., S-undersaturated melts originating at very shallow levels in the supra-subduction zone mantle (Peck et al. 1992). Laurite–erlichmanite + Pt–Ir–Os alloys are common in continental harzburgites and moderately depleted ophiolitic harzburgites (Lorand et al. 2009 2010). Cratonic peridotites have not been studied in detail for PGM. However, Mss is expected to be stable over a larger range of partial melting degrees because the solubility of S in partial melts is expected to decrease significantly (down to 600 µg.g−1) at pressure of > 3 GPa (Mavrogenes and O’Neill 1999). Type-e Mss containing Os + Ir + Ru + Rh concentrations > 1,000 µg.g−1 have been reported by Alard et al. (2000) and Griffin et al. (2002, 2004) in cratonic peridotites (Fig. 21). Of course, the inferences about the PGE mineralogy as a function of pressure is likely complicated due to insulation of BMS and/or PGM inside silicates and chromite. However, they are supported by the Se–Te whole-rock systematics of refractory peridotites that were recently retrieved by high-precision Te analyses (König et al. 2012, 2014; see Fig. 10). The BMS-poor refractory samples display suprachondritic Se/Te (> 20, up to 30), more than double the average ratio of lherzolites (cf. König et al. 2012, 2014). These ratios contradict the common belief that Te is more compatible than Se during partial melting (Hattori et al. 2002; Patten et al. 2013; Wang et al. 2013). Restitic Mss is likely to cause this Se/Te increase because it concentrates Se anions more readily than Te (Lorand and Alard 2001; Guo et al. 1999; Helmy et al. 2010). Selenium can also replace S in the anionic sublattice of pyrite-type disulfides like laurite which may contain up to 200 µg.g−1 Se (e.g., Hattori et al. 2004). Of course, the story should be different if Ru–Os–Ir alloys were stable. However, no refractory peridotites containing such alloys have been analyzed for Se and Te.

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Luguet et al. (2007) detected up to 15 ng.g−1 total PGE contents in Cr-spinel from BMS-free Pyrenean harzburgites. However, it should be stressed that Cr-spinel occurs in too low a modal abundance ( 2 wt.%) are plotted (redrawn after Lorand et al. 2013).

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al. (1999), and Lorand et al. (2003a,b) summarized the diagnostic features of S-related supergene weathering as (i) mantle-derived BMS replaced by low-temperature Fe-oxyhydroxide phases (Fig. 2D), (ii) sub-chondritic S/Se ratios indicating selective S loss as mobile sulfates with respect to Se, a much less mobile element (Fig. 6) (iii) Cu/S increasing dramatically above the mean value defined by sulfide modal compositions (0.05−0.15) due to the more limited mobility of Cu relative to S in ground water (Fig. 7) (iv) lack of a covariation trend between measured bulkrock S abundances and BMS modal abundances. Xenolith suites that fit these four criteria were documented in Eastern Australia (Handler et al. 1999), Sidamo, Ethiopia (Lorand et al. 2003b), French Massif Central (Lorand 1990; Lorand et al. 2003a) including a BMS-rich population (Montferrier, Alard et al. 2011) and Subei bassin xenoliths (China) (Reisberg et al. 2005). However, there is growing evidence to suggest that weathering was likely not the only process that removed S because some xenolith suites lacking evidence for secondary oxidation products of BMS (Fe oxyhydroxides, hematite) are strongly depleted in inert chalcophiles like Se and Pt (e.g., Ionov et al. 1992; Ackerman et al. 2009; Liu et al. 2010). A consensus has recently emerged to ascribe the overall chalcophile/siderophile element depletion to regional silicate melt percolation events that predate erupted lavas in rift-related subcontinental lithospheric mantle (e.g., Lee 2002; Lorand et al. 2003a,b; Ackerman et al. 2009; Liu et al. 2010). The distinctive PGE patterns of chalcophile–siderophile element-depleted peridotite xenoliths is ascribed to mechanical removal/dissolution of BMS in percolating melts. Thin sections of these xenoliths (which generally show granular textures), typically contain one or two tiny BMS grains, and most have none. This was noted for Massif Central xenoliths (Lorand and Conquéré 1983), Nògràd-Gömör, Hungary-Slovakia xenoliths (Szabó and Bodnar 1995), anhydrous and hydrous Eifel lherzolite xenoliths, Germany (Shaw 1997), Northern Bohemian Kozakov xenoliths (Ackerman et al. 2009), xenoliths from Taiwan (Wang et al. 2009) and Yangyuan xenoliths in China (Liu et al. 2010). In addition to this widespread BMS depletion, the observed BMS are mostly Type-e rounded Mss or multiphase assemblages of Mss + minor amount of pentlandite and/or chalcopyrite. Type-e Mss grains are typically metal-deficient ((Fe+Ni+Cu)/S ratios < 1, and as low as 0.87) two-phase (Ni-rich and Fe-rich) Mss, although less Ni-rich than Type-i pentlandite-dominated BMS assemblages (Figs. 2a, 5). The arch-shaped bulk-rock patterns of chalcophile–siderophile element depleted off-craton xenoliths closely resemble patterns of residual Mss (Fig. 27, cf. Fig. 21). LAM-ICPMS analyses of Type-e Mss from postcratonic peridotite xenoliths were reported by Alard et al. (2000), Lorand and Alard (2001), Powell and O’Reilly (2007) and Wang et al. (2009) (Fig. 21). Their strongly Pt- and Pd-depleted CI-normalized PGE patterns relative to compatible PGE (Os, Ir, Ru) were considered to be a fingerprint of Mss having survived partial melting. Lorand and Alard (2001) evaluated the bulk-rock PGE budget of four off-craton spinel lherzolite xenoliths and found that it could be accounted for by Mss for Ir, Rh, Pt, and Pd, but a systematic (20−30%) deficit for Ru was attributed to unrepresentative BMS sampling. Elsewhere, Os was found to be hosted in Mss (e.g., Harvey et al. 2010, 2011). All these observations are strong evidence for Mss-dominated BMS assemblages in chalcophile–siderophile element depleted mantle xenoliths. Basic parameters for the effect of melt percolation on chalcophile–siderophile elements were presented for strongly depleted dunitic wall-rock to pyroxenite in ophiolites (Büchl et al. 2002) and hornblendites in OIB-hosted xenoliths (Lorand et al. 2004), for dunitic channels in abyssal peridotites (Marchesi et al. 2013) and cratonic samples (Tanzania, Chesley et al. 1999; Greenland peridotites; Wittig et al. 2010). In principle, melt flowing through grain boundaries of a peridotitic matrix can remove pre-existing BMS at low pressure (1 GPa) where SCSS values in percolating silicate melt are at the highest levels. Moreover, regarding dunites, olivine-producing reactions are effectively open system melting reactions, i.e., melt-producing and sulfur-diluting (e.g., Büchl et al. 2002 and reference therein). For continental peridotites, the chalcophile–siderophile element depletion is independent of matrix mineralogy. It is the

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increasing SCSS in the percolating melt at decreasing pressure that is the driving factor of BMS dissolution. If a fluid evolves at low pressure, S should become incorporated into that fluid, thus reducing further the stability of corresponding solid BMS. Two other critical parameters for the stability of BMS in melt percolation process are temperature and oxygen fugacity. Type-e Mss are all metal-deficient Mss (Fe + Ni + Cu)/S < 1). Eggler and Lorand (1993) calibrated the sulfidation reaction of olivine + S2 = orthopyroxene + Mss + O2 as a sulfide oxybarometer. This latter indicates ƒO2 values > FMQ for several chalcophile–siderophile element-depleted xenolith suites, i.e., Massif Central samples (Eggler and Lorand 1993), Nògràd-Gömör Volcanic field (Szabó and Bodnar 1995) and Yangyuan peridotites (Liu et al. 2010). Moreover, in all these occurrences, the distribution coefficient for the Ni–Fe exchange reaction between olivine and Mss is also low (3 < Kd < 23), especially when compared to experimentally determined values (cf. > 40; Barnes et al. 2013). Possibly these low Kd might be indicative of oxidizing conditions for type-e Mss, unrepresentative of the mantle as a whole. It is assumed that pyrrhotite solid–solution incorporates Ni as NiS, i.e., the higher the oxygen content, the less ideal the incorporation of Ni into the pyrrhotite structures becomes, as a result of the increasing number of vacancies in the pyrrhotite lattice from parallel incorporation of Fe as Fe3+ (Kawakami et al. 2006 and references therein). Moreover, for ƒO2 > FMQ + 2 log units, sulfates become important S species. The solubility of SO4 in silicate melts is higher by a factor of 10 compared to that of reduced sulfides (Jugo et al. 2005), so that oxidized magmas are expected to more readily dissolve BMS compared to reduced magmas. The solubility of PGE in silicate melts also increases by a factor 10 for a modest increase of ƒO2 from WM (wüstite–magnetite reference buffer) to FMQ (Borissov and Palme 2000; Palme 2008; Brenan et al. 2016, this volume). So, hot, oxidized (ƒO2 > FMQ) basaltic melts may not be saturated with respect to PGE in their mantle source regions (especially Os which easily forms the oxide OsO4) and thus would more readily dissolve refractory PGM. The incompatible behavior of Re, as well as low Re/Ir in chalcophile–siderophile element depleted xenoliths, is also consistent with an overall oxidizing re-equilibration (Ackerman et al. 2009; Liu et al. 2010). Several authors have also pointed out that the most depleted samples are those equilibrated at the highest temperatures (> 1100 °C) and collected from the deepest part of the subcontinental mantle lithosphere (Lorand et al. 2003a,b; Ackerman et al. 2009). These temperature-dependent depletion trends may reflect the positive influence of temperature on the solubility of PGE in silicate melts (Borissov and Palme 2000 and references therein). On average, basalt-hosted peridotite xenoliths that are extremely depleted in PGE show CI-chondrite normalized PGE patterns characterized by positive Ru anomalies (Fig. 27). Such patterns suggest that in low-S environments, Ru shows the greatest affinity for Type-e Mss that survived melt percolation, in agreement with its strongly chalcophile behavior compared to the other PGE (Fig. 22). An alternative way by which molten BMS at mantle temperatures (at least in the shallowest SCLM; Fig. 16) may be removed from peridotites is by migrating through intergranular pore spaces (e.g., Lorand et al. 2003a,b). However, both experimental evidence and observations in nature have raised significant objections to large-scale migrations of BMS through the existing porosity of the upper mantle. All experimental studies suggest that sulfide melt will be non-wetting against olivine-rich lithologies (Ballhaus and Ellis 1996; Gaetani and Grove 1999; Rose and Brenan 2006). Hattori et al. (2002) and Liu et al. (2010) pointed out that liquid Fe–Ni monosulfides that were trapped as Mss in olivine bear evidence for these non-wetting properties with respect to silicate phases. Experiments on sulfide surface tension (Mungall and Su 2005) demonstrated that sulfide liquid does not form interconnected thin films on the surfaces of silicate minerals at least until their abundance increases above about 5−10%. Ackerman et al. (2009) and Liu et al. (2010) suggested that the bulk-rock chalcophile/ siderophile element budget of a melt-impregnated mantle column is more likely controlled

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by dissolution of any intergranular BMS component by percolating silicate melts. The low wetting properties of Fe-Ni monosulfide liquid indirectly account for the overabundance of Type-e Mss in basalt-hosted peridotite xenoliths, in addition to preferential recrystallization of olivine in the presence of melt films (Lorand 1987a). Griffin et al. (2002) described several generations of Type-e Mss in Udachnaya kimberlitic olivine xenocrysts (characterized by different chalcophile–siderophile trace element patterns) trapped by necking down processes of initially fine-grained olivine during recrystallization. The high PGE contents of Type-e metal-deficient Mss is also linked to oxidizing conditions. Li et al. (1996) and Barnes et al. (2001) experimentally demonstrated that NiAs pyrrhotite structure accumulates defects under increasing ƒO2 which distort octahedral sites in the metal sub-lattice such that even Pt and Pd, the most incompatible PGE, can be accommodated in metal-deficient Mss. Most bulk-rock analyses of PGE-poor xenoliths show low, sub-chondritic Os/Ir ratios which have tentatively been linked to the oxidizing conditions developed during porous flow melt percolation in the subcontinental lithosphere (Fig. 27). Handler et al. (1999) reported a positive correlation between Ir/Os and Cu/S and Lorand et al. (2013), a positive correlation between Os/Ir and bulk-rock S contents. These two correlations clearly show that Os and S losses are linked and took place only in pore spaces of the peridotites as Type-e Mss exhibits unfractionated Os/Ir (Fig. 21). Handler et al. (1999) suggested that both elements were lost during alteration of BMS during extrusion in their lavas and contact with superficial, highly oxidizing athmospheric conditions that may cause Os to be highly volatile as OsO4. However, Liu et al. (2010) argued that calculated ƒO2 for low-PGE peridotite xenoliths are still lower than those required to oxidize Os, i.e., ƒO2 = FMQ + 2 or greater, and Pd and Ru should also show some volatility which has not been detected. These authors propose that Os could be fractionated from Ir–Ru–Pt via incongruent sulfide breakdown during melt percolation. For oxidizing conditions Os in Mss may preferentially dissolve into the melt/fluid phase and would leave Ir–Ru–Pt to be taken up into Pt-alloys that exsolve from Mss through desulfidation reactions at decreasing sulfur fugacity and increasing oxygen fugacity (Peregoedova et al. 2004). In that case the most siderophile, less volatile PGE (Ir) tend to be strongly partitioned into the Pt alloys (Dalloy/sulfide melt = 135−324), while Pd is strongly rejected (D = 0.04−0.16; Fleet et al. 1991). There is no reported Pt–Ir alloy occurrence in basalt-hosted peridotite xenoliths, which supports the Mss-desulfidation origin for sub-chondritic Os/Ir. Platinum-group mineral occurrences in basalt-hosted peridotite xenoliths are exceedingly rare, or have simply been overlooked. Handler and Bennett (1999) identified an insoluble component enriched in Pt and Pd over Ru and Ir in chalcophile–siderophile element depleted xenoliths from eastern Australia. This Pt–Pd-rich component was assumed to be residing in discrete minerals which were identified as PtS and Pd–Sn–S exsolution phases inside fractured Type-e Mss (Keays et al. 1981). The lack of refractory alloys of Pt and Ir contradicts the Mss desulfidation model as being responsible for sub-chondritic Os/Ir ratios in basalt-borne xenoliths.

BMS precipitation associated with magma percolation/metasomatism Evidence from the subcontinental mantle. There is now a wealth of evidence to suggest that, following dissolution and removal of BMS phases, melt-consuming reactions at decreasing melt volume can take place in a cooling mantle which then cause precipitation of newly formed Pt-, Pd-, Au-enriched BMS once the vanishing silicate melt fractions reach S-saturation (Lorand et al. 2013 and reference therein). At the bulk-rock scale, numerous PGE-depleted non-cratonic xenoliths show positive PdN/PtN (N = CI-chondrite-normalized) ratios at odds with a residual origin (Fig. 27). All those features related to chalcophile–siderophile re-enrichment of the lithospheric mantle are usually grouped under the generic term metasomatism/late-magmatic refertilization reactions (Lorand et al. 2013; Bodinier and Godard 2014; Pearson et al. 2014 and references therein). At the thin-section scale, in addition to Type-e Mss, metasomatized samples contain Type-i metal-rich (M/S > 1) Cu–Ni BMS assemblages dominated by pentlandite,

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chalcopyrite, Ni-rich sulfides and/or metal-rich Mss (Fig. 5). Those Type-i BMS assemblages display morphologies very similar to Type-i BMS in PUM-type fertile lherzolites, i.e., large grains (up to 500 µm across) with polyhedral shapes and convex inward or convex outward grain boundaries. These angular sulfides thinly spread along silicate surfaces indicating wetting properties at silicate triple junction (Fig. 4). Some peridotite xenolith occurrences have been reported to be strongly enriched in Type-i BMS (up to 1 wt.% in Hungarian protogranulartextured lherzolites (Szabó and Bodnar 1995) and Taiwan xenoliths (Tungtchiyu suites) (Wang et al. 2009). Moreover, Type-i BMS are highly enriched in Ni (> 40 wt.% Ni). This extreme Ni-enrichment occurred under more reduced conditions than those preserved by Type-e Mss (ƒO2 = FMQ − 1 log units; Szabó and Bodnar 1995) in agreement with asthenosphere-derived metasomatic melts/fluids (cf. Woodland et al. 1992). Many SCLM peridotite xenoliths analyzed for minor and trace chalcophile–siderophile elements contain both Type-e and Type-i BMS. Compositional differences between Type-e and Type-i were first recognized by Lorand and Conquéré (1983) with regard to major elements and M/S. Burton et al. (1999, 2000), Alard et al. (2000, 2002) and Griffin et al. (2002, 2004) identified unrelated PGE and Os isotopic signatures for the contrasting BMS populations in xenoliths from Kilbourne Hole (USA), the Massif Central (France), south eastern Australia and Kaapvaal and Siberian cratonic peridotites. Type-e Mss all show features of residual sulfides, i.e., low abundances of Cu sulfides (mostly occurring as isocubanite), strong depletions of Pt, Pd and Os–Ir–Ru enrichment, coupled with unradiogenic 187Os/188Os (≪ 0.1290, the PUM value; Meisel et al. 2001) consistent with long-term Re removal by partial melting. Type-i BMS show radiogenic 187Os/188Os coupled with Os–Ir–Ru-depleted basalt-like CI-normalized PGE patterns (Fig. 21). Small-scale variations in the proportions of BMS populations readily explain the huge range of Pt/Pd ratios that characterize subcontinental mantle peridotites (Fig. 27). Melt percolation-related transfer of chalcophile–siderophile elements operates at the scale of a single melt column. A bottom to top increase of Type-i BMS modal abundances was documented in the Massif Central (Lorand 2003a) and at Sidamo, Ethiopia (Lorand et al. 2003b). Ackerman et al. (2009) for the Kozakov suite (Bohemia) provide evidence for Pt removal, transport and precipitation from the percolating melt at low melt/rock ratios during ascent through the overlying peridotitic lithologies. These authors reported an upward positive correlation between lithophile trace element fractionation (Ce/Tb and Nb/La) and Pt/Os. Ultimate products from melt-rock reactions are small-volume alkaline silicate melts strongly enriched in volatiles from the C–H–O–S system (“carbonated melts”; Bodinier and Godard 2014; Pearson et al. 2014). Such melts have been shown to precipitate two different Type-i BMS assemblages. One assemblage is made up from bleb-like metal-rich Mss coexisting with Ni-rich pentlandite (up to 44 wt.% Ni) and isocubanite. Lorand and Alard (2001), Alard et al. (2002) and Lorand et al. (2004) interpreted this metasomatic BMS assemblage as crystallization products from immiscible metal-rich sulfide melt that exsolved along with LREE-enriched poikilitic diopside (Fig. 28a). These Mss invariably display exsolution microtextures, possibly due to abrupt cooling in a devolatilization front (Fig. 28b,c). Such “quench-textured” Mss assemblages were also identified in many metasomatized cratonic peridotites (Lorand and Grégoire 2006). The other Type-i BMS population comprises large (up to 500 µm), low-Ni (< 10 wt.% Ni) pyrrhotite grains containing abundant exsolved pentlandite (Fig. 28d). This Cu-poor (< 3 vol. % Cu) BMS occurrence was documented in Montferrier lherzolites (France) and Kerguelen dunites, which have been strongly metasomatized by volatile-rich, carbonated melts (Lorand et al. 2003a, 2004; Alard et al. 2011; Delpech et al. 2012). In thin section, the lowNi pyrrhotite shows textural evidence for sulfidation reactions of olivine and carbonated melts while corresponding bulk-rock compositions are strongly enriched in S (3 × PUM abundances and S/Se up to 104; Fig. 6), Pd (PdN/PtN up to 6 × CI-chondites) and Os (OsN/IrN up to 2.7 × CIchondrites; Lorand et al. 2004; Alard et al. 2011). The latter authors concluded that sulfidation

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Figure 28. Photomicrographs of metasomatic BMS related to small-volume melts and associated PGM. A: cluster of interconnected Type-e BMS inside a poikiloblastic clinopyroxene (Montboissier xenolith, France; Lorand and Alard 2001). B: enlargment of A showing a two-phase metalrich Mss + isocubanite-chalcopyrite intergrowth (plane-polariser reflected light). C: BSE image of a metal-rich Mss displaying rod-shaped pentlandite exsolutions (Kerguelen xenolith; south west indian Ocean; Delpech et al. 2012); D: large, pyrrhotite-rich Type-i BMS related to sulfidation reactions; note the delicate pentlandite-pyrrhotite intergrowth (Montferrier xenolith, France; Alard et al. 2011). E: BSE image of platelet Pt–Te–Bi in pentlandite + chalcopyrite intergrowth; Type-i BMS from orogenic peridotites (Lherz; France, Lorand et al. 2008b); F: BSE image of Ni arsenides (maucherite (Mc) and Pt arsenideFig. (sperrylite 28 (S))—associated with a Type-i BMS in a Semail ophiolitic harzburgite (Oman, Lorand et al. 2009).

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reactions were produced by exsolved C–H–O–S fluids that transferred the most volatile chalcophile–siderophile elements (S, Os, Pd) in the ultimate stage of carbonate metasomatism. Evidence from the oceanic mantle. Evidence of metasomatic BMS postdating adiabatic partial melting events have now been documented in oceanic peridotites, whether abyssal peridotites or ophiolitic samples. Bulk-rock S enrichments are not easy to detect because they are obscured by widespread sea-floor hydrothermal alteration and serpentinization that precipitated abundant hydrothermal BMS (Fig. 17). Uyssal et al. (2012) and Aldanmaz et al. (2012) reported bulk-rock Cu contents in Tethyan ophiolitic peridotites greatly in excess of the values expected for the degree of melting the rocks experienced (Fig. 18). Rehkämper et al. (1999a) concluded that suprachondritic PdN/PtN (up to 2.53, N = chondrite-normalized) in ODP site 895 (East Pacific Rise) and abyssal peridotites from ODP site 620 (Kane Fracture Zone, Mid-Atlantic Ridge) were achieved by segregation of immiscible sulfide liquid during adiabatic melting, in line with textural evidence of incompletely extracted partial melts. Using LAM-ICPMS analyses, coupled with a petrographic study of BMS and of bulk-rock analyses of S, Se, PGE and Au in dredged samples from the MARK area (20−24°N), Luguet et al., (2001, 2003) were able to locate this Pd excess inside Type-i Cu-Ni-rich BMS intimately associated with a late-magmatic clinopyroxene that precipitated from incompletely extracted melt fractions (Fig. 29 a–c). Enclosed BMS in secondary clinopyroxene and spinel have very high Cu sbundances, as attested to by the assemblage bornite + chalcopyrite + pentlandite (Fig. 29d; Lorand 1988; Luguet et al. 2003), being replaced by native copper by serpentinization (Seyler et al. 2007; Fig. 29b). Like basalt-hosted spinel peridotites, oceanic peridotites show evidence for mixed chalcophile–siderophile element signatures at the hand sample scale, despite their uniform pentlandite-rich Type-i BMS compositions. Two populations of CI-normalized PGE pattern in BMS assemblages have been reported from the MARK area of the Mid-Atlantic Ocean, Ligurian ophiolite and Samail harzburgites, Oman (Luguet et al. 2001, 2003; Luguet et al. 2004; Lorand et al. 2009, respectively). One exhibits the Mss-like, arch-shaped, Os, Ir, Ruenriched pattern usually ascribed to residual BMS. The other BMS population, the most Cu-enriched one, is strongly enriched in Pt, Pd and Au and display “basalt-like” patterns as reported for metasomatic/magmatic Cu–Ni-rich sulfides in continental peridotites. These two populations are not cogenetic because they show contrasting Os isotopic compositions and very different Re-depletion ages (Alard et al. 2005). At the bulk-rock scale, the corresponding samples display mixing arrays between a Cu-Pd depleted refractory end-member and variously enriched end-members (Figs. 18, 19b). Such mixing trends were discussed by Lorand et al. (2009) for Maqsad harzburgites (Samail ophiolite, Oman) and Marchesi et al. (2013) for abyssal peridotites from Hole 1274a ODP Leg 209); see also Aldanmaz et al. (2012) for Tethyan ophiolites). They are all directly related to melt-consuming reactions that take place at the top of the lithospheric mantle due to abrupt cooling (e.g., Bodinier and Godard 2014). Decreasing temperature and vanishing melt fractions favor S-saturation and segregation of immiscible sulfide melts strongly enriched in incompatible elements rejected from Mss and PGM alloys (Lorand et al. 2013). This produces the observed trends in the Cu vs. S and Pd vs. S plots of Figures 18 and 19b which are steeper compared to partial melting trends.

Platinum-group minerals and magmatic percolation of the lithospheric mantle Magmatic percolation at decreasing melt mass may generate conditions suitable for PGM precipitation either from BMS or directly from silicate melts if the latter can reach PGE saturation, as documented above for some pyroxenites. Ackerman et al. (2009) reported strong Pd and Pt enrichments controlled by a trace phase other than BMS, as suggested by Pt and Pd concentrations peaks decoupled from S (Figs. 19, 27). The two BMS precipitation processes described in continental mantle xenoliths that have seen volatile-rich (carbonated) small melt fractions have been shown to result in contrasting PGM assemblages that reflect different partitioning behavior of PGE-alloying semi-metals (Te and As). In-situ analyses of metasomatic

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Figure 29. Plane polariser reflected light phomicrographs of late-magmatic/metasomatic BMS in oceanic peridotites. A: Fine-grained rims of clinopyroxene with large Type-i BMS around a large orthopyroxene crystal. B: enlargment of A showing a coarse-grained native copper grain (Cu) inside pentlandite. C: BMS grains (pentlandite-chalcopyrite intergrowth) enclosed in cleavage planes of a coarse clinopyroxene; note the very low dihedral angles with matrix silicates, a further evidence of the high wetting capacity of metalrich sulfide melts. D: negative crystal shaped Type-e BMS inside a disseminated Al-Cr spinel of a Semail harzburgite (ophiolite of Oman). Note the Cu-Ni-rich assemblage of pentlandite-chalcopyrite-bornite. Source of data: Lorand (1988) and Seyler et al. (2007).

BMS (Fig. 30) and observations of Pt–Pd–Te microphases (moncheite-merenskyite series) in a wide range of metasomatized peridotites provide evidence that metal-rich immiscible sulfide melts 29 behavior (Ohnenstetter 1992; Luguet et preferentially incorporate Te because of its Fig. semi-metal al. 2004; Lorand et al. 2010). According to experiments by Helmy et al. (2007), sulfide melts can dissolve several thousands µg.g−1 Te at upper mantle temperatures while mantle-derived BMS contain only a few µg.g−1 Te (Lorand and Alard 2010). The Pt–Pd–Te microphases do not form until the very end of solidification, at temperatures well below that of the solidus of BMS liquid (≪ 900 °C; Barnes et al. 2006). Their subsolidus origin is demonstrated by their distribution as elongated laths inside cleavage planes in pentlandite (Figs. 4d, 28e; Lorand et al. 2010). The low-Ni pyrrhotite identified as a product of sulfidation reactions between vapor and olivine preferentially concentrates As, a volatile non-metal which is readily transported in moderately saline, high-H2S aqueous vapor phases similar to those identified in carbonated Kerguelen xenoliths (Pokrovski et al. 2002; Lorand et al. 2004; Delpech et al. 2012). Arsenic behaves quite differently in reduced or oxidizing media (Helmy et al. 2010). Under the oxidizing conditions that are generated by melt percolation in continental lithosphere, As should be dissolved into Mss or pyrrhotite structures before exsolving as arsenides on cooling. Such Pt–Pd arsenides have been detected at the margins of low-Ni pyrrhotites in both Montferrier

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(Alard et al. 2011) and Kerguelen dunites that have pervasively reacted with H2S-rich aqueous vapor (Delpech et al. 2012; Fig. 30). Witt-Eikschen et al. (2009) also detected a missing component in the bulk-rock As budget of metasomatized Eifel peridotite xenoliths which could correspond to discrete PGE arsenides not detected by conventional petrographic methods. The place where As likely plays a critical role are supra-subduction mantle wedges. This was first suggested more than thirty years ago by the identification of disseminated Ni arsenides (nickelite, NiAs; orcelite, Ni5As2; maucherite, Ni11As8) in a wide range of mantle occurrences showing evidence of metasomatism by slab-derived melt/fluid (Lorand and Pinet 1984). As demonstrated above, ophiolitic pyroxenites commonly contain micrometer-sized Pt–Pd arsenides (Edwards 1990), which were also reported from a strongly metasomatized Kamchatka peridotite xenolith (Ishimaru and Arai 2008). At Horní Bory, Ackerman et al. (2013) also described PtAs from subduction related mantle as well as NiAs. Lorand et al. (2009) described micrometric PtAs2 and maucherite in the intergranular spaces of a refractory harzburgite from the Samail ophiolite (Oman; Fig. 28f). Melt-rock reactions associated with porous flow can create regional-scale As anomalies, as is the case for the Betico-Rifean orogenic peridotite bodies of Ronda, Spain, and Beni-Bousera, Morocco (Gervilla et al. 1990). In addition to widespread disseminated Type-i NiAs phases in both peridotites and pyroxenites (Lorand 1987), an unusually As-rich PGM assemblage was described in chromitite orebodies (Torres Ruiz et al. 1996; Guttierez Narbonna et al. 2003; Gonzáles-Jiménez et al. 2013). In those mantle occurrences, melt-rock reaction at decreasing melt fractions generated a highly mobile, strongly As-enriched, volatile phase (Guttierez Narbonna et al. 2003) which ultimately produced an immiscible melt enriched in As, S, Ni, Pt, Pd, and Au (Gervilla et al. 1990, 1996). Owing to their high-melting temperature (> 1400 °C), refractory PGM of the Os–Ir–Ru–Pt system are expected to be preserved by mantle sections affected by porous melt percolation processes, as is any Type-e Mss enclosed in olivine. Both PGM and Mss are good candidates for explaining the unradiogenic Os signatures corresponding to ancient (> 1.0 Ga-old) melting events preserved in fertile lherzolites that defined the PUM arrays in Figures 6, 7, 18−20 (Burton et al. 1999, 2000; Alard et al. 2002, 2005; Meisel et al. 2003; Lorand et al. 2013 and reference therein). Mss-like residual PGE patterns have been detected in some Type-i BMS, mostly in oceanic peridotites (see above). However, no Type-e Mss inclusions (or their subsolidus decomposition products) have been reported in any of these samples, nor in orogenic lherzolites. Lorand et al. (2010) suggested that, in Pyrenean orogenic peridotites, 1000

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Figure 30. Pd vs. Te diagram for metasomatic BMS in Kerguelen peridotite xenoliths (redrawn after Delpech et al. 2012). Immiscible Cu-Ni rich sulfides carry more Te than sulfidation-related sulfides which preferentially incorporate Pd, Pt, S, As, and Os. Fig. 30

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Laurite PtIrOs PtTeBi PdTeBi Pt (native, arsenides sulfarsenides, sulfides)

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spinel/garnet pyroxenites

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Figure 31. Evolution of PGM assemblages in the main lithologies of Pyrenean peridotites (after Lorand et al. 2010 and unpublished data).

instead of Mss, these are residual refractory PGM (laurite and Pt–Ir–Os alloy) inherited from harzburgitic protolith (cf. Luguet et al. 2007) that have transferred ancient Re-depletion ages to the lherzolitic BMS phases. Figure 31 shows that the PGM assemblage gradually evolves from the harzburgites to the lherzolites, with decreasing modal proportions of laurite and, to a lesser extent, Pt– Ir–Os alloys, balanced by increasing modal abundances of Pt–Pd–Te–Bi phases which become the only PGM so far recognized in pyroxenites that represent silicate melt conduits. Such an evolution trend has been interpreted as reflecting progressive dilution of the residual PGM inherited from refractory residual protolith into metalrich metasomatic sulfide melts precipitated by silicate melts en-route to the surface (Lorand et al. 2010).

Separate evidence for secondary entrapment of refractory PGM by metasomatic BMS were provided by Griffin et al. (2002) who found Type-e Mss in Udachnaya (Siberia) olivine macrocrysts to be mixtures between Pt–Ir–Os alloys of residual origin and younger cumulate Mss. Likewise, in Kerguelen harzburgites, Delpech et al. (2012) reported metalrich Mss compositions reminiscent of Os–Ir–Ru-rich PGM (Fig. 32). Bockrath et al. (2004b) and Ballhaus et al. (2006) pointed out that, in S-oversaturated systems, sulfide melts will preferentially wet discrete PGE nuggets: the huge solubility of PGE into BMS will dissolve any captured PGM nugget, except perhaps the Pt–rich alloys that show the lowest solubility in Mss (Peregoedova et al. 2004). However, Lorand et al. (2010) documented several shapes for laurites in Pyrenean lherzolites, either euhedrally shaped, concentrically zoned (Os-rich rim) crystals at the contact between BMS and silicates (Fig. 33 a,b) or rod-shaped 20 µm-long platelets inside pentlandite (Fig. 33d). Laurite platelets are likely subsolidus exsolution products; the interpretation that euhedral crystals may be primary precipitates from Ru-rich sulfide melt or undissolved crystals remains ambiguous. Regardless of the way in which they have survived in fertile mantle rocks, such micrometric PGM of residual origin are now considered to account for the large dispersion of Os isotopic compositions of ancient melt depletion events in refertilized peridotites, as well as in melt vein-conduits (e.g., chromitites, pyroxenites; Fonseca et al. 2012; Gonzáles-Jiménez et al. 2013).

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magmatic Mss + Pt-Ir-Os alloys (Siberian olivine macrocrysts, Griffin et al., 2002) Metasomatic Mss (Kerguelen dunite xenolith; Delpech et al, 2012)

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Figure 32. CI-chondrite normalized PGE patterns of mixed magmatic Mss + PGM grains (LA-ICPMS analyses). Note the close similarities between Kerguelen metasomatic Mss and laurite patterns. Fig. 32

A

B

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5 μm

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D

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Figure 33. BSE images showing the great diversity of laurite–erlichmanite crystals in Pyrenean peridotites (white). A: octahedral sectioning of an external laurite granule in contact with Type-i pentlandite; lowbrightness spongy material is secondary magnetite. B: euhedral laurite inclusion (La) in the core of a massive Type-i pentlandite bleb; C: polycrystalline inclusions inside pentlandite and in a former limit between two sub-grains; D: en-echelon rods looking like exsolution network inside Type-i pentlandite.

Fig. 33

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Chalcophile–siderophile elements are trace to ultra-trace elements that may be easily mobilized from their host minerals as a function of sulfur-saturation, redox conditions, pressure, temperature, fugacity of sulfur, and silicate-melt compositions. These elements are hosted in accessory BMS in the convecting upper mantle, at least above the level of Fe–Ni metal saturation. Mantle samples brought to the surface of the Earth surface contain only subsolidus decomposition products of high-temperature BMS, as well as a large range of trace phases containing siderophile/chalcophile elements as major elements. Some of these phases are exsolution products (e.g., Au, Pd–Te–Ni); other are products from the various petrogenetic processes operating inside the convecting mantle or in the rigid lithosphere. By consuming sulfur, adiabatic partial melting processes transfer the chalcophile–siderophile element inventory of mantle residues from the BMS phases to refractory Fe–Ni monosulfides, and then to refractory PGM (Os–Ir-Ru alloys/sulfides and Pt–Ir–(Os) alloys) which may account for the bulk-rock HSE budget. When S is totally eliminated, some elements (e.g., Se and Te, Pt and Pd) that are usually considered to be proxies for sulfur-saturated conditions, display decoupled behavior, depending on which PGM (alloys vs. sulfides) are formed. Basaltic melts flowing through vein conduits may precipitate BMS at high-pressure (> 2 GPa), where the solubility of S in silicate melt is low. Cumulate pyroxenites have preserved basalt-like chalcophile/siderophile element inventories inside pyrrhotite-dominated BMS assemblages. Websterites/wehrlites show evidence of assimilation of peridotitic wallrock by basaltic melts, which precipitated pentlandite-dominated Os–Ir–Ru–Ni-rich hybrid BMS. The saturation of PGM occurs when enough Te and/or As ligands are available in the mixed melt to collect PGE. The distribution of chalcophile–siderophile elements in the lithospheric mantle is intimately related to events of regional-scale melt migration. The overall depletion in chalcophile–siderophile elements, as shown by alkali-basalt hosted xenoliths, is related to partial to total consumption of BMS by porous flow percolation operating at high melt/rock ratios. Such reactions are balanced by precipitation of immiscible metal-rich Cu-Ni sulfide melts once silicate melts are consumed (via melt-rock reaction under decreasing temperature gradients). This metasomatic BMS precipitation has been shown to rejuvenate the inventory of the most incompatible chalcophile–siderophile elements (Cu, Te, Se, Pd, Pt, Au) in both the subcontinental and the suboceanic mantle lithosphere. Sulfur-rich fluids exsolved from carbonate/silicate melts play a major role in the redistribution of volatile chalcophile– siderophile elements (Pd, Os, S, As). Arsenic may substitute for S and Te as a PGE-alloying semi-metal, especially in suprasubduction settings. In spite of the long-term, complex evolution of the upper mantle, refractory PGM (Os–Ir–Ru alloys/sulfides; Pt–Ir–(Os) alloys) are long-lived minerals carrying a record of ancient melt depletion event(s) in seemingly young, so-called “fertile” lherzolites from continental and oceanic mantle. Whether hosted in metasomatic BMS or preserved as restitic crystals, their presence in fertile mantle samples that should not contain any high-temperature PGM is puzzling. Systematic reassessment of PGM microtextures (direct precipitation crystals, restitic PGM, quench crystals, exsolutions) using cutting-edge techniques (X-ray tomography, high-resolution TEM) is still required for unravelling such mantle heterogeneities carried by trace phases.

ACKNOWLEDGMENTS This chapter was greatly improved through thoughtful reviews of James Brenan, Lukas Ackerman, Bill Griffin and an anonymous reviewer. It also benefited from editorial comments by James Day and Jason Harvey who are warmly thanked for their help with the text.

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Reviews in Mineralogy & Geochemistry Vol. 81 pp. 489-578, 2016 Copyright © Mineralogical Society of America

Petrogenesis of the Platinum-Group Minerals Brian O’Driscoll School of Earth, Atmospheric & Environmental Science The University of Manchester Williamson Building, Oxford Road Manchester M13 9PL UK [email protected]

José María González-Jiménez Department of Geology and Andean Geothermal Center of Excellence (CEGA) Facultad de Ciencias Físicas y Matemáticas Universidad de Chile Plaza Ercilla #803, Santiago de Chile Chile [email protected]

INTRODUCTION The platinum-group minerals (PGM) are a diverse group of minerals that concentrate the platinum-group elements (PGE; Os, Ir, Ru, Rh, Pt, and Pd). At the time of writing, the International Mineralogical Association database includes 135 named discrete PGM phases. Much of our knowledge of the variety and the distribution of these minerals in natural systems comes from ore deposits associated with mafic and ultramafic rocks and their derivatives (see also Barnes and Ripley 2016, this volume). Concentrations of PGM can be found in layered mafic–ultramafic intrusions. Although they don’t typically achieve ore grade status, suprasubduction zone upper mantle (preserved in ophiolite) lithologies (i.e., chromitite [> 60 vol.% Cr-spinel], pyroxenite) characteristically host a diversity of PGM assemblages as well (Becker and Dale 2016, this volume). Occurrences of the PGM in layered intrusions, ophiolites, and several other important settings will all be described in this review. In keeping with the general theme of this volume, the focus of this chapter is on relatively high-temperature (magmatic) settings. This is not a straightforward distinction to make, as PGM assemblages that begin as high-temperature parageneses may be modified at much lower temperatures during metamorphism, hydrothermal processes or surficial weathering (e.g., Hanley 2005). However, the vast majority of the published literature on PGM petrogenesis is based on occurrences from magmatic environments, an understandable bias given the importance of the major ore deposits that occur in some layered mafic–ultramafic intrusions, for example. For that reason, the emphasis of this review will be on high-temperature magmatic settings, with the understanding that lower temperature (sub-solidus; < 600 °C) processes can modify primary PGM assemblages. The geochemical behavior of the platinum-group elements (PGE) in magmatic settings is highly chalcophile and not, as might be expected, highly siderophile. This is because most terrestrial magmatic systems are relatively oxidized, such that native Fe is not stable. A consequence of this is that the PGE have very high partition coefficients for sulfide in magmatic environments. From the high-temperature perspective, the degree of mantle melting and the manner in which magma fractionation proceeds, 1529-6466/16/0081-09$10.00

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whether or not sulfide separation and fractionation has occurred, and the way in which latestage (metasomatic) alteration has affected the rocks, if at all, are each potentially important processes in determining the PGM assemblage that forms. In particular, the relevance of the separation, accumulation, and crystallization of sulfide liquid (and/or arsenide liquid) in silicate magma systems, with respect to the timing of PGM crystallization, is a crucial issue. The PGE constitute most of the highly siderophile elements (HSE; Os, Ir, Ru, Rh, Pt, Pd, Re, Au). While the HSE have and are being applied to addressing first order problems in Earth and planetary evolution, as outlined throughout this volume, outstanding concerns remain regarding the small-scale siting of the HSE in rocks. Specifically, the control exerted by the different types of PGM on bulk concentrations and relative distributions of the PGE in rock samples is presently an avenue of active investigation. Of critical but relatively unknown importance is the interplay between the high-temperature processes listed above, and the way in which these operate to produce a particular PGM assemblage. The PGE are generally considered to be relatively resistant to geochemical processes that fractionate and modify the lithophile elements in silicate and non-silicate rocks (Walker 2009). However, the petrogenetic links between the PGM and other non-silicate minerals (i.e., oxides such as chromite, base-metal sulfides, arsenides, and antimonides) remain unclear in many natural samples. One reason for this is that some of the PGE in a sample may hypothetically exist in solid solution in minerals such as sulfides or oxides, whilst others form discrete PGM. Progress in unravelling these issues has been made with the advent of high precision microbeam techniques, such as LA-ICP-MS, where trace element concentrations and isotopic information (e.g., 187Os/188Os) can be collected from individual PGM, providing new information on their chronology and petrogenesis (Pearson et al. 2002; Malitch et al. 2003; Ahmed et al. 2006; Shi et al. 2007; Marchesi et al. 2011). Unfortunately, in most relevant natural materials, PGM grain sizes are too small (< 1 μm) for such analyses and uncertainties surrounding PGM crystallization processes persist. The principal goals of this contribution are twofold: (1) to provide an overview of the principal PGM assemblages/parageneses in natural rock samples, including those from terrestrial as well as extraterrestrial environments. This includes highlighting, where possible and/or relevant, how these control whole-rock PGE budgets; (2) to provide the reader with a flavour of the processes (or combination of processes) responsible for the crystallization of PGM in different petrological environments. As noted above, this chapter is principally concerned with synthesizing observations on PGM occurrences that have been welldocumented, so the emphasis is on mid-upper crustal mafic–ultramafic intrusions (including layered intrusions) and ophiolitic mantle. However, studies of secondary concentrations of PGM that are considered to be derived from high-temperature source material (i.e., placer deposits associated with ophiolites and zoned Uralian–Alaskan–Aldan-type complexes) have also contributed valuable information to our understanding of high-temperature PGM petrogenesis. For this reason, an appendix that summarizes some of the major findings of placer deposit studies is included with this chapter. The reader is also referred to reviews by Cabri and Feather (1975), Cabri (1981a, 2002) and Brenan and Mungall (2008).

PHASE RELATIONS AND ORIGIN OF THE PGM Chemical properties of the PGM The PGE are Group VIII transition metals, together with Fe, Cu, and Co. The PGE have melting points above that of Fe (1665 K), ranging from 1828 K for Pd to 3306 K for Os (Table 1). In their elemental form, Rh, Pd, Ir, and Pt (arranged in increasing Z) are cubic in symmetry (fcc), with Os and Ru being hexagonal (hcp). The electron configurations of the PGE are given in Table 1. They can exist in multiple valence states (0 to +8) and have

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Table 1a. A compilation of some of the common natural PGM discussed in this review, organized according to mineral composition and complexity of bonding with other elements. Melting points, crystal structure and electron configuration details are also provided for the PGE. This table has been adapted and updated from Cabri (1981). Osmium

Iridium

Ruthenium

Melting T (K)

3306

2739

2607

Crystal structure

hcp

fcc

hcp

[Xe]4f145d66s2

[Xe]4f145d76s2

[Kr]4d75s1

Erlichmanite (OsS2)

Xingzhongite [(Pb,Cu,Fe)(Ir,Pt,Rh)2S4]

Laurite (RuS2)

Electron configuration S

Kashinite (Ir,Rh)2S3 Cuproiridsite CuIr2S4 As

Omeiite [(Os,Ru)As2]

Iridarsenite [(Ir,Ru)As2]

Anduoite [(Ru,Os)As2] Ruthenarsenite [(Ru,Ni)As]

AsS

Osarsite [(Os,Ru)AsS]

Irarsite [(Ir,Ru,Rh,Pt)AsS]

Ruarsite (RuAsS)

Native metal

Osmium

Iridium

Ruthenium

PGE Alloys

Ir–Os

Os–Ir Pt–Ir Ru–Os–Ir

Os–Ir–Ru

Chengdeite (Ir3Fe)

bonding behavior influenced by their overlapping d-orbitals. Their d-orbitals, in particular, are important for the formation of PGM because they consist of unpaired electrons that can form metal–metal bonds with other PGE, or covalent bonds with electron acceptors such as S. Thus, despite their differences in symmetry, the PGE can exist in solid solution with one another (Cabri 1981a), as well as combining with other siderophiles (e.g., Fe), chalcogenides (e.g., S, Te), and semi-metals (e.g., As, Sb). Native metals (minerals) composed of the PGE do not typically undergo phase transitions within temperature and/or pressure ranges of geological significance (Cabri 1981a) and the principal variable property, in addition to melting temperature (see Table 1), is the compositional range of solid solutions. Berlincourt et al. (1981) provide a detailed synthesis of the phase relations of the PGE. Examples of unary phases include the complete solid solution of Ir, Rh, Pd and Pt with Ni, and the complete solid solution of Os with Ru. More commonly, binary phases are formed between the individual PGE and each other, and elements including As, Bi, Cu, Fe, Hg, In, Ni, Pb, S, Sb, Se, Sn, and Te (see Table 1). Examples of known natural ternary and quaternary phases are also listed in Table 1. The PGE are characterized by their extreme affinity for metal phases, instead of oxides or silicates, and together with Au and Re, are referred to as the highly siderophile elements (HSE); the distribution coefficients of the PGE between metal and silicate are greater than 104 (Jones and Drake 1986; Holzheid et al. 2000; Fortenfant et al. 2003). The siderophilic character of the PGE is manifested by numerous occurrences (see also Brenan et al. 2016, this volume; Lorand and Luguet 2016, this volume), e.g., the high concentrations of PGE in iron meteorites, Fe–Ni metal alloys found in chondrites and the occurrence of many metal alloys as PGM (Table 1). The PGE may also have a tendency to exhibit chalcophile behavior, readily bonding with S, As, and other Group Va and VIa ligands.

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Table 1b. A compilation of some of the common natural PGM discussed in this review, organized according to mineral composition and complexity of bonding with other elements. Melting points, crystal structure and electron configuration details are also provided for the PGE. This table has been adapted and updated from Cabri (1981).

Melting T (K) Crystal structure Electron configuration S

Rhodium

Platinum

2237

2041

fcc

fcc

[Kr]4d85s1

[Xe]4f145d96s1

Prassoite (Rh17S15)

Braggite (Pt,Pd,Ni)S

Bowieite (Rh,Ir,Pt)2S3

Cooperite (PtS)

Cuprorhodsite (CuFe)Rh2S4

Malanite [Cu(Pt,Ir,Co)2S4]

Te

Moncheite [(Pt,Pd)(Te,Bi)2] Maslovite (PtBiTe)

As AsS

Sperrylite (PtAs2) Hollingworthite [(Rh,Pt,Pd)AsS]

Platarsite [(Pt,Rh,Ru)AsS]

Daomanite (CuPtAsS2) Sb

Genkinite [(Pt,Pd)4Sb3] Geversite [Pt(Sb,Bi)2] Stumpflite [Pt(Sb,Bi)]

Bi

Insizwaite [Pt(Bi,Sb)2]

BiTi Native metal

Maslovite (PtBiTe) Rhodium (Rh)

PGE Alloys

Platinum (Pt) Hongshiite (PtCu) Tetraferroplatinum (PtFe) Niggliite (PtSn) Rustenburgite [(Pt,Pd)3Sn] Isoferroplatinum (Pt3Fe) Tulameenite (Pt2FeCu)

Extraterrestrial occurrences of the PGM The distribution of the PGE during condensation of the solar nebula was governed by their relatively low vapor pressure (Palme 2008; Yokoyama and Walker 2016, this volume). Five of the six PGE are refractory, with condensation temperatures higher than that of Fe–Ni alloy. Only Pd, with the lowest melting temperature of the PGE, is non-refractory, condensing in solid solution with Fe–Ni. Platinum-group minerals may be found in carbonaceous chondrites (e.g., Fig. 1a,b), where they occur as refractory metal nuggets (RMN) that have traditionally been considered to be amongst the earliest condensates in the Solar System (Palme 2008). The RMN were first documented in calcium–aluminium inclusions (CAIs) in the CV3 chondrite, Allende (Palme and Wlotzka 1976; Wark and Lovering 1976). The RMN contain high but variable concentrations of the PGE (except Pd), together with Mo, W, Ni, and Fe. The presence of significant Mo

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Table 1c. A compilation of some of the common natural PGM discussed in this review, organized according to mineral composition and complexity of bonding with other elements. Melting points, crystal structure and electron configuration details are also provided for the PGE. This table has been adapted and updated from Cabri (1981). Palladium Melting T (K)

1828

Crystal structure

fcc [Kr]4d10

Electron configuration S

Vysotskite [(Pd,Ni,Pt)S]

Vasilite [(Pd,Cu)16(S,Te)7]

Se

Oosterboschite [(Pd,Cu)7Se3]

Palladseite [Pd17Se15]

Te

Keithconnite [Pd3–xTe(x = 0.14–0.43)]

Merenskyite [(Pd,Pt)(Te,Bi)2]

Kotulskite [Pd(Te,Bi)]

Telluropalladinite (Pd9Te4)

Telargpalite [(Pd,Ag)3Te] As

Atheneite [(Pd,Hg)3As]

Palladoarsenide (Pd2As)

Guanglinite [Pd11Sb2As2]

Stillwaterite (Pd8As3)

Majakite (PdNiAs) As,Sb

Arsenopalladinite [(Pd8(As,Sb)3)]

Palladodymite [(Pd,Rh)2As]

Isomertieite (Pd11Sb2As2)

Mertieite-II [Pd8(Sb,As)3]

Sb

Stibiopalladinite (Pd5Sb2)

Sudburyite [(Pd,Ni)Sb]

SbAs

Isomertieite (Pd11Sb2As2)

Mertieite-II [Pd8(Sb,As)3]

SbTe

Borovskite (Pd3SbTe4)

Testibiopalladite [PdTe(Sb,Te)]

Froodite (PdBi2)

Sobolevskite (PdBi)

Padmaite (PdBiSe)

Urvantsevite [Pd(Bi,Pb)2]

Bi

Polarite [Pd,(Bi,Pb)] BiTe

Michenerite ([Pd,Pt)BiTe]

AsBi

Palladobismutharsenide [Pd2(As,Bi)]

Pb

Plumbopalladinite (Pd3Pb2)

Hg

Potarite (PdHg)

HgTe

Zvyagintsevite (Pd3Pb)

Temagamite (Pd3HgTe3)

Native Metal

Palladium (Pd)

PGE Alloys

Skaergaardite (PdCu)

Nielsenite (PdCu3)

Atokite [(Pd,Pt)3Sn]

Cabriite Pd2SnCu

Palarstanide [Pd8(Sn,As)3]

Taimyrite [(Pd,Cu,Pt)3Sn]

Paolovite [Pd2Sn]

Stannopalladinite [(Pd,Cu)3Sn2]

(up to 40 wt.%) and W in the RMN distinguishes them from terrestrial PGM and was originally interpreted as reflecting their formation via condensation from the solar nebula under relatively reducing conditions (possibly several orders of magnitude lower than the Iron–Wüstite fO2 buffer; Simon et al. 2005), in the temperature range 1600–1400 K at ∼ 10–4 bar (Palme and Wlotzka 1976; Campbell et al. 2001; Berg et al. 2009). This is because if the RMN were residues of extensive heating and vaporization, W and Mo would likely be lost as volatile oxides (Palme 2008).

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Figure 1. (a) Back scatter electron micrograph depicting Pt–Ru–Fe nugget (bright phase at image center) included in compositionally zoned melilite (Allende CV3). Image courtesy of Dr Glenn MacPherson, Smithsonian Institution (USA). (b) Numerous refractory metal nuggets (white spots) enclosed in a type B1 CAI from the Allende meteorite. The equant dark gray minerals are spinel and the light gray phase is melilite. Image courtesy of Dr Daniel Schwander, University of Mainz (Germany).

In documenting the occurrences of CAI-hosted RMN in the CV3 meteorites Allende and Leoville, El Goresy et al. (1978) noted two principal textural occurrences. The first is as isolated Pt-dominated nuggets, the second as ‘Fremdlinge’, a term introduced by El Goresy et al. (1978) to describe a complex opaque mineral-bearing paragenesis. In addition to RMN, Fremdlinge contain Fe–Ni metal, sulfides (e.g., Fe, Ni-dominated sulfides, MoS2, WS2), Ca-phosphates, and silicates (including volatile-bearing silicates such as nepheline and sodalite), oxides, and tungstates. It was originally speculated that Fremdlinge might have a presolar origin (El Goresy et al. 1978), but this has been subsequently shown to be unlikely, for the Allende examples at least, which yielded solar isotopic ratios of Mg, Fe, Mo, Ru, and W (Hutcheon et al. 1987). Isotopic data for the texturally isolated RMN are more limited (i.e., one for Ru and another for Os; Hutcheon et al. 1987 and Berg et al. 2009, respectively), but no isotopic anomalies were found. Palme et al. (1994) argued for condensation of the RMN as a single refractory metal alloy phase. They interpreted Fremdlinge as the product of oxidation and sulfidation of initially homogeneous grains, with the associated molybdenite (MoS2) and scheelite (CaWO4) being considered as secondary metasomatic products. A contrasting model suggests that since refractory metals occur in one of three different crystal structure types (hcp alloys of Ru and Os, bcc alloys of Mo and W and fcc alloys of Ir, Fe, and Ni), Fremdlinge may represent the product of subsequent mixing of these components (Sylvester et al. 1990). Additional studies of meteorites

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including the CM2 Murchison, CV3 Allende and CM Paris chondrites (Blander et al. 1980; Berg et al. 2009; Schwander et al. 2011, 2012, 2013; Harries et al. 2012; Hewins et al. 2014) revealed that they are all observed or inferred to be associated with CAIs. Harries et al. (2012) carried out detailed crystallographic and compositional studies of RMN from the Murchison meteorite and argued that their typically homogeneous nature along with their well-defined crystal habits and crystal structures supported equilibrium condensation in the solar nebula, consistent with the proposals of Palme et al. (1994) and Berg et al. (2009). In a detailed textural and mineralogical study of CAIs from the Paris meteorite, Hewins et al. (2014) report that RMN occur exclusively inside CAI or attached to various disrupted CAI minerals, e.g., hibonite ((Ca,Ce)(Al,Ti,Mg)12O19), perovskite (CaTiO3), and spinel. There is an interesting compositional variation in the RMN depending on the host mineral. Specifically, the most refractory RMN (Os and Ir ≈ 47.4 and 37.2 wt.%, respectively) are observed in hibonite. The RMN in the perovskite + Al-spinel host are Os–Ir–Mo–Ru alloys, whilst those in spinel (remote from perovskite) are enriched in Pt (i.e., up to 29 wt.%) and Rh. The RMN observed hosted in forsterite are Pt–Fe-rich alloys. The largest CAI studied by Hewins et al. (2014) contained tens of evenly distributed RMN (100 nm to < 1 µm in diameter) that appear to occur preferentially in Y-rich perovskite. The experimental data of Schwander (2014) suggests that the condensation model may not be universally appropriate for RMN, and indicates that a silicate liquid enriched in refractory metals might have precipitated these phases (see also Cottrell and Walker 2006). This notion is supported by Rudraswami et al. (2014), who observed that RMN in cosmic spherules (collected from the Indian Ocean seafloor) formed during melting and oxidation on atmospheric entry. Croat et al. (2013) argued that if the RMN are the products of high-temperature condensation processes, they might also be expected to be present as inclusions in pre-solar graphite. Samples of the carbonaceous chondrites Murchison (CM2) and Orgueil (CI) yielded four Os-, Ru-, and Mo-rich RMN inclusions from four different pre-solar graphites. The RMN themselves are too small (30–50 nm) for isotopic analyses. However, Croat et al. (2008) presented C isotope data that indicate a pre-solar origin for the graphite, which in conjunction with the textural configuration suggests that the RMN inclusions are pre-solar too. Croat et al. (2013) consequently argued against the assumption that all isolated RMN in carbonaceous chondrites are necessarily solar in origin. Several studies have reported the presence of RMN in martian meteorites, and attributed these to relatively late-stage processes (with respect to Solar System formation). For example, Lorand et al. (2012) have documented Fe–Os–Ir–(Ru) alloys several hundred nm in diameter associated with Ni-poor troilite in the martian meteorite NWA 2737, which they ascribed to a reduction of S in Fe-sulfides, driven by sulfur degassing. Lorand et al. (2014) also reported Ir– (Os)–As–S- and Os–S-rich compounds in the martian regolith breccia (NWA 7533), that may represent remnants of heavily bombarded ancient martian crust. The latter studies indicate that processes operating on the crusts and mantles of other planets in the Solar System are capable of forming PGM too. More generally, it is worth noting that the parent bodies of many differentiated meteorites are considered to have lower fO2 than Earth, such that Fe–Ni metal may be present for the PGE to partition into (Day et al. 2016, this volume). This might result in significantly different PGM assemblages than those found on Earth.

Origin of the terrestrial PGM: Mantle melting, metasomatism, and metal transfer The crustal abundance and distribution of the PGE in the present-day Earth are the result of a combination of processes that occurred during mantle melt extraction and the geological evolution of plate tectonics. Specifically, the way in which mantle melting has mobilized and transported the PGE is a significant factor in accounting for the distribution of these elements in natural (accessible) geological materials (see also Lorand and Luguet 2016, this

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volume). Despite the overwhelming preponderance of silicate minerals, particularly olivine (55–90 vol.%) in the upper mantle (Lorand et al. 2008), there is a breadth of experimental and geochemical evidence suggesting that silicates do not significantly contribute to the overall PGE budget of the mantle (Hart and Ravizza 1996; Burton et al. 1999; Harvey et al. 2010). For example, even though the partitioning of Ir into olivine (as Ir2+; Brenan et al. 2005) has been experimentally demonstrated, only trace concentrations (0.03–0.1 ppb; ng.g–1) of the PGE have been measured in mantle olivine (Lorand et al. 2008). It has been estimated that < 10% of the PGE abundances of typical mantle peridotites is controlled by their constituent silicate phases (Handler and Bennett 1999; Burton et al. 2002; Lorand et al. 2008). However, it is possible that the fO2 conditions under which the substitutions above were recorded are significantly higher than those governing mantle melt production (∆FMQ = +1.4 to 5.4 vs. –1 to –3; Ballhaus 1995). The compatibility of the PGE in olivine (i.e., Ir and Ru; Mungall and Brenan 2014) thus continues to be considered as viable, highlighting that the role of silicate in hosting mantle PGE requires further investigation. Accessory chromite is also a refractory phase during mantle melting and has been considered as a host for the PGE. The substitution of Ru3+ (or Ru4+) and Rh3+ into chromite has been proposed (Capobianco and Drake 1990; Righter et al. 2004) and absolute concentrations of Ir, Ru, and Rh in the range of tens of ppb in chromite have been reported (Capobianco and Drake 1990; Righter et al. 2004; Brenan et al. 2005, 2012). In addition, concentrations of ∼ 500 ppb Ru have been measured in komatiite chromite (Locmelis et al. 2011). However, as chromite typically occurs as an accessory phase (1–2 vol.%) in mantle peridotites, it is probably responsible for only a few percent of the bulk rock mantle PGE budget (Carlson 2005; Luguet et al. 2007). It is therefore unlikely that chromite is the dominant control on PGE in the upper mantle. The upper mantle, particularly the sub-continental lithospheric mantle (SCLM) also commonly contains base-metal (Fe–Cu–Ni) sulfides as accessory phases. In the absence of metal phases, the PGE are chalcophilic and sulfides are widely considered to be significant carriers of PGE in the mantle (Cabri 1981a; Mitchell and Keays 1981; Barnes et al. 1985; Alard et al. 2002; Carlson 2005; Lorand et al. 2008). The abundances of the PGE in common mantle sulfide phases are typically at the ppm (µg.g–1) level, three orders of magnitude higher than typical bulk rock mantle peridotite concentrations (Morgan 1986; Pattou et al. 1996; Burton et al. 2002; Lorand et al. 2008). Sulfide melt–silicate melt partition coefficients of 103–106 have been reported for the PGE (Fleet et al. 1993; Roy-Barman et al. 1998; Ballhaus et al. 2006; Mungall and Brenan 2014), signifying that it is the behavior of sulfide during mantle melting that exerts the greatest control on PGE concentration and distribution. Thus, the degree of partial melting and tectonic setting must exert considerable influence on whether sulfides and their PGE are retained in the mantle or not. High degrees of mantle melting, such as are thought to have occurred during the production of Archean komatiites and which are also manifest in some supra-subduction zones in the geological record, have the potential to liberate mantle sulfides originally held as Fe-rich monosulfide solid solution (mss). Experimental work indicates that mss is the first phase to crystallize from a sulfide liquid, followed by intermediate solid solution (iss), which crystallizes from a more fractionated Cu-rich sulfide liquid. By contrast, the low degrees of partial melting that are typical of mid-ocean ridge basalts (MORB) would favor the retention of sulfides in the mantle residue, as the melts produced are already S-saturated. Thus, MORB magmas tend to be PGE-poor (Bézos et al. 2005). However, these issues are further complicated by the fact that mantle peridotites have been observed to contain multiple sulfide populations, separable on the basis of microstructural, chemical, and isotopic criteria (see also Harvey et al. 2016, this volume). For example, numerous studies (e.g., Alard et al. 2000; Lorand and Alard 2001; Pearson et al. 2002; Bockrath et al. 2004a,b; Harvey et al. 2011) have reported Os–Ir–Ru and Rh–rich mss inclusions in olivine with depletions in the PPGE (i.e., Pt and Pd). The formation of these sulfides has been attributed to separation

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of an immiscible sulfide melt fraction (i.e., droplets, tens of µm in diameter) during an early silicate melt extraction event (Holzheid 2010). The interstitial and intergranular regions of the peridotites may contain a second population of Pd–Re-enriched sulfides (Alard et al. 2002; Bockrath et al. 2004a,b; Harvey et al. 2011), whose occurrence has been attributed to relatively late-stage metasomatism. Unlike the former variety, these interstitial sulfides are more prone to subsequent mobilization and processing, as they are not ‘armored’ by olivine crystals. Building on the work of Alard et al. (2002), Harvey et al. (2011) carried out a detailed study of Kilbourne Hole spinel lherzolites and showed that the secondary metasomatic sulfides readily contribute their enhanced PPGE + Re (and consequently radiogenic 187Os) to partial melts. By contrast, it is only via protracted and high degrees of partial melting that the silicate-hosted sulfides will be dissolved in the silicate melt, providing a mechanism by which materials enriched in the IPGE (i.e., Os, Ir, and Ru) can be included in mantle melting. Experimental studies have suggested that the PGE are not equally soluble in base-metal sulfides (Mackovicky et al. 1986; Barnes and Francis 1995; Tredoux et al. 1995; Mungall et al. 2005). The following increase in order of the PGE partition coefficients into sulfide liquid has been suggested: Au~Os~Ir~Ru < Pt< Rh< Pd (Barnes and Francis 1995; Tredoux et al. 1995). Mackovicky et al. (1986) reported that the sulfide melt precursor of pentlandite [(Fe,Ni)9S8] and Cu-sulfide can accommodate up to 15 wt.% Pt. Borisov and Palme (2000) showed that the PGE content of mantle-derived magmas crystallizing at or below FMQ is still high enough for them to be saturated in Ru + Ir-bearing alloys, and Pt–Fe saturation is also likely (see also Mungall and Brenan 2014). Ballhaus (1995) suggested that in the shallow fertile mantle, the bulk of the IPGE reside in mss, this being governed by high sulfide/silicate partition coefficients. However, he suggested that following partial melting and the removal of sulfide with the melt, metal phases might be stabilized with decreasing fO2 and fS2. Support for this comes from the detailed study of highly depleted (sulfide-free) spinel harzburgites from the Lherz massif by Luguet et al. (2007), which showed that whilst the major minerals (silicates and oxides) account for up to 30% of the bulk rock PGE budget, intergranular µmto-sub-µm scale Ru–Os ± Ir sulfides and Pt–Ir ± Os alloys account for 50–100% of the PGE. Luguet et al. (2007) interpreted these alloys and PGE-sulfides as being residual phases that developed during complete extraction of base-metal sulfides by up to 24% partial melting of the protolith (see also Fonseca et al. 2012). The presence of PGM is not restricted to the SCLM. Luguet et al. (2003) reported Os–Ru alloys in Mid-Atlantic Ridge (MAR) sulfidepoor abyssal peridotites. Further evidence for the presence of PGM in oceanic peridotites comes from studies of ophiolite dunites, pyroxenites and chromitites (see below). It has been shown that at least some of the PGM population developed in ophiolite chromitites may be ‘recycled’ as xenocrystic grains during melt transport in the conduits that form the chromitite (Gonzalez-Jimenez et al. 2013a; Yang et al. 2014). The transfer of the PGE from the mantle to crust is manifested in the variety of settings in which these precious metals are enriched. The following six sections are concerned with outlining the characteristics of the principal terrestrial settings of PGM mineralization and the diversity of the PGM observed therein. It is not possible in a review such as this to discuss every single reported example of PGM mineralization. However, we have endeavoured to be as thorough as possible with the details of important examples of ophiolites, layered mafic– ultramafic intrusions, subcontinental lithospheric mantle, Uralian–Alaskan–Aldan Complexes and Ni–Cu–PGE–sulfide deposits. Examples of less conventional PGM mineralization are also included, for breadth, but in keeping with the focus of this volume, exclusively low temperature PGM mineralization environments are not discussed in detail.

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O’Driscoll & González-Jiménez PGM IN LAYERED MAFIC–ULTRAMAFIC INTRUSIONS

Layered mafic–ultramafic intrusions (LMI) are generally taken to represent solidified magma chambers, and are typically the focal points of significant magmatic activity in anorogenic tectonic settings. The type examples of LMI are therefore associated with large igneous provinces (LIPs), which are considered to have developed during periods of significant continental break-up in the geological record. There are arguably fewer significant PGE deposits (in terms of variety of PGM) associated with LMI than there are with ophiolite chromitites and mantle peridotites (Barnes and Ripley 2016, this volume). However, there are an exceptional few LMI that contain the greatest concentrations of PGE on Earth. Indeed, the largest known layered intrusion, the Bushveld Complex (South Africa) contains > 75%, > 50% and > 80% of the world’s exploited Pt, Pd, and Rh, respectively (Naldrett 2004; Mungall and Naldrett 2008). Layered intrusions can be sub-divided on the basis of whether they represent open or closed system magma chambers. The most economically significant LMI in terms of PGE concentration predominantly contain cumulate sequences that preserve evidence for open system behavior, including the development of stratiform chromitite deposits. Type examples of such chromitites occur in the Bushveld Complex and the Stillwater Complex (Montana, USA). In both of the latter intrusions, the PGE are frequently, but not ubiquitously, associated with these chromitite seams (see below). Less commonly, closed-system LMI may also contain significant PGE enrichment (e.g., the Platinova Reef of the Skaergaard Intrusion, Greenland). Consideration of the styles of PGE mineralization in LMI is important, in light of the issues of mantle melting discussed above. Layered intrusions are mostly the result of large degrees of partial melting of mantle that underlies continental crust, rather than the oceanic mantle sampled by ophiolite peridotites. Thus, they offer a complementary perspective on the way in which the PGE are transferred from mantle to crust relative to that observed in oceanic peridotites. Broadly speaking, the whole-rock PGE patterns associated with the most economically significant PGE-enriched LMI cumulates, including chromitites, exhibit positive slopes on chondrite-normalized plots (e.g., Fig. 2; see also Day et al. 2008; O’Driscoll et al. 2009a). This is primarily a function of the relative geochemically incompatible behavior of Pt and Pd compared to the IPGE so that magma chamber processes that form LMI cumulates concentrate the PPGE relative to their mantle source. However, it is worth mentioning that some chromitite deposits with relatively low total abundances of the PGE (e.g., the socalled sulfide-poor chromitites of the Bushveld Complex; Scoon and Teigler 1994) also have IPGE > PPGE. This highlights the danger of over-generalization in the interpretation of these deposits, i.e., layered intrusion chromitites may also exhibit PGE patterns with a negative slope on chondrite-normalized plots. Given the importance of the presence or absence of chromitite as a host lithology for PGE mineralization (Supplementary Table 1), it is useful to use this distinction as a criterion for sub-division of the following discussion of PGE reefs in LMI.

Chromitite-hosted layered intrusion PGM The UG2 chromitite, Bushveld Complex. The 2054.4 ± 1.3 Ga Bushveld Complex is the largest known LMI on Earth, with an aerial extent of ~70,000 km2 (Scoates and Friedman 2008). The stratigraphy of the Rustenburg Layered Suite of the Bushveld Complex is divided into the Marginal, Lower, Critical, Main and Upper Zones. The Upper Group 2 (UG2) chromitite occurs in the Upper Critical Zone and is the single largest PGE resource on Earth, with Pt grading at ∼ 2.7 ppm and containing a combined 6,636 × 106 tons of measured, indicated and inferred ore (from D. Causey, quoted in Zientek 2012). The UG2 chromitite is broadly stratiform and is approximately 1 m thick, though it commonly comprises more than one seam. There is apparent intrusion-wide lateral and vertical heterogeneity in the absolute PGE abundances of the UG2 chromitite that is reflected in the variability of the PGM present (Schouwstra et al. 2000; Voordouw et al. 2010; Junge et al. 2014). However, where it has

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Figure 2. Chondrite-normalized PGE abundances of selected layered mafic–ultramafic intrusion PGE-ore bodies, divided into chromitite-hosted (a) and non-chromitite-hosted (b) reefs. The data plotted have been sourced as follows: UG2 chromitite (Bushveld Complex; Gain 1985, Voordouw et al. 2010, Junge et al. 2014); Merensky Reef (Bushveld Complex; Prichard et al. 2004); J-M Reef (Stillwater Complex; Godel and Barnes 2008a); Muskox chromitite reef (Barnes and Francis 1995); The Great Dyke Main Sulfide Zone (Oberthür et al. 2003; including samples from the position of the Pt peak, the Pd peak and the lowermost base-metal sulfide accumulation peak); Rum Layered Suite (the Unit 11/12 and Unit 7/8 samples are labelled; O’Driscoll et al. 2009a). Chondrite normalization values are from Naldrett and Duke (1980).

been studied closely, it appears that Pt sulfide, Pt–Pd sulfide, laurite (RuS2), ferroplatinum (Pt,Fe), and Pt–Rh–Cu are consistently important components of the UG2 PGM population (Kinloch 1982; McLaren and De Villiers 1982; Penberthy and Merkle 1999; Schouwstra et al. 2000; Cawthorn et al. 2002; Voordouw et al. 2010; Junge et al. 2014). For example, McLaren and De Villiers (1982) documented ~6,000 PGM grains from 10 drill cores of both

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the eastern and western limbs of the Bushveld. The most abundant PGM they documented were laurite, cooperite (PtS), Pt–Ir–Rh–Cu sulfide, braggite [(Pt,Pd,Ni)S], Pt–Pb–Cu sulfide, vysotskite [(Pd,Ni,Pt)S], and Pt–Fe alloy (see their Table 5 for relative distributions). In a recent study of drill core from the Karee Mine in the western limb of the Bushveld Complex, Junge et al. (2014) documented ∼ 355 discrete PGM with grain sizes in the range < 5–20 µm. The dominant phases are Pt–Fe alloys (30%), laurite (29%), cooperite–braggite (26%), and rare zvyagintsevite (Pd3Pb) and potarite (PdHg). Petrographic examination of the PGM reveals that they typically occur at the margins of sulfide grains, which are themselves located at chromite triple junctions and grain boundaries (Junge et al. 2014). Only laurite is a commonly observed inclusion within chromite crystals, although it is more commonly found at the edges of chromite and sulfide grains. Junge et al. (2014) argued against early precipitation of laurite, suggesting that the textural relationships observed indicated coprecipitation with sulfide and chromite. A significant amount of the UG2 budget of Pd and Rh are hosted in pentlandite (maxima of 2.2 wt.% and 3 wt.%, respectively), to a much greater extent than the other PGE. Junge et al. (2014) also note the presence of rare malanite [Cu(Pt,Ir,Co)2S4], irarsite [(Ir,Ru,Rh,Pt)AsS], platarsite [(Pt,Rh,Ru)AsS], and sperrylite (PtAs2). Together with laurite, Junge et al. (2014) suggest that Pt–Fe alloys and cooperite– braggite co-precipitated with chromite and sulfide, and were later modified by annealing at the postcumulus stage. Voordouw et al. (2010) carried out a detailed SEM–EDS study of the UG2 Reef from the Two Rivers Platinum Mine on the eastern limb of the Bushveld Complex. They documented > 7,000 PGM which were divided among eight classes: cooperite, Pt–Pd sulfide (braggite, vysotskite), Pt–Rh–Cu sulfide, laurite, Pt–Fe alloy, PGE tellurides, PGE sulfarsenide phases, and PGE alloys. Voordouw et al. (2010) emphasized the vertical changes in PGM composition and distribution across their studied section (Fig. 3). The base of the main chromitite seam (1–2.5 m thick, at Two Rivers) contains > 10 wt.% of each of Pt sulfide, laurite, Pt–Fe alloy, Pt–Pd sulfide, and Pt–Rh–Cu sulfide. The middle and top portions of this seam have > 10 wt.% of the first four of these phases, but not Pt–Rh–Cu sulfide. The leader seam, a ∼ 16 cm chromitite in the hanging wall of the UG2, contains a different PGM assemblage (Fig. 3), that is enriched in PGE–sulfarsenides and PGE alloys, relative to the main seam. The leader seam is also noteworthy for its high pentlandite and chalcopyrite (CuFeS2) contents, whereas the main seam base–metal sulfide population is characterized by pyrrhotite (Fe1–xS) and pyrite (FeS2). Voordouw et al. (2010) conclude that the UG2 main seam represents a predominantly primary magmatic PGM assemblage, whereas the leader seam contains PGM that have been modified by secondary (metasomatic) processes. Peyerl (1982) also showed that the primary PGM mineralogy of the UG2 chromitite (i.e., Pt–Pd sulfide) may have been extensively modified (to a Pt–Fe and Pt–Pd–As–Sb-dominated mineralogy) as a result of volatile fluxing associated with emplacement of the Driekop dunite pipe in the eastern limb of the Bushveld. Given that the leader seam hosts a significant component of the mineable UG2 ore at certain localities, the observation that metasomatic processes may upgrade ore potential, by modification of the primary magmatic PGM assemblage and redistribution of the PGE, is an important one. The Merensky Reef, Bushveld Complex. Like the UG2 horizon, the stratiform Merensky Reef occurs in the Upper Critical Zone of the Bushveld Complex, typically between 20–400 m above the UG2 chromitite. The Merensky Reef comprises 3373 × 106 tons of ore (measured, indicated and inferred), grading ∼ 2.9 ppm Pt (from D. Causey, quoted in Zientek 2012). Strictly speaking, the bulk of the PGE mineralization does not occur within chromitite; the Merensky horizon comprises a coarse-grained (pegmatoidal) base-metal sulfide-bearing melanorite, typically on the order of 10–30 cm thick, sandwiched by an upper and a lower (1–3 cm) chromitite seam. However, those base-metal sulfides that occur in the chromitite seams contain approximately twice as much PGE as in the pegmatite (Godel et al. 2007). There has been a plethora of studies on the PGE-bearing phases of the Merensky Reef (e.g.,

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Figure 3. Stratigraphic section of the UG2 chromitite at the Two Rivers Platinum Mine, with corresponding stacked histograms illustrating the variation in whole-rock PGE abundances and PGM assemblage variation across the main seam and leader seams. This image is reproduced from Voordouw et al. 2010, having been previously published as their Figure 3. They analyzed 23 samples for each peak, except for the top of the main seam, which represents data collected from 10 samples. The image is reproduced with the permission of the author and Mineralium Deposita.

Brynard et al. 1976, Vermaak and Hendricks 1976, Kinloch 1982, Mostert et al. 1982, Peyerl 1982, Kinloch and Peyerl 1990; Prichard et al. 2004; Godel et al. 2007; Rose et al. 2011; Wirth et al. 2013) and the reader is referred to these studies for more detail than is provided here. A summary Table adapted and updated from that provided as Table II in Schouwstra et al. (2000) is provided as Table 2 here. Broadly, the dominant PGM are Pt–Fe alloys, Pt–Pd sulfides, laurite and Pt–Pd tellurides. However, like the UG2 chromitite, significant heterogeneity in the relative proportions of the PGM is observed along the lateral extent of the Merensky Reef. Some localities (i.e., at Union Mine on the western limb of the Bushveld) are dominated by (∼ 80 vol.%) Pt–Fe alloys, whereas the Impala Mine to the south (also on the western limb of the Bushveld) is dominated (∼ 56 vol.%) by Pt–Pd sulfides (Table 2). Mostert et al. (1982) investigated the PGM mineralogy of the Merensky Reef at the Impala platinum mines and identified 17 PGM phases in a population of 800 grains. They reported cooperite (44 vol.%), laurite (21 vol.%), moncheite ([(Pt,Pd)(Te,Bi)2]; 17 vol.%), and braggite

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Table 2. Volume percentage distribution of PGM in the Merensky Reef (Bushveld Complex) at a regional scale, organized by mine locality (arranged north to south for each limb)*. The bulk of the data is taken from Table 2 of Schouwstra et al. (2000), but updated with data from other sources (see table footnotes for details). Union

Boschkoppies (Bafokeng)

Impala Platinum**

Rustenburg

Marikana***

—

6.1

Western Platinum

Western Bushveld limb Amandelbult

6

PGM Groups

—

—

51

—

50

9.6

75% of the upper seam PGM (by area) comprise Pt–Fe alloys and Pt–Pd tellurides (Rose et al. 2011). The principal difference is thus in the respective abundances (by area) of the Pt–Pd sulfides, which comprise > 48% of the PGM in the lower seam and ∼ 16% of the upper seam PGM, and the Pt–Fe alloys; the latter comprising only ∼ 13% of the lower seam but > 45% of the upper seam. Cawthorn et al. (2002) make the point that there appears to be a greater concentration of PGE alloys in the north of the Bushveld intrusion, in comparison to the southern and eastern areas where PGE-sulfides and tellurides are more significant, an observation borne out by Table 2. Kinloch (1982) correlated the PGE character of the UG2 chromitite and the Merensky Reef at the regional scale, and showed that where the Merensky Reef shows a predominance of Pt–Pd sulfides in one section of the Bushveld, then the PGM population of the UG2 chromitite is also likely to be dominated by Pt–Pd sulfides. A similar observation holds for the distribution of Pt– Fe alloys. Kinloch (1982) interpreted this observation in terms of metasomatism of a primary magmatic PGM assemblage. He noted that close to intrusive pipe-like features (interpreted to be feeder conduits) and also to reef disturbances (i.e., the Merensky Reef potholes), Pt–Fe alloys tend to dominate the PGM assemblage. Kinloch (1982) suggested that the primary Pt–Pd sulfides had been converted to Pt–Fe alloys in these zones of enhanced volatile activity. In general, the Merensky Reef PGM may be associated with sulfides or with primary and/ or secondary silicates and occur both as inclusions within crystals and along grain boundaries (Cawthorn et al. 2002). For example, Prichard et al. (2004) note that there is no particular tendency for laurite to occur as inclusions within chromite; instead all 16 grains they observed were found in close proximity to base-metal sulfides. In their high-resolution X-ray computed tomography (CT) study of the Merensky Reef chromitites, Godel et al. (2010) reported a preponderance of PGM at chromite–silicate–sulfide triple junctions. The Merensky Reef base-metal sulfides have total average concentrations of the PGE of ∼ 500 ppm, but given the fact that they occur in modal abundances greater than those of the PGM, their contribution to the total PGE budget of the ore deposit is deemed to be significant (Ballhaus and Ryan 1995). This generalization is supported by the observations of Godel et al. (2007), who showed that between 65% and 85% of the PGE budget of the reef is accounted for by PGM. Of the sulfides, pentlandite is the most important host for the precious metals, with total PGE concentrations of up to 600 ppm. Several workers have carried out Re–Os [187Re → 187Os + b–; t½ = 41.6 × 109 yr] and Pt–Os [190Pt → 186Os + a; t½ = 469 × 109 yr] isotopic studies on base-metal sulfides and PGM from the Merensky Reef (Hart and Kinloch 1989; Schoenberg et al. 1999; Coggon et al. 2011a). In an extensive study, Hart and Kinloch (1989) obtained consistent and relatively radiogenic 187 Os/188Os for 36 laurite grains (in the range 0.17–0.18). Additionally, they analyzed two erlichmanite (OsS2) grains that yielded 187Os/188Os compositions of 0.11, a value consistent with that of the chondritic mantle at 2.06 Ga. Coggon et al. (2011a) analyzed several different PGM from the Merensky Reef (including laurite, cooperite, sperrylite, and Pt–Fe alloy) by LAMC-ICPMS and found that their data defined a Pt–Os isochron with an age of 1995 ± 50 Ma, 186 Os/188Osinitial = 0.11982 ± 0.00001 (2s, MSWD = 1.16). The latter authors also used a single PGM grain of cooperite to calculate a 190Pt–186Os model age of 2024 ± 101 Ma. The model age and the isochron age are 30–59 Ma younger than the U–Pb zircon age for the Merensky Reef, explained by Coggon et al. (2011a) as reflecting late stage metasomatism of the ore body.

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Petrogenesis of the Merensky Reef and the UG2 chromitite: ‘Uppers’ or ‘Downers’? The classic models for PGE mineralization in LMI have addressed the issue of how magmas with relatively low initial concentrations of the PGE (e.g., < 5 ppb) could form crystalline (cumulate) products possessing enrichments of up to 103 ppm of these precious metals. In this respect, the presence of base-metal sulfides is significant. Under the conditions typical in many basalt magma chambers, the PGE have chalcophile tendencies, meaning that they will be preferentially partitioned into a sulfide phase, if present. The process by which small volumes of sulfide, envisaged to occur in magma chambers as immiscible droplets of sulfide melt, can equilibrate with many times their own volume of silicate magma so that the PGE may be ‘scavenged’ remains the outstanding question (see discussion of R-factors in Barnes and Ripley 2016, this volume). Two opposing schools of thought have dominated this debate over the past ∼ 30 years (cf. Mungall and Naldrett 2008). Importantly, both involve transport of the PGE to the site of mineralization, from above or below. The ‘downers’ models invokes separation (or unmixing) of sulfide droplets from the magma, following magma mixing. These sulfides settle out through the magma column, scavenging PGE as they do so, before accumulation at the site of mineralization and exsolution of the PGM from the sulfide, typically envisaged to be at or close to the magma chamber floor. The Merensky Reef was originally interpreted in the context of the classic downers model, having developed as a result of a large scale mixing event between new and resident magma, followed by downward settling of chromite crystals and sulfide droplets through a column of magma several km thick (Campbell et al. 1983; Naldrett 1989). A revised version of the ‘downers’ model invokes fractional crystallization and gravity settling of chromite and sulfide droplets from a basally emplaced layer of magma only ~16 m thick (Naldrett et al. 2011). Similarly, the UG2 chromitite and attendant mineralization was traditionally considered to be the product of mixing of resident and replenishing magma, following the classic models for stratiform chromitite development proposed by Irvine (1977a,b). Building on the work of Eales (2000), Mondal and Mathez (2007) subsequently attributed the formation of the UG2 chromitite to emplacement of new magma with much of the chromitite crystal cargo already entrained. The ‘uppers’ model calls for the upward percolation through the (footwall) cumulate pile of a chloride-rich aqueous fluid capable of stripping out (dissolving) PGE-enriched sulfides, before re-precipitating these as a PGE-reef at a stratigraphically higher level (cf. Boudreau and McCallum 1992a, for the Stillwater Complex J-M Reef). The ‘downers’ model is therefore a primary magmatic phenomenon, whereas the ‘uppers’ model occurs later in the solidification history of the magma chamber, when high-temperature metasomatism is more likely to be an important process. Cawthorn et al. (2002) drew attention to the Merensky Reef laurite 187 Os/188Os data of Hart and Kinloch (1989), suggesting that there is no obvious overlap with the underlying anorthosite 187Os/188Os isotopic compositions and therefore there is a problem with the uppers model for the Merensky Reef. The latter authors also summarized some of the variations present in these general models, with particular relevance to the formation and concentration of the PGM in LMI. In particular, they drew attention to the ideas of Tredoux et al. (1995), which describe clustering of metal ions in magma to form PGM, followed by their entrainment in sulfide. The initial formation of the PGM as Fe–PGE alloys or discrete PGM was also suggested by Merkle (1992), Brenan and Andrews (2001) and Gornostayev et al. (2001) with later capture by either chromite (Hiemstra 1985) or sulfide (Ballhaus and Sylvester 2000). Cawthorn et al. (2002) pointed out that early formed PGM assemblages should be dominated by the IPGE, relative to the PPGE; something that does not generally hold for either the Merensky Reef or the UG2 chromitite (as described above). The latter authors referred to the study of Barnes and Maier (2002), which proposed processes involving the early development of discrete PGM, followed by their later modification involving basemetal sulfide, were responsible for the relative abundances of the PGE observed. Indeed, low temperature modification of PGM assemblages (cf. Ballhaus and Ryan 1995) has been

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suggested (to temperatures as low as 80–90 °C), illustrating that for a given section of either the Merensky or UG2 horizons, the specific PGM assemblage should be considered as a combined product of the time-integrated magmatic and high-to-low temperature metasomatic processes. Recent experimental measurements carried out by Fonseca et al. (2009) and Mungall and Brenan (2014) suggest very high sulfide/silicate partition coefficients for the PGE in the Merensky Reef and the UG2 chromitite. Specifically, Mungall and Brenan (2014) propose sul DPGE of ∼ 106 for the Merensky Reef and the UG2 chromitite (Fonseca et al. 2009 suggest values of ∼ 107 for formation of the Merensky Reef). The implication of these results is that whatever the causal mechanism for sulfide separation (and accumulation), there is no requirement for upgrading or modification of the ore deposits to explain the very high metal tenors. Instead, they are in the range of what might be expected for the high R-factors associated with separation of sulfide melt in both the Merensky and UG2 reefs. The Stillwater Ultramafic Series chromitites. The Stillwater Complex (Montana, USA) is a 2705 ± 4 Ma LMI emplaced into mid-late Archean metasediments in association with a major crust-formation event (McCallum 1996). The exposed part of the complex covers 180 km2. It comprises three series with the chromitites occurring in the lowermost Ultramafic Series, at the base of the cyclic harzburgite–orthopyroxenite units (Todd et al. 1982). The chromitite seams are named A (stratigraphically lowest) to K (stratigraphically highest). The association of the chromitite seams with the Ultramafic Series has been considered to reflect chromite crystallization from mixing of the resident magma with repetitive influxes of olivine-saturated, high-Mg melts (interpreted by some authors as boninites; Boudreau et al. 1997) into an open-system magma chamber (McCallum 1996; Horan et al. 2001). Talkington and Lipin (1986) documented the types and distribution of PGM in the (A, C, E, G, J and K) chromitites. Whole-rock abundances of the PGE range from several ppb to ∼ 16 ppm, with Pd> Pt> Rh> Ru> Ir. The latter authors did not analyze Os, but Horan et al. (2001) subsequently reported concentrations of up to ∼ 78 ppb in the Stillwater Ultramafic Series chromitites. The whole-rock abundances are not well accounted for by the PGM observed, which are predominantly IPGE-rich phases. Talkington and Lipin (1986) observed PGM (≤20 µm in size) both as inclusions in chromite and as interstitial phases in massive and disseminated chromitite lithologies. In the former case, inclusions contained within unfractured chromite crystals are predominantly laurite. The interstitial phases are predominantly sperrylite and isoferroplatinum. Talkington and Lipin (1986) favored a magmatic origin for the Ru–Ir–Os-rich phases, implying that PGM such as laurite precipitated relatively early, whereas they considered that Rh, Pt, and Pd may have precipitated at a later stage in the development of the chromitites. The primary rationale for this interpretation was the occurrence of laurite as inclusions within chromite crystals, and the interstitial textural position of the PPGE-rich phases. They noted the close spatial proximity of PPGE-rich PGM and interstitial sulfides, arsenides, antimonides, and mercurides, and proposed partitioning of the PPGE into an immiscible sulfide fraction at the magmatic stage, together with late-stage hydrothermal fluid processing of the primary PGM assemblage, to explain their observations. The Rum Layered Suite chromitites. The Rum Layered Suite (NW Scotland) is a ∼ 60 Ma open-system LMI, emplaced during the onset of opening of the Northeast Atlantic. The eastern portion of the intrusion (Eastern Layered Intrusion) contains chromitite seams at the bases of cyclic peridotite–troctolite units; each unit is considered to be the product of emplacement of a fresh batch of basaltic or picritic magma into the chamber. The chromitites are laterally extensive for 100’s of meters, rarely exceeding ∼ 2 mm in thickness and their formation has been attributed to the assimilation of feldspathic (+ clinopyroxene) cumulate by replenishing picritic magmas (O’Driscoll et al. 2010). They are characterized by significant whole-rock enrichments in the PGE (ppm levels; Fig. 2a) compared to the cumulate above and below them (O’Driscoll et al. 2009a). Power et al. (2000) carried out a detailed documentation of the Rum PGM and reported a considerable abundance and diversity of mineral species. More than 70% (by number) of all

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of the PGM observed (∼ 850 grains over 5 unit boundary chromitite seams) are associated with base-metal sulfides, typically at the grain boundaries between sulfide and other mineral phases rather than as sulfide-hosted inclusions. The remainder (∼ 28%) are either hosted in Fe-oxide and Fe-hydroxide phases or associated with silicates. One of the most remarkable features of PGE mineralization in the Rum chromitite seams is that the PGM populations are highly variable from one unit to the next, presumably reflecting the open system character of the intrusion, and perhaps the compositional heterogeneity of the replenishing magmas (Power et al. 2000). For example, at the base of the Eastern Layered Intrusion, Unit 1 contains electrum (Au,Ag) grains, but there are no reported PGM. The Unit 5/6 and 6/7 boundaries are dominated by Pt–Pd tellurides and bismuthides, whereas the Unit 11/12 chromitite seam is dominated by arsenide phases, especially sperrylite. Approximately 70% of the PGM from a chromitite seam in Unit 14 are tellurides or arsenides. The Unit 7/8 chromitite, with 374 documented PGM, is the most PGE-enriched (Fig. 4). In this case, Pt–Fe alloys and PGE sulfides (braggite and cooperite), laurite, and Pt–Pd alloys dominate the assemblage. This assemblage would appear to record a relatively high-temperature development. In this light, it is worth noting that O’Driscoll et al. (2009a) also recorded the highest whole-rock PGE concentrations from the Unit 7/8 horizon and suggested that the magma replenishing the Rum chamber at this level was particularly primitive (magnesian) in character (i.e., picritic). Power et al. (2000) argued that the tight spatial control exerted by the ∼ 2 mm chromitite seams on the PGM abundances implied a magmatic origin, although they did not rule out localized in situ alteration/re-distribution of some of the phases at the postcumulus stage. On the basis of mineral compositional and textural evidence, O’Driscoll et al. (2009b, 2010) argued against the magma mixing hypothesis proposed by Power et al. (2000) and proposed that the Rum chromitite seams, sulfides, and PGM developed in situ. Latypov et al. (2013) expanded on the latter idea, with an in situ crystallization model that invoked nucleation of sulfides onto chromite crystal surfaces before scavenging the PGE from the convecting magma in the chamber. Power et al. (2003) also documented an extensive array of PGM associated with another ultramafic intrusion on the island of Rum. The PGM predominantly occur within and at the edges of disseminated base-metal sulfide grains and include paolovite (Pd2Sn), and Pd bismuthotellurides. The ultramafic intrusion occurs as a satellite plug to the Rum Layered Suite and further highlights the potential for the Rum parental magmas to produce PGE mineralization. Other layered intrusion chromitite-hosted PGM. It is generally accepted that magma chamber conditions and the processes that lead to the crystallization of chromitite seams in layered intrusions are also conducive to enrichment of the PGE, thus leading to high concentrations of the PGM in chromitite. The Muskox intrusion is an important LMI in the literature in this regard, as some of the classic ideas of chromitite petrogenesis were developed for the seams that occur there (Irvine 1977a,b). The Muskox intrusion is located on the northwestern edge of the Canadian Shield, and its chromitite seams are positioned in the middle of cyclic unit 21 and at the base of cyclic unit 22. The formation of these chromitites has been attributed to magma mixing in an open system magma chamber by Irvine (1977a; 1977b). No detailed documentation of the PGM has been reported in these rocks, but Barnes and Francis (1995) reported whole-rock base and precious metal concentrations throughout the Muskox stratigraphy. They showed that the enrichments of the PGE were not significant in the chromitite reef at the boundary between cyclic units 21 and 22 (i.e., 300–1000 ppb S PGE), compared to major chromitite-hosted PGE deposits, suggesting that chromitite PGE abundances can be quite variable from one deposit to the next. Notably, however, these PGE abundances are not dissimilar to those in the Rum chromitites. In the case of the Muskox intrusion, Barnes and Francis (1995) attributed the relatively low abundances of the PGE to a low R-factor, compared to that experienced by the chromitites from the Merensky or UG2 reefs (i.e., ∼ 1,000 at the Muskox intrusion compared to > 10,000 for reefs such as the Merensky Reef; see Barnes and Ripley 2016, this volume, for additional detail on the R-factor principle). However, it is worth noting that Day et al. (2008) reported Os concentrations of > 200 ppb in unit 22 chromitite

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Figure 4. Grayscale QEMSCAN® image of the Unit 7–8 boundary, Rum Layered Suite. The ∼2 mm thick chromitite seam running east–west across the image separates overlying Unit 8 feldspathic peridotite from the underlying Unit 7 anorthosite. The different silicate minerals are not distinguished in this image, but the distribution of Cr-spinel (colored black) in the seam and in the overlying peridotite is visible. QEMSCAN® maps samples at electron beam stepping intervals down to 1 µm and interprets the mineralogical composition of a sample at the pixel scale (see O’Driscoll et al. 2014a for further details). In the area shown by this image (adapted from O’Driscoll et al. 2014a), 31 PGM and Au/Ag-rich mineral phases were mapped by QEMSCAN® analysis (the positions of some of these are shown by the white stars). The close spatial association of these minerals with the chromitite seam is noteworthy. The white stars include both PGM and Au/Ag phases. Note that 3 grains are located out of the area of the sample.

(an order of magnitude greater than that reported by Barnes and Francis, 1995), suggesting that there may be a degree of intra-reef variability too. Other detailed PGM studies, such as those carried out by Pirrie et al. (2000) on chromitites in the ∼ 60 Ma Mull and Skye layered intrusions (British Paleogene Igneous Province) support the general point however, that LMIhosted (stratiform) chromitite seams tend to be enriched in PGM. Furthermore, Halkoaho et al. (1990) describe the Sompujärvi PGE Reef at the boundary between the third and fourth megacyclic unit in the Penikat layered intrusion (Finland), noting that where associated with disseminated chromite, the PGE grade is distinctly higher. The link between chromite and PGM was explored by Finnigan et al. (2008), who carried out a series of experiments in which they documented the formation of PGM (especially IGPE-rich PGM) at the chromite-melt interface. They argued that the selective uptake of Cr3+ and Fe3+ from the melt by the growing spinel created a boundary layer across which a redox gradient developed. Finnigan et al. (2008) showed that PGE solubilities in the melt can consequently decrease dramatically (by as much as 20%) in these localized reduced (fO2) melt films, offering an explanation for the close association between the PGM and chromitite in natural environments, but also for why the PPGE are only typically enriched in chromitites that contain abundant base-metal sulfide.

Non-chromitite-hosted PGM in layered intrusions The J-M Reef, Stillwater Complex. The J-M Reef of the Stillwater Complex contains bulk rock PGE at some of the highest grades (∼ 18 ppm) of any deposit on Earth (Zientek et al. 2002; Godel and Barnes 2008a). Zientek (2012) reports ∼ 149 × 106 tons of reserves and mineralized material for the deposit, grading at 3.7 ppm Pt and 12.9 ppm Pd. The mineralization occurs

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within the olivine-bearing cumulate (OB-1) of the Lower Banded Series (Zientek et al. 1985). Its PGM population includes Pt- and Pd-sulfides, tellurides, and Pt–Fe alloys (Heyse 1983; Zientek and Oscarson 1986). Zientek and Oscarson (1986) reported that all of the Pt and ∼ 20% of the Pd are contained in PGM, with pentlandite also being a significant host for Pd. Godel and Barnes (2008b) investigated the PGM phases in four different samples of J-M Reef cumulate (troctolite, anorthosite, leuconorite, and olivine melagabbronorite) and reported ∼ 850 PGE-rich grains, which they divided into eight groups. There is variability between samples, reflecting the compositional heterogeneity of the J-M Reef in general (Figs. 5, 6). Godel and Barnes (2008b) reported that Pd–Pt sulfides and Pt–Fe alloys dominate the numbers of grains observed (∼ 44% and ∼ 31%, respectively) with Pd–Pt tellurides also representing a significant component (∼ 18% of total grains). The Pd–Pt sulfides are predominantly braggite–cooperite and vysotskite, representing 59% of the total area measured across the four samples. Most of the PGM occur as vermicular-type structures within but close to the margins of base-metal sulfides, especially chalcopyrite. The Pd–Pt sulfides are also observed enclosed within secondary silicates (i.e., chlorite, amphibole) and oxide grains (magnetite; Godel and Barnes 2008b). In the latter instance Pd–Pt sulfides are recorded from the centers of secondary magnetite-bearing veins that crosscut the base-metal sulfides, related to late Cretaceous–early Paleogene Laramide orogenesis (McCallum 1996; Fig. 5b). The Pt–Fe alloy is predominantly isoferroplatinum, which may contain up to ∼ 5.6 wt.% Pd (Godel and Barnes 2008b). This mineral is mostly associated with base-metal sulfides, either as inclusions within pyrrhotite and pentlandite (∼ 48% by area) or at the grain boundaries between silicate and sulfide (∼ 39% by area). Godel and Barnes (2008b) report that the isoferroplatinum grains are relatively large compared to other PGM, with an average grain size of ∼ 50 µm2. A lesser number of isoferroplatinum grains (∼ 12% by area) are observed to occur in close association with magnetite (Fig. 5a). Tellurides of Pd and Pt account for 12% by area of the total PGM population, but their contribution to the overall PGE budget is estimated at only 0.1% and 5.1%, for Pt and Pd, respectively. Telluropalladinite (Pd9Te4), keithconnite [Pd3–xTe(x = 0.14–0.43)] and kotulskite [Pd(Te,Bi)] are the dominant PGM, typically occurring in close association with (either at the edges of or as inclusions within) secondary silicate phases. Laurite accounts for 1% (by area) of the total PGM observed by Godel and Barnes (2008b), but constitutes up to 50% of the whole-rock Os, Ir, and Ru contents. Most laurite (80% by area) is present as inclusions in base-metal sulfides (e.g., pentlandite and chalcopyrite), the rest occurring at contacts between secondary magnetite and base-metal sulfides, where they are usually smaller grains. Other PGM recorded in the J-M Reef include Pd–Pb and Pd–Cu alloys (zvyagintsevite and skaergaardite [PdCu], respectively), Au–Pd–Ag alloys, and native Pd (∼ 3.5% by area of total PGM), all typically at the contacts between basemetal sulfides and (typically secondary) silicate phases. From the textural relationships and mineralogical associations, Godel and Barnes (2008b) proposed that the PGE, Te, Bi, and basemetals were initially contained in an immiscible sulfide liquid fraction. At least two subsequent alteration events were invoked by these authors as follows: 1) Initial desulfurization of the base-metal sulfides due to an upwardly mobile S-undersaturated silicate melt which formed the Pt–Fe alloy and some of the Pd sulfide, and 2) A low temperature (∼ 250–465 °C) metasomatism event which led to the formation of secondary magnetite and the other Pd-alloys, including the Pd–Cu PGM. As a caveat to the latter, the observation of Barnes and Naldrett (1986) that wholerock PGE contents correlate positively and strongly with S concentrations suggests that any S redistribution was relatively localized. In addition, the remobilization of the J-M Reef PGM assemblage at low temperatures during magnetite formation shows that such modification can occur significantly after initial crystallization. The Picket Pin deposit occurs up stratigraphy from the J-M Reef, in the upper portion of Anorthosite subzone II (AN II), and is traceable along 22 km of strike of the Stillwater intrusion. Although there has not been a detailed mineralogical characterization of the PGM associated with this horizon, Boudreau and McCallum (1992b) report that arsenide and antimonide PGM dominate.

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Figure 5. Textural relationships of PGM illustrated in backscatter electron micrographs of the Stillwater J-M Reef as follows: (a) Irregularly shaped bright Pt–Fe alloy grains (example arrowed) are included in pyrrhotite and crosscut by ∼10 µm thick secondary magnetite veinlets (arrowed). (b) Sieve-textured magnetite rim (on Cr-spinel crystal) containing Pd-telluride inclusion. These images have previously been published in Godel and Barnes (2008b) as their Figures 6d and 8c, and are reproduced here with the permission of the authors and under the Fair Use Provision of Economic Geology.

The Main Sulfide Zone, Great Dyke. The Great Dyke (Zimbabwe) is a 2575 ± 0.7 Ma intrusion emplaced into Archean granites and greenstone belts of the Zimbabwean craton (Oberthür et al. 2002). Its lower Ultramafic Sequence contains chromitites with sub-economic concentrations of the PGE. The Main Sulfide Zone (MSZ) contains disseminated sulfide mineralization hosted predominantly in pyroxenites. The MSZ has measured, indicated and inferred resources of 2136 × 106 tons, at typical grades of 2.7 ppm Pt and 1.8 ppm Pd (from D. Causey, quoted in Zientek 2012). The MSZ is situated several meters below the transition between what are referred to as the lower Ultramafic and upper Mafic Sequences, and has economic concentrations of the PGE (Oberthür 2011) that occur as (Pt,Pd-) bismuthotellurides, PGE-arsenides, -sulfides, and sulfarsenides (Johan et al. 1989a; Oberthür et al. 1997, 1998, 2000). Oberthür et al. (2003) carried out a detailed study of a suite of mineralized bronzitites at Hartley Platinum Mine, where the MSZ is several meters thick and comprises a lower sub-economic PGE-subzone and an upper base-metal sulfide subzone. The MSZ has a finescale geochemical structure, whereby peaks in concentrations of different PGE and in the abundance of base-metal sulfides are stratigraphically offset from one another. As described in Oberthür (2011), the main peak in Pd concentrations occurs in the PGE subzone (plus a minor Pt peak), followed by the major Pt peak several meters above. The lower portion of the base-metal sulfide subzone overlaps with the top of the PGE subzone, and the first

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Figure 6. Pie charts illustrating the proportion of PGM observed in the four samples from the J-M Reef (troctolite, anorthosite, leuconorite, and olivine melagabbronorite with ~1 to ~5 vol.% base-metal sulfides) studied by Godel and Barnes (2008b) by area (a) and number of grains (b). The samples contain ~49 to 419 ppm bulk rock PGE (see also Fig. 2). Summary pie charts of the textural distribution of the different PGM by (c) area and (d) number of grains have been replotted using the data presented in Figures 3 and 4 of Godel and Barnes (2008b).

peak in sulfide concentration occurs < 1 m above the Pt peak (Oberthür 2011). Oberthür et al. (2003) interpreted the structure of the MSZ as reflecting successive episodes of primary sulfide accumulation, PGE scavenging and fractionation. They documented 181 PGM grains throughout the MSZ (e.g., Fig. 7), predominantly Pt-rich bismuthotellurides and Pd-rich bismuthotellurides (55% and 16% of the total number of PGM grains, respectively). The most common PGM in this category are moncheite, maslovite, merenskyite [(Pd,Pt)(Te,Bi)2], and michenerite [(Pd,Pt)BiTe]. Other phases recorded include sperrylite (11% by number) and cooperite/braggite (2% by number). Rarer PGM include laurite, Pt–Fe alloy, and an unnamed phase (PtSnS). Oberthür et al. (2003) also report electron microprobe data that suggest much of the Pd and Rh is hosted in pentlandite (maximum concentrations of ∼ 2,500 ppm and 550 ppm, respectively) in the PGE subzone. In contrast, Pt is concentrated in PGM phases. The PGM occur in discrete zones that match the geochemical zonation of the MSZ. Sperrylite occurs throughout the PGE subzone, cooperite/braggite in the lower part only and (Pt,Pd)-bismuthotellurides are concentrated at the top of the PGE-subzone. Oberthür et al. (2003) interpreted their observations as signifying that the PGE were primarily concentrated in base-metal sulfides at magmatic temperatures, referring to the observations of Ballhaus and Ryan (1995) on Merensky Reef base-metal sulfides. Subsequent expulsion of the PGE (Os, Ir, Ru, and especially Pt) occurred during recrystallization of the sulfides to pentlandite, pyrrhotite, chalcopyrite, and pyrite. Oberthür et al. (2003) suggested that Te and Bi were also

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Figure 7. (a) Reflected light image (oil immersion) of braggite and moncheite grains (labelled in image) in the MSZ ore of the Great Dyke. (b) Backscatter electron micrograph of euhedral laurite grain intergrown with base-metal sulfides (chalcopyrite and pyrrhotite) from the MSZ. A grain of moncheite (labelled) also occurs. (c) Ternary (PtS–PdS– NiS mol.%) diagram illustrating the range of compositions of cooperite and braggite from pristine and altered (oxidized) MSZ base-metal sulfide ores. Panels (a) and (b) are reproduced from Figures 5a and 5e of Oberthür et al. (2003) and (c) is a reproduction of Figure 7 of that paper. All panels are reproduced with the permission of the author and of Mineralium Deposita.

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expelled from the sulfide lattice at this stage, whereupon they combined with the PGE to give the observed parageneses. Low temperature (< 500 °C) interstitial ‘magmatic-hydrothermal’ fluids were invoked by Oberthür et al. (2003) as a potentially important agent in redistributing the PGE during sulfide recrystallization, though the latter authors indicate that the precise role of metasomatism in PGM formation needs further investigation. The Platreef, Bushveld Complex. The Bushveld Complex Platreef crops out in the northern limb of the intrusion and has measured, indicated and inferred resources of 3418 × 106 tons at grades of 2.8 ppm Pt and 3.4 ppm Pd (Mudd 2012). The stratabound (but not stratiform) Platreef is located at the base of the Rustenburg Layered Suite, in direct contact with underlying Archean metasedimentary and meta-igneous country rocks and reaches thicknesses of ∼ 40 m. Building on earlier syntheses of the PGM populations of this economically significant deposit (e.g., Kinloch 1982; Viljoen and Schürmann 1998; Hutchinson et al. 2004), Holwell et al. (2006) carried out a detailed study of the PGM at the Sandsloot Mine in the Platreef. Base-metal sulfides (including pyrrhotite, pentlandite, chalcopyrite) are common interstitial components of the reef and are characterized as being anhedral and exhibiting a tendency to be intergrown with plagioclase and secondary phases such as actinolite, epidote and micas. Holwell et al. (2006) note that PGE grades in the reef itself are variable and are quite erratic in the footwall. They identified > 1000 PGM grains in 58 polished thin sections and blocks of reef, hanging wall, and footwall samples. Nine broad groupings of the PGM were established by these authors as follows: (1) Pt/Pd tellurides; (2) Pt/Pd bismuthides; (3) Pt/Pd arsenides; (4) Pt/Pd antimonides; (5) Pt/Pd germanides; (6) PGE sulfides; (7) PGE sulfarsenides; (8) PGE alloys with Fe, Cu, Sn, Pb, and Tl; and (9) Au, and Ag bearing phases. Holwell et al. (2006) highlighted the paucity of S-dominant PGM in the Platreef at Sandsloot, in direct contrast to both the UG2 chromitite and the Merensky Reef. Overall, footwall and hanging wall PGM populations in the Platreef are dominated by Pt/Pd tellurides, alloys with lesser PGE-bearing arsenides (especially sperrylite) and antimonides. In particular, the Platreef is dominated by Pt/Pd tellurides including moncheite and kotulskite with sperrylite being common throughout the deposit too. For the most part, the PGM in the reef pyroxenites and pegmatites are surrounded by silicates or situated at silicate–sulfide grain boundaries. Holwell et al. (2006) showed that many are texturally associated with the replacement products of sulfides, such as secondary amphiboles (tremolite and actinolite). Some of the pyroxenite at the Platreef has undergone a replacement that introduced Fe-rich olivine to the rock, and these portions of the reef are characterized by a distinct PGM assemblage. Specifically, this includes Pd-tellurides and alloys and Pt–Fe alloys, as well as rare sperrylite. Another feature of note is that 55% (of the total area of PGM studied) of the Platreef hanging wall material comprises an unusual Pd-germanide with a composition close to Pd2Ge, attributed by these authors to processing of the reef by newly replenishing Main Zone magmas. The complex PGM mineralogy of the Platreef package is likely (at least in part) a function of the interaction of parental magmas with the calc-silicate floor rocks (Schouwstra et al. 2000). Holwell et al. (2006) suggested that formation of an initial telluride-dominated PGM assemblage developed in the pore spaces of the primary cumulate texture, possibly controlled by the interaction of the intercumulus melt with primary magmatic base-metal sulfides. They argued for localized modification of the PGE-assemblage in the reef when a late-stage Fe–Pb-rich fluid, possibly derived from serpentinization of the footwall lithologies, was introduced. Holwell et al. (2006) suggested that the telluride assemblage was then converted into one dominated by Pt–Fe and Pd–Pb alloys. This interpretation was based partly on the observation that zvyagintsevite (Pd3Pb) is unusually common in these rocks and that kotulskite has enhanced Pb contents (up to 12 wt.%), reflecting replacement of Te. Holwell et al. (2006) highlighted similarities between this process and the metasomatism considered to have modified the PGM population around the Bushveld dunite pipes and the Merensky Reef potholes (Kinloch 1982; Kinloch and Peyerl 1990). Episodes of hydrothermal fluid alteration and serpentinization are considered

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to have produced the arsenide and antimonide-dominated assemblage in the footwall, based on the extremely different character of the PGM mineralization compared to that in the reef itself. For example, Holwell et al. (2006) highlighted the difference in whole-rock Pt/Pd ratios between the reef and footwall (0.79–1.94 vs. 0.54–0.98, respectively), which they argued reflected formation of the footwall from enhanced metasomatism as Pd is more mobile than Pt. The Platinova Reef, Skaergaard Intrusion. The Skaergaard Intrusion lies at the unconformity between Precambrian gneisses and an overlying sequence of Eocene plateau lavas in Kangerdlugssuaq, southeastern Greenland. It formed via the successive emplacement of a series of magma pulses during opening of the Northeast Atlantic Ocean (∼ 55 Ma) and is the archetypical example of the closed system fractionation of a basaltic magma (Wager and Brown 1968). Crystallization of the magma resulted in the accumulation of three different series: the Layered Series, the Marginal Border Series and the Upper Border Series. The so-called Platinova Reef occurs in the Triple Group, which constitutes the upper ∼ 100 m of the Middle Zone of the Layered Series. The Platinova Reef is hosted in an Fe-Ti oxide-rich tholeiitic gabbro, and the associated mineralization is broadly stratiform. Inferred resources of 1,520 × 106 tons of ore, grading at 0.04 ppm Pt, 0.61 ppm Pd, and 0.21 ppm Au, have been reported (Mudd 2012). The base-metal sulfides of the Platinova Reef are dominated by Cu-rich minerals, contrasting with Fe–Ni–Cu sulfides that are typical of some of the PGE reefs already discussed above. The sulfides are predominantly bornite (Cu5FeS4), chalcocite (Cu2S), and digenite (Cu9S5) and comprise ∼ 0.05 mod.% of the reef. They occur in close textural association with Ti-magnetite/clinopyroxene grain boundaries, occasionally being intergrown with H2O-bearing silicates (chlorite-group minerals, hornblende, actinolite, and a member of the annite–phlogopite series). The PGM assemblage has been documented in detail by Nielsen et al. (2003a–e) and is dominated (∼ 90% of the total PGM observed) by skaergaardite (PdCu; Rudashevsky et al. 2004). Other PGM documented as both discrete grains and inclusions include keithconnite, vasilite [(Pd,Cu)16(S,Te)7], and zvyagintsevite. The PGM are characteristically partially or completely enclosed by the Cu–Fe sulfides (e.g., Rudashevsky et al. 2004; Godel 2013), an observation that led Andersen (2006) to suggest that the PGM are genetically related to formation of the sulfides. Rudashevsky et al. (2004) suggested that skaergaardite crystallized from a disordered Pd–Cu-rich alloy precursor that had exsolved from a Cu–Fe sulfide melt under conditions of relatively high fS2 [log fS2 > 7 units] at ∼ 600 °C. Unnamed Cu–Pd–Au–Pt–Fe alloys were also documented by Rudashevsky et al. (2004). In 2008, McDonald et al. also reported the occurrence of another new Pd-dominant alloy (or intermetallic), nielsenite (PdCu3), in the Platinova Reef (see also Rudashevsky et al. 2009). Karup-Møller et al. (2008) carried out an experimental study of the phase system Cu–Fe– Pd–S, at temperatures of 1000 °C, 900 °C, and 725 °C. They suggested that the skaergaardite precursor first formed as a Pd–Cu alloy at relatively high temperatures (≥1000 °C) in association with bornite (both crystallized from a metal-rich sulfide melt). In broad agreement with Rudashevsky et al. (2004), Karup-Møller et al. (2008) considered that skaergaardite (low skaergaardite; b-CuPd) formed following cooling of Pd–Cu alloy below 600 °C. These observations re-emphasize an important point already made above for the Merensky Reef PGM; that high-temperature PGM assemblages are capable of continuously reacting and reequilibrating to significantly lower temperatures than the magmatic conditions under which they first formed. Indeed, Karup-Møller et al. (2008) suggest a lower temperature limit of 300 °C for PGM formation in the Platinova Reef. The latter authors extrapolated their observations to nielsenite formation and speculated on its high-temperature crystallization at ∼ 1100 °C, with subsequent re-equilibration to temperatures ≤ 500 °C. Other non-chromitite-hosted PGM. There are numerous layered intrusions, in addition to those named above, with non-chromitite-related PGE mineralization. Examples such as the Duluth Complex (USA) and the Sudbury Igneous Complex (Canada) are more appropriately

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discussed below in the section on PGE–Ni–Cu sulfide-rich deposits, in terms of the style of PGM mineralization. Other published PGM studies include work on LMI such as the River Valley Intrusion (Canada; Holwell et al. 2014), the Freetown Layered Complex (Sierra Leone, Africa; Bowles et al. 2013), the Konttijaervi (associated with the basal contact of the Suhanko intrusion) and Kevitsa (also a Ni–Cu–PGE deposit) intrusions (northern and central Finland, respectively; Vuorelainen et al. 1982; Gervilla and Kojonen 2002), the Bird River Sill (Canada; Talkington et al. 1983), the Imandra layered intrusion (northwestern Russia; Barkov and Fleet 2004), the Ivrea-Verbano Basic Complex (lower crustal LMI, Italy; Garuti and Rinaldi 1986; Ferrario and Garuti 1990), the Munni Munni layered intrusion (western Australia; Mernagh and Hoatson 1995) and the Sompujärvi PGE Reef in the Penikat intrusion, which is not associated with chromitite, sensu stricto (Finland; Barkov et al. 2005). Supplementary Table 1 contains additional details of these intrusions. The Freetown Layered Complex and the Munni Munni layered intrusion are discussed further below. The Freetown Layered Complex is a ~193 Ma, open-system intrusion associated with opening of the equatorial portion of the Atlantic Ocean. It comprises four major cyclic units, associated with magma replenishment events; within one of these units (Zone 3), four PGEbearing horizons (B, C, D, and M; in order of descending stratigraphy) have been reported. Bowles et al. (2013) studied 21 thin sections of Horizon B and one of each of the others and documented a total of 169 PGM grains (Fig. 8). They noted that Horizons B and D contain Pt–Fe alloys and cooperite. The Pt–Fe alloys range in composition from isoferroplatinum to tetraferroplatinum (PtFe) and dominate the number of PGM grains observed in Horizon B (~52%; or 71% by area), with cooperite comprising ~21% of the total number of grains (and ∼ 21% by area too). Rare tulameenite (Pt2FeCu), bowieite [(Rh,Ir,Pt)2S3], Pt–Ir–Rh base-metal sulfides, and laurite are also present in Horizon B. Horizons C and M are dominated by Pd-bearing PGM. Horizon C contains Pd-bearing native Cu alloys (~75% of total PGM area) and nielsenite (~25% by area), with Pd-antimonide–arsenide dominating in Horizon M. Overall, a large proportion of the PGM (∼ 35%) occur within (but close to the margins) of sulfide grains (Fig. 8c). The majority of the PGM (∼ 50%) are associated with silicate minerals, and especially in interstitial positions to olivine, pyroxene, and plagioclase, and as inclusions within late-stage (hydrous) silicates such as amphibole, chlorite, and phlogopite (Figs. 8a,c). Bowles et al. (2013) invoked a model involving early-formed cooperite that underwent alteration (desulfurization) to Pt–Fe alloy with Pt-oxide as the final product. [It should be noted that the existence of Pt-oxide has been called into question by Hattori et al. (2010). Reference to this mineral phase here and throughout this work is solely for the purposes of acknowledging the observations of other studies, and does not bear on the views of the present authors.] The presence of arsenides and antimonides in the stratigraphically lowest layer was considered to be an effect of metasomatism. Platinum-group element mineralization in the Freetown Layered Complex probably has more in common with PGE reefs in other open system intrusions, than it does with the Platinova Reef of the Skaergaard intrusion. The study of Bowles et al. (2013) built on earlier studies (Bowles 1986; Bowles et al. 2000) of alluvial PGM from the Freetown peninsula, in which the origin of Pt–Fe alloys and Os-rich phases of the laurite–erlichmanite series was investigated by measuring their 187Os/188Os compositions, by ion microprobe. The PGM were found to have isotopic compositions extending from 0.1181 (± 0.0010) to very radiogenic values of 0.2820 (± 0.0040). One example of a Pt–Fe grain containing laurite inclusions with different 187Os/188Os compositions was reported, suggesting that at least some of the PGM formed in a surficial environment (Bowles 1986; Bowles et al. 2000; see also Appendix), although this was disputed by Hattori et al. (1991). Mernagh and Hoatson (1995) described the PGM mineralogy of the Munni Munni Layered Intrusion, Western Pilbara Block (western Australia). The Munni Munni Complex is late Archean in age and its stratigraphy resembles that of the Great Dyke, discussed above.

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Figure 8. (a) Back scattered electron micrograph of cooperite (PtS) and magnetite (labelled) inclusions in amphibole from Horizon B in the Freetown Layered Complex. Image adapted from Figure 5c of Bowles et al. (2013). (b) Back scatter electron micrograph of Pt–Fe alloy inclusion in chalcopyrite from Horizon D, adapted from Figure 5n of Bowles et al. (2013). (c-d) Pie charts showing (c) the associations between the PGM and their host phases, and (d) the varying abundance of the types of different PGM. BMS = basemetal sulfides. Data are taken from Figures 4a and 4b of Bowles et al. (2013). Data and images are reproduced and presented with the permission of the author and Canadian Mineralogist.

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It comprises a lower Ultramafic Zone (~ 1850-m thick), and an overlying Gabbroic Zone, ≥ 3600-m thick (Barnes and Hoatson 1994). Mernagh and Hoatson (1995) described the PGM (n = 87) from the porphyritic websterite layer at the Ultramafic zone–Gabbroic zone contact. They found that the PGM are dominated by both Pt- and Pd-rich phases (n = 45, n = 42, respectively). Specific phases include sperrylite (19% of total grains), telluropalladinite (15%), potarite + atheneite [(Pd,Hg)3As] (12%), moncheite (9%), platarsite (8%), and native Pt (5%), Pd (5%). With the exception of Pt–Pd sulfarsenides, all PGM were found to be < 10 μm in size. Approximately 78% of the PGM were found to be associated with silicates, either as inclusions or along silicate–silicate grain boundaries. Approximately 15% of PGM occur within chalcopyrite or pentlandite. Although the PGM are thus predominantly hosted in silicate, they are still often spatially associated with base-metal sulfides. For example, PGM are commonly associated with hydrous silicate assemblages (amphibole, biotite) that appear to be a replacement product of sulfides. In addition, PGM occur in embayments or along cleavage traces of silicate (e.g., pyroxene) grains, close to where sulfides occur. Mernagh and Hoatson (1995) suggested that based on their observations, the PGM likely developed during the postcumulus replacement of primary igneous textures/mineralogy by metasomatic fluids.

PGM IN OPHIOLITES Ophiolites are fragments of the oceanic (s.l.) crust and upper mantle, tectonically emplaced onto continental crust during orogenic events. The lowermost sections of many ophiolites offer unrivalled opportunities to study how the HSE are processed in the oceanic mantle via melt extraction, melt-percolation and metasomatism (see also Becker and Dale 2016, this volume). Ophiolite PGM are typically hosted in peridotites and chromitites, as well as placers associated with rivers draining the ophiolitic massifs (see Appendix for a description of the latter). Most of the ophiolites discussed below are considered to be derived from supra-subduction zone settings.

PGM in ophiolite peridotites Occurrence and frequency of PGM in ophiolitic peridotites. Most mantle peridotites from ophiolites are variably serpentinized harzburgites, dunites and lherzolites. These lithologies usually contain PGE in concentrations of only a few ppb, which mostly reside in base-metal sulfides (Leblanc 1991). However, it has been shown that micrometric PGM also contribute to the PGE budget of ophiolite peridotites. For example, Augé (1988) studied PGM inclusions in accessory chromite from dunites at Tiébaghi (New Caledonia) and reported they comprise 50% laurite, ~13% native Os, ~13% unidentified (Ir,Cu)2S3, ~13% prassoite (Rh17S15), and ~13% Pt–Pd–Fe alloys. Chromite-hosted inclusions in Vourinos (Greece) dunites comprise 60% laurite and 40% Os–Ir–(Ru) alloys (Supplementary Table 2). Augé (1988) noted that the morphology and chemistry of the PGM in dunites from both localities are similar. Ohnestetter (1992) reported very small (< 5 µm) PGM in olivine-rich pyroxenites forming the Monte Maggorie Massif (Corsica). They identified 44 PGM grains included in, and attached to, the rims of larger base-metal sulfides (pentlandite, chalcopyrite and bornite) and awaruite (Ni3Fe). Most PGM identified at Monte Maggorie are alloys, including 32% tetraferroplatinum, 11% native Pt and Pd, 9% native Os, and < 2% Ir and Cu–Pd, accompanied by 14% tellurides including merenskyite and moncheite, 9% complex amalgams of Au–Ag–Pt–Pd–Cu, 7% potarite, and < 2% PdSnCu (Supplementary Table 2). Hutchinson et al. (1999) recorded Ptand Pd-rich PGM in mantle peridotites of the Lizard (SW England) and Troodos (Cyprus) ophiolites. They noted that the PGM are associated with Cu-rich sulfides within poikilitic clinopyroxene-rich lenses in both mantle lherzolite (Lizard) and harzburgite that is considered to have formed just below the Moho (Troodos). The PGM identified at the Lizard are small grains (< 6 µm) of laurite, sperrylite, potarite, and unidentified tellurides. At Troodos, the PGM are even smaller (< 2 µm) grains of Pd–Cu, Pd–Pt–Cu, Pt–Pd, Pt tellurides and sulfides

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and Pd-arsenides (Hutchinson et al. 1999). However, these authors did not report the relative abundances of each type of PGM. Kogiso et al. (2008) used X-ray fluorescence combined with microbeam analysis (micro-XRF) to search for PGM in spinel lherzolite from the Horoman peridotite complex (Japan). The analyses were conducted with a broad beam on large Fe–Ni– Cu sulfides and spinel grains and revealed several ∼ 1 to 10 µm PGM alloy grains, including Os–Ir, Os–Ir–Pt, Pt–Au, and Pt. Crystallization of PGM in upper mantle ophiolite peridotites. Augé (1988) interpreted the association of PGM with chromite in dunites and chromitites at Vourinos and Tiébaghi as evidence for the mineralogical control exerted by the chromite on fractionation of the PGE in the upper mantle. He proposed that discrete PGM of the IPGE group should crystallize earlier than chromite in a mantle melt. The refractory PGM could remain suspended in the melt, serving as sites upon which subsequent chromite grains could nucleate, explaining the common spatial occurrence of both together. He also noted that the similarities in PGM assemblages observed in dunite and chromitite PGM populations at Vourinos and Tiébaghi suggested that the conditions prevailing during their precipitation were similar, thereby linking the genesis of chromitites and dunites. In contrast, Hutchinson et al. (1999) suggested that the common association of PGM with base-metal sulfides in peridotites from both the Lizard and Troodos ophiolites reflects the collection of PGE from the silicate melt by the sulfide phase. The PGE are 104–105 times more compatible in sulfide than in chromite and if sulfide saturation is achieved prior to chromite crystallization in the mantle, the PGE will be scavenged by the sulfide phase. The fact that the sulfide-bearing peridotites from the Lizard ophiolite include both IPGE and PPGE-dominated PGM thus suggests that sulfur saturation was achieved, limiting the mineralogical control that the crystallization of chromite could exert on the fractionation between the two groups of PGE. Hutchinson et al. (1999) also noted the close association of the PGM with Cu-rich base-metal sulfides. They interpreted this association as reflecting the equilibration of sulfide droplets with large volumes of infiltrating silicate melts, resulting in the capture of the PGE. According to the latter authors, the fact that PGM-bearing base-metal sulfides are associated with pyroxene-rich lenses highlights the importance of melt infiltration in the peridotites. The model they proposed suggests that meltrock reaction associated with the percolating silicate melts promoted the formation of sulfide melt droplets, which remained entrained in the partially molten peridotite matrix. These sulfide droplets equilibrated with variable volumes of infiltrating silicate melts, scavenging the available PGE. Although the exact mechanism controlling the crystallization of PGM from this early melt was not elucidated by the authors, the common presence of the PGM at the boundaries of the larger base-metal sulfides could reflect the sub-solidus exsolution of the PGE from base-metal sulfides after their crystallization. It has been demonstrated in experiments that base-metal sulfides may contain PGE in solid solution at high-T, which can be exsolved as discrete PGM upon cooling (e.g., Makovicky et al. 1986, 1988; Ballhaus and Ulmer 1995; Peregoedova and Ohnestetter 2002). The observation that PGM are commonly associated with the Cu-rich base-metal sulfides plausibly reflects either the segregation of the PGM at the latest stages of re-equilibration of the solid products of the original Fe–Ni–Cu sulfide liquid, or direct crystallization of PGM at high-T from the residual Cu-rich sulfide melt. An alternative hypothesis for the formation of the PGM observed by Augé (1988) and Hutchinson et al. (1999) is that they are refractory phases already present in the peridotite, later incorporated by either crystallizing chromite or sulfide melts during their passage through the mantle. The results obtained in the experiments of Peregoedova et al. (2004) and Fonseca et al. (2012) suggest that high temperature desulfurization of PGE-bearing base-metal sulfides promotes the formation of discrete PGM, particularly Os- and Ir-bearing alloys. Fonseca et al. (2012) proposed that during partial melting of a sulfide-bearing peridotite, the extraction of sulfur into silicate melts leads to a decrease in fS2 in the residue that triggers the exsolution of

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Os–Ir alloys from the refractory matte (sulfide partial melt forming a solid/melt mush). Mantle base-metal sulfides may also lose S if they react with S-undersaturated melt/fluids such as those that commonly migrate through peridotites in ophiolite environments (e.g., Lorand and Conquéré 1983; Kelemen et al. 1992; Peregoedova et al. 2004). A series of small-scale melt extraction and/or melt–rock reaction events could therefore produce a sequential decrease of fS2 in the mantle source region, promoting the breakdown of PGE-bearing sulfides, leading to formation of a series of residual PGM sulfides and PGE alloys. The residual PGM could remain attached to the original sulfides or be transported as discrete grains in the silicate melt. Mechanical entrainment by either chromite, as observed in Vourinos and Tiébaghi, or by migrating sulfide melts (as occurred in the Lizard and Troodos) are consequent possibilities. Ohnenstetter (1992) ruled out a magmatic origin for the PGM identified in mantle peridotites from the Monte Maggorie Massif. It was suggested that the precipitation of the PGM was linked to a hydrothermal system created by the ascension of a fluid-enriched magma under a midoceanic ridge during asthenospheric upwelling. In this model, the migration of magma towards the surface as well the sudden opening of fractures may have caused a pressure drop responsible for the oversaturation of the magma in a fluid phase. The hydrothermal fluid that separated from the silicate melt could conceivably concentrate significant amounts of Pt, Pd, and Rh, together with Au, Ag, Cu, Ni, Pb, Sn, Hg, Te, and Bi. The metal-enriched hydrothermal fluid would have a relatively low viscosity, so could be injected into fractures and precipitate the PGM along cracks. It was estimated that the crystallization of the PGM from the hydrothermal fluid took place at temperatures lower than ~600 °C, based on the following criteria: (1) the crystallization of PGM after pentlandite, the latter having a maximum thermal stability of 610 °C; (2) a maximum temperature of < 470 °C for the crystallization of tetraferroplatinum co-existing with Pt–Cu and Au–Cu alloys, based on phase-equilibria data for the stability fields of binary Pt–Fe, Pt–Cu, and Au–Cu compounds; (3) a maximum temperature of < 400 °C for the crystallization of tellurides associated with the Cu-rich base-metal sulfides. Such low temperatures do not support direct crystallization from the melt, but favor metasomatic processes. Ohnestteter (1992) ruled out the potential involvement of fluids associated with the serpentinization of the hot peridotites, as the anomalous enrichments in Pt and Au (expressed in the PGM abundances) were not found to correlate with the degree of serpentinization. In addition, PGM were not typically observed in networks of serpentinized veins and fractures.

PGM in ophiolite chromitites The mantle portion and the mantle–crust transition zone of many ophiolites host bodies of chromite of both high-Cr and high-Al type (i.e., podiform chromitite). These chromitites are usually characterized by relatively high abundances of all of the PGE, generally on the order of 101–103 ppb (e.g., González-Jiménez et al. 2014a,b). Differences in chromitite morphology, structural relationships, and geochemical signatures allow the recognition of three distinct chromitite types in ophiolites (González-Jiménez et al. 2014a,b). Type I is the most abundant and is characterized by bulk-rock enrichment in the refractory IPGE over the PPGE, independent of the bulk-rock content (usually between 10–3–101 times CI-chondrite; Fig. 9). This type of chromitite may show a range of geometries (e.g., pods, boudins, fusiforms, veins, dikes, stockworks, and schlieren) and is found throughout the lower ophiolite pseudostratigraphy. Type IIA chromitites include concordant layers, bands, and seams, as well as less extensive discordant pods or irregular bodies that are generally confined to the shallower zones of the oceanic lithosphere. Chromitites of Type IIA show higher bulk-rock PGE contents than Type I as they are significantly enriched in Pt and Pd. The third type (Type IIB) of chromitite shows the same spatial distribution and PGE distributions as Type IIA but has a more limited range in Cr# (Cr# = Cr/Cr + Al) and a wider range of Mg# (Mg# = Mg/Mg + Fe2+) that overlaps the compositional range of chromitites from LMI (Fig. 9).

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Figure 9. CI-chondrite (Naldrett and Duke 1980) normalized patterns of ophiolitic chromitites. Moa-Baracoa, Sagua de Tánamo, Mayarí and Potosí are examples of Type I chromitites with high-Cr and high-Al chromite hosted in the mantle section and Moho Transition Zone in the ophiolite belt of Mayarí-Baracoa (Proenza et al. 1999, 2001; Gervilla et al. 2005; González-Jiménez et al. 2011b). Other examples of Type I chromitites with high-Cr chromite include Al’Ays, Saudi Arabia (Prichard et al. 2008); Pindos, Greece (Economou-Eliopoulos 1993; Economou-Eliopoulos and Vacondios 1995); Oman (Ahmed and Arai 2003); Skyros Island, Greece (Economou 1983); Thetford Mines, Canada (Gauthier et al. 1990). Type I chromitites with high-Al include Golyamo Kamenyane, Bulgaria (González-Jiménez et al. 2014b) and Oman (Ahmed and Arai 2003). Data sources for sulfide-poor Type II chromitites: Herbeira, Spain (Moreno et al. 2001); Kraubath, Austria (Malitch et al. 2003); Leka, Norway (Pedersen et al. 1993), Shebenik, Albania (Kocks et al. 2007); Thetford Mines, Canada (Gauthier et al. 1990). Data sources for sulfide-rich Type II chromitites: Acoje, Philippines (Bacuta et al. 1990); Ceruja, Albania (Karaj 1992); Cliff and Mid-Unst in Shetland, Scotland (Prichard et al. 1986); Eretria, Tsangli area, Greece (Economou-Eliopoulos and Naldrett 1984) Faeoy, Finland; Haylayn Block, Oman (Lachize et al. 1991).

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The in situ study of chromitite from a large number of ophiolites reveals that PGM are generally rare, with grain sizes typically < 30 µm (Fig. 10a–f; see Ahmed and Arai 2003; González-Jiménez et al. 2014a,b, for review). In most cases, the PGM occur as inclusions in chromite grains and less frequently in the interstitial silicate matrix (unaltered or altered) between the chromite grains. Nevertheless, PGM can be also found in fractures cutting the chromite grains (where they may exhibit irregular shapes and/or alteration rims) or alteration zones around chromite edges (e.g., Fe2+ or Fe3+-rich chromite with either porous or homogeneous textures) produced in chromite during post-magmatic alteration events such as serpentinization or metamorphism (e.g., ferrian chromite; Fig. 10e). Examples of PGM forming planar arrays (linear trails, corresponding to former grain boundaries or healed fractures) with base-metal minerals, silicates and fluid/melt inclusions have been observed. The grain morphology of PGM in ophiolitic chromitites is variable. Generally, PGM hosted in primary chromite show polygonal shapes. In contrast, those PGM found within fractures or altered zones of the chromitite may show a wide range of morphologies (i.e., subhedral or anhedral shapes, typically with internal sieve-texture and/or corroded grain boundaries) as a result of different rates of mechanical deformation or reaction with metasomatic fluids. PGM distribution in ophiolitic chromitites. González-Jiménez et al. (2009a), and more recently González-Jiménez et al. (2014b), compiled a list of PGM in chromitites from worldwide occurrences, including reporting on the absolute number of grains found. An updated list of the PGM identified in both Type I and Type II chromitites, including previously unpublished data, is presented in Supplementary Table 2. Most PGM identified in Type I chromitites occur in unaltered zones (i.e., cores) of chromite grains (66%), sometimes being manifest as inclusion trails. They consist predominantly of members of the laurite–erlichmanite solid solution series (76%), irarsite (9%), and Os–Ir–Ru alloys (9%). Substitution of Os by Ru in disulfides of the laurite– erlichmanite series is extensive-to-complete in most ophiolitic chromitites (Fig. 11). This is frequently reflected in the form of a range of styles of chemical zonation within single grains (Fig. 10a–d), including normal, reverse and oscillatory zonation. The coexistence of Os-poor laurite with Ru-poor, Os–Ir alloys has been reported in unaltered chromite grains, consistent with the assemblages reproduced at high-T in experiments (e.g., Brenan and Andrews 2001; Andrews and Brenan 2002). Grains of irarsite from within unaltered chromites show substitutions of Ir by Os (osarsite; OsAsS), Ru (ruarsite; RuAsS), and to a lesser extent, Rh (hollingworthite; (Rh,Pt,Pd)AsS). Rare Pt–Pd–Rh minerals observed in these unaltered zones include sperrylite and unidentified Pt–Pd–Rh alloys. Laurite–erlichmanite (18%) and irarsite (11%), characterized by S-deficiency are the most abundant PGM in the altered zones of Type I chromitites. Almost complete desulfurization and/or oxidation of these minerals has resulted in the formation of secondary Ru-rich Os–Ir alloys (13%) and oxides (5%) that are often observed in cracks, at the altered edges of chromite or in the serpentinized/ weathered chromitite silicate matrix (Fig. 10e). Most of the remaining Pt–Pd–Rh PGM in the altered zones of chromitites comprise Pd–Cu–Sn, stibiopalladinite (Pd5Sb2), and sperrylite. These minerals are particularly frequent in chromitites from the metamorphosed ophiolites of Kraubath in the Austrian Alps (Thalhammer et al. 1990; Malitch et al. 2001), the Great Serpentinite Belt in Australia (Yang and Seccombe 1993), the Pampean Ranges in Argentina (Proenza et al. 2008) and the small Dobromirtsi massif in the Bulgarian Rhodopes (González-Jiménez et al. 2010). The abundance of Pt–Pd–Rh minerals in the altered zones of chromitites has been traditionally associated with the infiltration of either late magmatic or post-magmatic fluids, resulting in the localized redistribution of these PGE (e.g., Graham et al. 1996). However, it may simply reflect the physical fractionation of the PGE; Ru–Os–Ir PGM have a tendency to wet and/or nucleate along chromite surfaces while Pt–Pd–Rh PGM have a tendency to remain suspended in residual silicate melts (Matveev and Ballhaus 2002; Ballhaus et al. 2006; Finnigan et al. 2008).

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Figure 10. Images of PGM included in chromite from chromitites hosted in (a) the ultramafic complex of the Vizcaino Peninsula, Baja California Sur (Mexico), (b-e) the mantle of the Mayarí-Baracoa ophiolite belt in eastern Cuba and (f) in the silicate matrix of chromitite from the mantle-crust transition zone in the New Caledonia ophiolite. (a) Backscattered electron micrograph of a laurite and an Os-Ir grain at the edge of a silicate inclusion in chromite. (b) Backscattered electron micrograph of a composite grain made up of an intergrowth of zoned laurite with PGE-rich mss and cuproiridsite. (c) and (d) are grains of zoned laurite-erlichmanite intergrowth with irarsite (Monte Bueno chromite deposit; Gervilla et al. 2005; González-Jiménez et al. 2009b). (e) Backscattered image of partially desulfurized irarsite attached to a larger grain of a pseudomorphed Ru-Os-IrRh-Fe-Ni alloy after desulfurization of laurite in a chromite crack (Tres Amigos Mine, Mayarí-Baracoa Ophiolite, Cuba). (f) Backscattered image of Pt-oxide after cooperite (Ouen Island; González-Jiménez et al. 2011a).

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Figure 11. Composition of inclusions of the laurite–erlichmanite solid solution series in chromite grains in chromitites from different ophiolite complexes, plotted in Ru-Os-Ir (atomic %) ternary diagrams. The data for these plots are sourced as follows: Acoje block, Philippines (Bacuta et al. 1990; Orberger et al. 1988); Dobromirtsi, Rhodope Complex in southern Bulgaria (González-Jiménez et al. 2010, 2013b, 2014c); Finero (Ferrario and Garuti 1990; Garuti and Zaccarini 1994); Great Serpentinite Belt, Australia (Yang and Seccombe 1993); Hochgrössen (Thalhammer et al. 1990; Malitch et al. 2003); Josephine (Stockman and Hlava 1984); Kempirsai, Kazakhstan (Melcher et al. 1997); Kraubath, Austria (Thalhammer et al. 1990; Malitch et al. 2003); Mayarí-Baracoa, eastern Cuba (including Moa-Baracoa, Sagua de Tánamo and Mayarí; Gervilla et al. 2005; González-Jiménez et al. 2011b); New Caledonia (including Tiébaghi, Massif du Sud, Ouen and Pirogues; Augé 1988; Augé et al. 1998; González-Jiménez et al. 2011a); Nurali, Russia (Zaccarini et al. 2004b; Grieco et al. 2006); Ojén, southern Spain (Torres-Ruiz et al. 1996; GutiérrezNarbona et al. 2003; González-Jiménez et al. 2013a) Oman (Augé 1987; Ahmed and Arai 2003); Ortaca, southeastern Turkey (Uysal et al. 2005); Othrys, Greece (Garuti et al. 1999a); Ray-Iz, Russia (Garuti et al. 1999b); Troodos, Cyprus (Legendre and Augé 1986; Augé and Johan 1988; McElduff and Stumpfl 1990), Vourinos, Greece (Augé 1985; Augé 1988; Garuti and Zaccarini 1997).

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In accord with their whole-rock enrichment in the PPGE, Type II chromitites show a preponderance of Pt, Pd, and Rh PGM (Fig. 9). Most Type II chromitites described are associated with metamorphosed ophiolites and the Pt–Pd–Rh PGM are situated near chromite rims (usually altered) or in the interstitial silicate matrix (63%), rather than in the unaltered portions (i.e., chromite crystal cores). These features make it difficult to interpret the fractionation of the PGE in terms of high-temperature magmatic processes. In their recent review, González-Jiménez et al. (2014b) showed that sperrylite (29%), stibiopalladinite (10%) and isoferroplatinum (8%), and members of the solid solution series cooperite–braggite are the most common minerals in the unaltered cores of chromite. The Pt–Pd–Rh ± base-metal alloys (23%), isoferroplatinum (12%), sperrylite (9%), and Pt-oxides (derived from the alteration of Pt-sulfides such as malanite) are the most abundant PGM in the altered zones of the chromitite. In chromitites that have been altered by high-temperature hydrothermal fluids (i.e., Ouen Island, New Caledonia), grains of isoferroplatinum in a chlorite-dominated interstitial matrix show corroded outlines and enrichment in Pd, whereas malanite in a serpentine-rich matrix is partially corroded and S-poor due to desulfurization. When in contact with chlorite, the malanite has rims of Pt-oxide (González-Jiménez et al. 2011a). The Os–Ir–Ru PGM are the same as those that predominate in Type I chromitites, whether they are observed as inclusions in chromite or they occur in the groundmass: laurite–erlichmanite, Os–Ir–Ru alloys and irarsite. As observed in Type I chromitites, when IPGE-rich sulfides occur in the altered silicate matrix they show evidence of post-magmatic desulfurization or oxidation, such as corroded outlines, a S-deficiency or rims of Pt-group alloys/oxides. Re–Os isotopes in ophiolitic chromitite PGM. The Re–Os isotope systematics of PGM from ophiolitic chromitites were initially examined on individual grains separated mechanically from chromitite samples and more recently on polished thin sections and blocks, using N-TIMS (Walker et al. 1996; Malitch et al. 2003), ion probe (Ahmed et al. 2006), and LA-MC-ICPMS (Pearson et al. 2007; Shi et al. 2007; González-Jiménez et al. 2014b and references therein). The PGM in ophiolite chromitites generally yield relatively low 187Re/188Os ratios and show a significant dispersion in 187Os/188Os ratios (Fig. 12). The in situ LA-MC-ICPMS analysis of Re–Os isotopes of PGM included in unaltered chromite grains have revealed significant heterogeneity in the Os compositions of PGM inclusions within single chromite grains, separated several mm apart. This isotopic heterogeneity is observed within composite aggregates of the PGM smaller than 50 μm in diameter and even within single zoned laurite grains (e.g., Marchesi et al. 2011; Gonzalez-Jimenez et al. 2012a). Although Re-depletion model ages (TRD) of some chromitite-hosted PGM broadly coincide with the timing of known tectonomagmatic mantle or crust-forming events (e.g., Shi et al. 2007; Marchesi et al. 2011; Belousova et al. 2015; McGowan et al. 2015), more commonly chromitite-hosted PGM yield TRD ages older than the supposed age of formation of the host chromitite/ophiolite (Fig. 12). For example, the ∼ 21 Ma chromitites of the Ojen massif in southern Spain contain PGM that yield TRD ages of ~1.4 Ga, and the Group II Os-Ir alloys analyzed by Shi et al. (2007) in the chromitites of the Mesozoic Donqiao ophiolitic massif reveal a range of TRD between 1130 and 240 Ma, with age peaks at ~870, 930, 1089, 1100, and 1130 Ma. Interestingly, González-Jiménez et al. (2013a) observed that grains of laurite and base-metal minerals (including sulfides and arsenides) in chromitite and in the host peridotite of the Ojen massif (southwestern Spain) have similar values and ranges of heterogeneity in Os isotopes. Most laurites (n = 104) and base-metal minerals (n = 27) have 187 Os/188Os between 0.1181 ± 0.000003 and 0.1364 ± 0.0006, roughly matching those of the base-metal sulfides of the peridotite (0.1159 ± 0.0007 to 0.1374 ± 0.0014; n = 33). GonzálezJiménez et al. (2014a) proposed that the presence of multiple populations of PGM in the chromitites reflects the combination of (1) direct crystallization of PGM from the basaltic parental melts of the chromitites and (2) entrainment of some older PGM present in the peridotite during the migration of such a melt through the source/host mantle.

Figure 12. 187Os/188Os isotopic ratios and corresponding TRD model ages for PGM analyzed in situ in chromitites from ophiolite-type ultramafic complexes; updated after González-Jiménez et al. 2014a. Data sources: Coolac Serpentinite Belt, Australia (Belousova et al. 2015); Dobromirtsi (Bulgaria; González-Jiménez et al. 2013a, 2014c); Eastern Desert, Egypt (Ahmed et al. 2006); Hochgrössen and Kraubath massifs, Austria (Malitch 2004; Malitch et al. 2003); La Cabaña (this study); Loma Baya, Guerrero State, Mexico (González-Jiménez et al. 2014b); Luobusa and Donqiao, Tibet (Shi et al. 2007; McGowan et al. 2015); Mayarí-Baracoa, eastern Cuba (Marchesi et al. 2011; González-Jiménez et al. 2012a); Neyiriz (this study); Ojén (González-Jiménez et al. 2013a); Ouen Island (González-Jiménez et al. 2014b); Outokumpu ophiolite, Finland (Walker et al. 1996); Thetford Mines, Canada (this study); Vizcaino, California State, Mexico (this study).

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The analysis of Re–Os isotopes of individual PGM from highly metamorphosed chromitites yields somewhat contradictory results. González-Jiménez et al. (2012b) observed that in highly metamorphosed chromitites of the Dobromirtsi peridotite (Central Rhodope Complex, Bulgaria), laurite grains wholly enclosed in unaltered chromite have 187Os/188Os compositions very distinct from the secondary laurite in the altered zones (Fig. 12). Given the low 187Re/188Os compositions in both populations (< 0.024), these authors suggested that the wider range of 187Os/188Os observed in the secondary laurite population formed by interaction with fluids of variable 187Os/188Os composition during metamorphism. Malitch et al. (2014) analyzed a similar assemblage of PGM containing laurite ± Os–Ir alloy ± Ru-pentlandite in chromitites from the Shetland Ophiolite Complex (Scotland), but they did not observe a comparable dispersion of 187Os/188Os values. Instead, the latter authors note that the PGM in the Shetland chromitites exhibit unusually homogeneous 187Os/188Os that they interpreted as evidence for the resistant nature of the 187Re–188Os system in Os-rich minerals to metamorphic alteration. González-Jiménez et al. (2014a) suggested that polyphase metamorphism might not just disperse the 187Os/188Os signatures of individual PGM, but may act to homogenize them also. Following this line of reasoning, alteration of the Shetland PGM may have occurred in a closed system and was not directly related to metamorphism but to a separate ‘autohydrothermal’ event, perhaps during the late stages of magmatic evolution of the chromitite ore system (cf. Derbyshire et al. 2013). González-Jiménez et al. (2011a) proposed a similar scenario to explain the origin of secondary ‘desulfurized’ PGM sulfides within ‘primary hydrous silicates’ in the chromitites of Ouen Island, New Caledonia. Petrogenesis of PGM during chromitite formation in the upper mantle. As noted above, Type I ophiolitic chromitites usually contain two types of PGM assemblage (Supplementary Table 2): (1) inclusions in chromite composed of almost exclusively Os-, Ir-, and Ru-rich PGM, which are commonly associated with small amounts of Ni–Fe–Cu sulfide (even in the same inclusion); (2) an interstitial assemblage consisting mainly of PPGE-rich minerals. The latter assemblage may also contain secondary Os–Ir–Ru PGM, produced following alteration of primary PGM via hydrothermal alteration or weathering. The PGM included in chromite may therefore be considered as primary, as these grains are ‘shielded’ from post-magmatic alteration. The origin of these two assemblages in chromitite has been the subject of much debate during the last four decades (see González-Jiménez et al. 2014a for review). Some authors have suggested that PGE-rich chromitites could result from a PGE-rich melt produced following high-degree partial melting of a PGE-enriched mantle source (e.g., Prichard et al. 2008). However, experimental data (e.g., Ballhaus et al. 2006; Finnigan et al. 2008) show that the PGE-enrichment in chromite and the fractionation between IPGE and PPGE could simply reflect the local effect of chromite crystallization in PGE-bearing S-undersaturated melts. Gijbels et al. (1974) and Naldrett and Cabri (1976) suggested that the PGE could enter the structure of chromite at high temperature and subsequently be exsolved as discrete PGM upon cooling. Later experimental work showed that Os, Ir, Ru, and Rh could substitute in the chromite structure at high temperature and suitable fO2 (Capobianco 1998; Capobianco and Drake 1990; Capobianco et al. 1994; Righter and Downs 2001; Righter et al. 2004). Although later experiments indicated that such high values of fO2 are probably unrealistic at mantle conditions (Brenan et al. 2012), the potential role of high-temperature formation of inclusions of Os-, Ir- and Ru-rich PGM in chromite is worth further discussion. Another problem with the idea is that the textures exhibited by the PGM are not typical of exsolution (e.g., they don’t occur along the crystallographic planes of the chromite). Instead, chromite-hosted PGM are characteristically euhedral and faceted, as shown by multiple studies employing optical (conventional reflected light microscopy, SEM) and laser ablation techniques. An important alternative model suggests that both sub-µm and visible PGM grains form initially as nano-size metal clusters (cf. Tredoux et al. 1995; Helmy et al. 2013), which later

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coarsen to larger PGM. Coarsening presumably takes place by diffusion-driven, solid-state Ostwald ripening, similar to that observed for native Au nanoparticles in pyrite (Reich et al. 2006). Micronuggets of PGM have been reported for chromites from a range of settings (Ballhaus and Sylvester 2000; Sattari et al. 2002; Locmelis et al. 2011; Pagé et al. 2012). However, their thermal behavior in the chromite matrix is not well understood due to a lack of experimental data on their physical and chemical properties. Coarsening of nano-scale metal clusters could operate at the sub-µm scale in chromite at high-temperature when diffusion is easier. However, macroscopic PGM that exhibit chemical and Os isotopic zoning characterized by precipitation–dissolution microstructures (e.g., Fig. 10) are better explained by the involvement of fluid/melts during mineral growth. Locmelis et al. (2011) and subsequently Pagé et al. (2012) conducted laser ablation ICP-MS analysis on komatiite chromites and observed an almost homogeneous distribution of Ru and sub-µm Ir-bearing alloys, which they took to imply that Ru exists in solid solution in chromite at high temperatures. However, they did not carry out high-resolution transmission electron microscopy (HRTEM) to test the potential existence of native Ru particles at near atomic scales. It is possible that the LA-ICP-MS technique, with typical beam diameters of ∼ 20–40 µm, does not operate at sufficiently high resolution to detect nano-scale PGM. This might in turn imply a different range of stability of Ru nanoparticles relative to metal nanoparticles of the other PGE. In light of this, the linear correlations observed between Os + Ir + Ru and Cr# in chromites from natural samples and experiments might simply reflect nugget fractionation, rather than PGE substitution in the crystal lattice of chromite (Ballhaus et al. 2006). Precipitation of metallic nanoparticles of the PGE from silicate melts has been documented in a series of experiments (Mungall 2005; Finnigan et al. 2008). As described above, local reduction of fO2 at the edges of chromite crystals may cause the saturation of the most easily oxidized PGE (Os, Ir, Ru) in the form of metallic nanoparticles (Finnigan et al. 2008). These results are comparable to those obtained by Matveev and Ballhaus (2002), Bockrath et al. (2004b) and Ballhaus et al. (2006), who also reported the nucleation of abundant nanoparticles of Os–Ir–Ru alloys on chromite surfaces. They also crystallized fewer but much larger-sized metallic alloys of Pt and Pd in the silicate glass. These experimental results clearly demonstrate that the precipitation of chromite may cause fractionation of the PGE from one another, as it provides a nucleation substrate for the most refractory IPGE alloys. Nugget formation is likely only possible at S-undersaturated conditions. These experimental constraints are also consistent with the observation that natural chromitites precipitated from S-undersaturated arc melts in the upper mantle contain an assemblage of Os-, Ir-, and Ru-rich PGM included in chromite crystals and another mostly dominated by Pt- and Pd-rich PGM in the interstitial silicate matrix (e.g., González-Jiménez et al. 2014a). Another excellent and historically significant example of this sort of distribution of the PGM (i.e., Os-, Ir-, and Ru-rich phases in chromite, and Pd-rich phases in interstitial sulfarsenides and arsenides, is described from Heazlewood, Tasmania (Peck and Keays 1990). A similar mechanism of chromitite genesis can be assumed for the formation of PGM in the Type II chromitites. In this model, ascending basaltic melts crystallize IPGE-rich chromitites in the lower parts of the mantle through which they migrate, producing Pt- and Pd-rich fractionated melts (Fig. 13). The fractionated melts would migrate upwards giving rise to Pt- and Pd-rich chromitites in the uppermost mantle portion or the lower crustal (cumulate) section of the ophiolite. Usually, Type II chromitites have much lower Mg# than Type I chromitites, which is taken as evidence that they crystallized from melts that underwent higher degrees of fractionation. This might also induce sulfur saturation relatively early during magmatic evolution, explaining the coexistence of oscillatory zoned laurite with Pt-rich alloys, Pt ± Rh ± Ir sulfides and Ni–Cu–Fe (e.g., Ouen Island and Pirogues in the New Caledonia ophiolite; Augé and Maurizot 1995; González-Jiménez et al. 2011a).

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Figure 13. Sketches illustrating how chromite and PGM crystallize from hybrid melts, following the mixing/mingling of magmas with different degrees of fractionation. The sketches also show how Os isotope heterogeneity can develop in the upper mantle conduits where podiform chromitites are considered to form. These conduits can be centimeters to tens of meters wide, but are typically on the order of tens of centimeters to several meters wide. Details are given in the main text. The parameter γOs signifies the percent deviation of 187Os/188Os from a chondritic reference at the time of formation, and BMS is base-metal sulfide.

Many PGM inclusions in chromite exhibit euhedral morphologies, an observation originally interpreted as evidence for their direct precipitation from fluids or melts before chromite crystallization (e.g., Constantinides et al. 1980; Legendre 1982; Stockman and Hlava 1984; Augé 1985). The mechanical trapping of PGM-saturated melt droplets by chromite was suggested by Melcher et al. (1997). In this scenario, fast-growing chromite can mechanically trap molten inclusions, explaining the drop-like textures that the inclusions sometimes exhibit. Melcher et al. (1997) also suggested that even euhedral inclusions may have originated as melt inclusions, suggesting that they filled negative crystal cavities in the chromite host. During the 1980s and 1990s the formation of chromitite was interpreted as the result of a simple mechanism of fractional crystallization of basaltic melt. In this model, the PGM represented early refractory minerals crystallized before chromite. The coexistence of PGM alloys, sulfides, and sulfarsenides simply reflected different stages of the cooling history of the chromitite body. The PGM alloys crystallized first at higher temperature and under lowest fS2,

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which increased upon cooling to give rise to a series of PGM sulfides and/or Ni–Cu–Fe sulfides (Stockman and Hlava 1984; Augé and Johan 1988; Leblanc 1991; Torres-Ruiz et al. 1996; Augé et al. 1998; Garuti et al. 1999a). The results obtained in subsequent experiments (Brenan and Andrews 2001; Andrews and Brenan 2002; Bockrath et al. 2004b) showed that Os-free laurite (RuS2) may co-precipitate in equilibrium with Os–Ir alloys at the T–fS2–fO2–P conditions prevailing during the crystallization of chromitites in the upper mantle [1200–1300 °C, log fS2 from –2 to –1.3, and ~ 0.5 GPa]. Furthermore, a simple trajectory of cooling does not explain why many chromitite bodies contain PGM with a range of patterns of compositional zoning, including normal, reverse, and oscillatory zonation (Fig. 10a–d; González-Jiménez et al. 2009b). Fractional crystallization also doesn’t explain the wide dispersion of 187Os/188Os isotope compositions amongst individual PGM (and base-metal sulfides) that coexist within single grains of chromite or intra-crystal Os isotopic heterogeneity in oscillatory-zoned laurites (Ahmed et al. 2006; Marchesi et al. 2011; González-Jiménez et al. 2012a). Crystallization of chromitite by melt mixing (or mingling) in open conduits in the upper mantle may create a heterogeneous environment with variations in T–fO2–fS2–aAs over short time spans (GonzálezJiménez et al. 2014a,b). Local changes in fO2 during crystallization of chromite could cause the precipitation of small metallic nanoparticles of PGM, which could then attach to the surface of chromite in equilibrium with the ambient silicate melt, as has been observed experimentally (Fig. 14; Matveev and Ballhaus 2002; Ballhaus et al. 2006; Finnigan et al. 2008). The influx of fluids/melts into the open conduit, producing sudden changes of fS2and/or aAs could destabilize discrete alloys, leading to the formation of PGE-sulfides or PGE-sulfarsenides (e.g., Bockrath et al. 2004b). Such a dynamic environment could also promote the immiscible segregation of droplets of sulfide melt, producing the base-metal sulfides commonly associated with the PGM in chromite (Garuti et al. 1999b) which could scavenge any PGE remaining in the melt (Moreno et al. 1999; Proenza et al. 2001; González-Jiménez et al. 2012a). Incorporation of PGM from the host peridotite: an alternative model? As noted above, some ophiolite chromitites contain PGM with TRD ages much older than the supposed age of formation of their host chromitite/ophiolite. For example, Luobusa chromitite ‘Group II’ alloys yield TRD ages (≥1.1 Ga) much older than the age of formation of the ophiolite (177 ± 31 Ma; Zhou et al. 2002), which has led to them being considered as xenocrystic in the chromitites (Shi et al. 2007). The latter authors stated that ‘Group II alloys were scavenged from the surrounding oceanic lithospheric mantle during melt-rock reaction’. The model of Shi et al. suggests that not all PGM hosted in ophiolitic chromitites necessarily form by direct precipitation from their parental melts. According to González-Jiménez et al. (2014b), incorporation of PGM present in the host peridotite might explain the ancient (> 1 Ga) TRD ages of some chromitite-hosted PGM from Phanerozoic ophiolites (Figs. 12, 13). GonzálezJiménez et al. (2014b) suggested that such heterogeneity in 187Os/188Os signatures of PGM could reflect ‘…a series of small-scale events of melt extraction and melt-rock reaction that has produced a progressive but stepped sequence of decreasing fS2 in the mantle, promoting the breakdown of their PGE-bearing base-metal sulfides into ‘residual’ laurites, and to a major extent Os–Ir alloys…’. The latter authors invoked this as a process to explain ‘…the broad coincidence of the TRD ages of many PGM with known magmatic events related to ophiolite formation’. Some PGM could therefore reside in the mantle for hundreds of millions of years, explaining the coexistence of PGM with ages close to that of the chromitite/ophiolite host and those with older ages. Chromitite-hosted PGM also show similar ranges in 187Os/188Os as peridotite-hosted base-metal sulfides. According to Gonzalez-Jiménez et al. (2014b), Os may be transported as submicroscopic Os-bearing alloys which could later react with S to form sulfides like laurite in water- and volatile-rich melts at high temperature, low pressure and low fS2 (cf. Bockrath et al. 2004a). This is an alternative mechanism that may produce contemporaneous PGM and chromite, where the PGM preserve older Os signatures from their host mantle, analogous to partial recent recrystallization of zircons that can also preserve older 176 Hf/177Hf isotope signatures (cf. Belousova et al. 2015; McGowan et al. 2015).

Figure 14. Sketches showing the localized (grain-scale) effect that chromite crystallization may have on PGE fractionation and PGM formation. Details are given in the main text (see also Finnigan et al. 2008).

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PGM in sulfide-rich ophiolite lithologies The association of massive base-metal sulfide ores with mafic and ultramafic rocks is an uncommon feature of ophiolites. The sulfide-rich ores may be either of magmatic or hydrothermal origin and located in any part of the ophiolite stratigraphy. Magmatic basemetal sulfides representing up to 5% by volume of the rock have been described in the crustal portions of some ophiolites, including the sheeted dike complex of Faeoy, Norway (Barnes et al. 1988; Boyd et al. 1988), the layered gabbros of the Haylayn block in Oman (Lachize et al. 1991) and the pyroxenite cumulates of ophiolites of south western Oregon, USA (Foose 1985). Horizons more enriched in magmatic sulfides than this have been described in dunite seams and pods of the mantle–crust transition and upper mantle at Ceruja (Albania; Karaj 1992), Unst (Shetland; Prichard and Lord 1993; O’Driscoll et al. 2012; Derbyshire et al. 2013), and chromite-rich dunites (including chromitites) at Acoje (Philippines; Bacuta et al. 1990), Potosí (eastern Cuba; Proenza et al. 2001) and northen Oman (Negishi et al. 2013). Overall, sulfide mineralized lithologies exhibit enrichment in PPGE over IPGE, represented by steep positive chondrite-normalized PGE patterns (Fig. 9), irrespective of the relative PGE abundances. This has been interpreted as reflecting immiscible segregation of sulfide melts after extensive fractionation of the silicate melt. However, Proenza et al. (2001) showed that the origin of the PGE-rich sulfide mineralization associated with the chromitites at Potosí (Cuba) was associated with the modification of the chromite ores by the late intrusion of olivine-norite and gabbro dikes. According to these authors, the interaction between pre-existing, sulfidepoor chromitite and the intruding volatile-rich melts drove brecciation, partial dissolution, and recrystallization (coarsening) of chromite. The sulfide assemblage formed by fractionation of an immiscible sulfide melt that segregated from the volatile-rich silicate melt; the segregation process itself has been attributed to the reaction between the intruding melts and the host chromite (Proenza et al. 2001). Evidence for the magmatic nature of the sulfide mineralization comes from sulfur isotopes (d34S), which range from –0.4‰ to +0.9‰. According to Proenza et al. (2001), the variable extent of the melt–rock reaction produced chromite ores with variable sulfide ratios, which preferentially collected the relatively incompatible PPGE. Massive sulfides interpreted as hydrothermal precipitates, related to serpentinization of the host rocks, have been reported in the Tsangli area, Eritrea (Greece; Economou and Naldrett 1984), and in the serpentinized harzburgites and dunites of the Limassol Forest (Cyprus; Foose et al. 1985; Thalhammer et al. 1986). Unlike magmatic ores, hydrothermal sulfides show relatively flat PGE patterns, with a depletion in Pt in the sulfide-rich chromitite ores described from Eretria (Fig. 9). A detailed PGM documentation of sulfide-rich ophiolite peridotites has only been provided for the Shetland Ophiolite Complex (Prichard et al. 1994; Derbyshire et al. 2013) and the pertinent details are summarized as follows: 1.

Some sulfide-rich dunites on the island of Unst have > 1 ppm Pt + Pd (O’Driscoll et al. 2012). These lithologies contain stibiopalladinite, sperrylite, geversite [Pt(Sb,Bi)2], genkinite [(Pt,Pd)4Sb3], Pt–Fe–Cu alloys, hongshiite (PtCu) and Pt–Pd ochres. Antimonides of Pd occur in silicates surrounding the clusters of sulfide, in fractures crosscutting the chromite grains, as rims to breithauptite (NiSb), or in Ni–Fe alloy rims on heazlewoodite or pentlandite. Geversite and the Pt–Fe–Cu alloys were found in a chlorite rim surrounding a cluster of sulfides associated with chromite grains.

2.

Sulfide-bearing wehrlites and pyroxenites contain grains of sperrylite, Pd-antimonides and Pt-ochres, located at the margins of aggregates composed of chalcopyrite (replaced by bornite) and pyrrhotite (hosted within clinopyroxene).

3.

Gabbro-hosted wehrlite in the uppermost part of the dunite transition zone (i.e., immediately below the geophysical Moho) of the Shetland Ophiolite is rich in base-metal sulfides (pyrrhotite, pentlandite, pyrite, chalcopyrite ± gersdorffite [NiAsS], and niccolite [NiAs]). Prichard et al. (1994) reported one grain (~5 µm) of sperrylite at the junction between wehrlite and pyroxene layers.

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Prichard et al. (1994) interpreted the close spatial association of Pt- and Pd-rich PGM and base-metal sulfides as evidence that the precipitation of the PGM occurred after the collection of the PGE from a silicate melt by an immiscible sulfide fraction. They suggested that the Pt preferentially entered the structure of chalcopyrite whereas pyrrhotite and pentlandite incorporated Pd more easily. On cooling and alteration, these PGE were released from the sulfides and recrystallized as PGM in close proximity to the sulfides. Although there is evidence to suggest that Pd exists in solid solution in pentlandite, the balance of evidence suggests that incorporation of Pt into chalcopyrite and Pd into pyrrhotite as suggested by Prichard et al. (1994) is less likely.

PGM IN PERIDOTITES OF THE SUBCONTINENTAL LITHOSPHERIC MANTLE The subcontinental lithospheric mantle (SCLM) formed in the mid-late Archean and constitutes the lower part of the continental lithosphere (see also Aulbach et al. 2016, this volume, and Luguet and Reisberg 2016, this volume). By providing stable continental cratons on an Earth floored by oceanic crust, the SCLM has facilitated a significant reorganization of plate tectonics over geological time (Griffin et al. 2013). Micrometric PGM have been identified in tectonically emplaced peridotite massifs and mantle xenoliths, both considered to represent SCLM materials.

Subcontinental lithospheric mantle peridotite massifs Mineralogical distribution and petrogenesis of PGM in peridotite massifs. Sub-µm PGM grains have been identified in lherzolites of the Lherz massif (southern France; Luguet et al. 2007; Lorand et al. 2010; König et al. 2015) The PGM identified in these studies are micronuggets (0.5–0.3 µm diameter) and occur as inclusions in or at the margins of larger base-metal sulfides. The PGM consist primarily of the laurite–erlichmanite solid solution series, Pt–Ir–Os-rich alloys, and Pt–Pd–Te–Ni minerals of the moncheite–merenskyite series. Other PGM found in these peridotites include sperrylite, minerals of the malanite– cuprorhodsite [(Cu,Fe)Rh2S4] solid solution series, braggite, unidentified Pt–Cu alloys, Pt– Fe alloys, and Au. Lorand et al. (2010) also identified smaller PGM during ablation runs of base-metal sulfides using LA-ICP-MS. In the Lherz peridotites, laurite occurs with Pt–Ir–Os alloys and members of the malanite– cuprorhodsite series in base-metal sulfide-poor, and base-metal sulfide-free harzburgites. The Pt- and Pd-rich bismuthotellurides are predominantly found in lherzolite and to a lesser extent in clinopyroxene-bearing harzburgite. Thus, the relative abundances of bismuthotellurides and base-metal sulfides increase significantly in the relatively fertile rocks (i.e., lherzolites and clinopyroxene-bearing harzburgite), whilst that of laurite drops. Lorand et al. (2010) interpreted the textural relationships between the PGM and base-metal sulfide in the different rock types as the result of mixing between refractory PGM (laurite/Pt–Ir–Os alloys) inherited from older refractory SCLM and late-stage metasomatic sulfides precipitated from a melt of tholeiitic composition. The formation of laurite and Pt–Ir–Os alloy is attributed to a series of partial melting events of the (~2 Ga) SCLM, that removed base-metal sulfide and left behind refractory PGM that were stable under S-undersaturated conditions. These PGM were subsequently mechanically collected by droplets of sulfide melt during refertilization reactions that converted the harzburgite protolith to lherzolite (see also Riches and Rogers, 2011). According to Lorand et al. (2010), the textural relationships and the positive correlation between base-metal sulfide and the abundance of Pt–Pd–Te–Bi minerals in the lherzolite indicate that these minerals are cogenetic and that their origin is related to the process of refertilization, which also added chalcogenes and semi-metals like Se, As, Sb, Te, and Bi.

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Marchesi et al. (2010) identified sulfide-hosted Pt-rich nuggets up to several µm across in peridotites of the Ronda massif (southern Spain). These Pt-rich minerals were inferred by inspection of time-resolved signals collected during laser ablation ICP-MS analyses.

SCLM peridotite xenoliths Mineralogical distribution and petrogenesis of PGM in mantle peridotite xenoliths. As with the Pt-rich nuggets identified by Marchesi et al. (2010) in the Ronda massif, discrete Ptrich minerals also occasionally occur within the base-metal sulfides found in mantle xenoliths. Examples include mantle xenoliths that have sampled the SCLM beneath the Massif Central in France (Lorand and Alard 2001), the South China Block in Taiwan (Wang et al. 2009), Central Iberia (Fig. 15; González-Jiménez et al. 2014c) and those associated with kimberlitic olivine from the Siberian lithospheric mantle (Griffin et al. 2002). The Pt-rich micronuggets that are found are generally considered to have formed at relatively low-temperature during sub-solidus decomposition of the base-metal sulfide. Lorand and Alard (2001) and Alard et al. (2002) noted that in some cases the PGM micronuggets associated with base-metal sulfides in mantle xenoliths are too small to be properly identified by LA-ICP-MS. Alard et al. (2011) used a Scanning Nuclear Microprobe to analyze sulfides in mantle xenolith material from Montferrier (southern France) and detected micronuggets of sperrylite. According to Alard et al. (2011), the sperrylite crystallized at temperatures below or near 900ºC (i.e., below the solidus of the sulfide) from volatile-rich fluids that refertilized the Montferrier mantle. The fact that Alard et al. (2011) were able to detect micronuggets of sperrylite included in the Monteferrier sulfides highlights the advantages of the Scanning Nuclear Microprobe relative to conventional techniques for identifying PGM. As reported by these authors, this instrument analyses a greater volume of the sulfide grain (imaged down to ∼ 70 μm depth) and can be very useful for identifying sulfide-hosted PGM. In contrast, LA-ICP-MS analyses are commonly carried out on areas of ∼ 50 mm in diameter (e.g., thick sections, polished blocks) and depths ranging between 40–60 μm, so may not provide sufficient coverage for PGM inclusion identification. The paucity of documentation of PGM nano-scale grains in the literature might be a problem of volume sampling using LA-ICP-

Figure 15. Images of Pt-alloy micronuggets in Fe-rich mss and magnetite (Mgt), interstitial to olivine in a mantle (lherzolite) xenolith from the Calatrava Volcanic Field, central Spain.

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MS, as well as a function of the limitations of conventional SEM microscopy. Therefore, the presence of PGM micronuggets in base-metal sulfides could be more frequent than reported, as also suggested by recent experimental results where PGM nanograins were crystallized from sulfide melts prior to larger base-metal sulfides (Helmy et al. 2013).

PGM IN CONCENTRICALLY ZONED URALIAN–ALASKAN–ALDAN-TYPE COMPLEXES Concentrically zoned, Uralian–Alaskan–Aldan-type complexes (CUAAC) are intrusions of relatively small size and elliptical shape, typically with a concentrically zoned structure where dunite is embedded in an envelope comprising pyroxenite ± wehrlite ± hornblendite ± magnetitite, passing outwards into gabbroic rocks (Supplementary Table 3). Minor associated lithologies include chromitite and magnetitite, and there are reported occurrences of gabbro/diorite, monzonite, monzodiorite, and hornbende–feldspar ± quartz ± biotite pegmatite. Concentrically zoned complexes occur within narrow structural belts of several hundreds of km length and develop at convergent margin settings (Uralian–Alaskan-type) or at the periphery of stable cratons (Aldan-type). Intrusions hosted in orogenic belts in Russia are referred to as Uraliantype (cf. Noble and Taylor 1960; Taylor 1967; Foley et al. 1997; Garuti et al. 1997, 2002, 2003; Malitch and Thalhammer 2002; Augé et al. 2005a), whereas those located in similar tectonic settings in southeastern Alaska (and more generally the North American Cordillera) are termed Alaskan-type (Findlay 1969; Himmelberg et al. 1986; Nixon et al. 1990; Patton et al. 1994; Himmelberg and Loney 1995; Fedortchouk et al. 2010). Examples of CUAAC have also been reported from all over the world (e.g., Australia, Colombia, Ecuador, Egypt, Madagascar, and Papua New Guinea), and can collectively be termed Uralian–Alaskan-type complexes (see Johan 2002 and Supplementary Table 3). Zoned Aldan-type complexes are located in the Aldan Province (southern Siberian Craton) and from the Koryak–Kamchatka belt and the Aluchin Horst (Chukota Pensinsula in eastern Russia; Podlipskiy et al. 1999; Gornostayev et al. 1999, 2000; Malitch and Kostoyanov 1999; Malitch et al. 2011a,b,c; Scheka et al. 2004; Tolstykh et al. 2004; Nazimova et al. 2011). Most authors agree that CUAAC from both orogenic belts and stable cratonic platforms crystallize from mantle-derived melts. Geochemical and petrological studies of zoned complexes indicate parental melts characterized by low aSiO2, giving rise to intrusions rich in clinopyroxene with little or no orthopyroxene. Most authors suggest that the parental melts of the CUAAC were generated in supra-subduction zones, or within marginal continental crust. From the standpoint of PGE geochemistry, an important feature of CUAAC rocks is positive anomalies of Pt and Ir with negative anomalies of Ru and/or Pd producing ‘M’ shaped patterns on standard chondrite-normalized PGE diagrams (Fig. 16). This characteristic distribution of the PGE distinguishes CUAAC from ophiolites and layered mafic–ultramafic intrusions. Nevertheless, the PGE patterns of CUAAC are variable from one locality to another. The PGE mineralization in CUAAC dominantly occurs as lodes, associated with chromitite, dunite– pyroxenite and/or concentrations of base-metal sulfides. Another distinctive characteristic is that streams draining CUAAC form placers of alluvial PGM that may attain local economic importance (see Appendix). Such occurrences constituted the main source of Pt up until the early twentieth century, when the PGE deposits in the Bushveld Complex were discovered. Broad overviews of PGE-mineralization in selected examples of CUAAC are provided by Yuan (2001), Johan (2002) and Economou-Eliopoulos (2010). In general, the bulk of the literature on PGE in CUAAC does not provide detailed information on the textural and compositional characteristics of PGM; many authors simply refer to the relative abundances of the PGM in terms of high- and low-temperature assemblages. Supplementary Table 3 summarizes CUAAC PGM information, as reported in the literature.

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Figure 16. CI-chondrite (Naldrett and Duke 1980) normalized patterns of different rock types from CUAAC. Data plotted in the top three graphs correspond to rocks from the Nizhny Tagil zoned complex in the Urals (Augé et al. 2005a; Malitch et al. 2011c). Average PGE concentration data for chromitites and dunites from other CUAAC are presented as follows: the Kytlym, Uktus Flekistovsky and Kyrlyn complexes of the Urals (Garuti et al. 1997, 2003; Zaccarini et al. 2011), Serpentinite Hill in Tasmania (Brown et al. 1998), Tulameen in British Columbia, Canada (Nixon et al. 1990) and Viravira and Condoto in Colombia (Tistl 1994). Data for sulfide-rich rocks of the Gabbro Akarem complex are from Helmy and Mogessie (2001).

PGM in dunite, pyroxenite and gabbro PGM in peridotites and mafic rocks of CUAAC. Accessory PGM are found dispersed in dunites (Alto Condoto, Colombia; Kondyor and Galmoenan, Russia), or other mafic lithologies, including pyroxenites and gabbros (Galmoenan, Russia), and magnetite-rich clinopyroxenites (Kachkanar, Russia; Fifield, Australia). The PGM identified by Tistl (1994) in the dunites of the zoned complex of Alto Condoto are mainly Pt–Fe alloys (possibly isoferroplatinum) containing lamellae of native Os (Supplementary Table 3). Tistl (1994) reported a second type of primary PGM mineralization (Viravira-type) within fragments of largely serpentinized, orthopyroxenebearing peridotites (dunite, harzburgite, lherzolite) hosted in high-Mg basalts. The PGM associated with dunite in the Kondyor massif are Pt–Fe alloys located in the interstices of chromite, olivine, and pyroxene aggregates (Rudashevsky et al. 1982; Ukhanov et al. 1997; Cabri et al. 1998). Primary accessory PGM in dunites, pyroxenites and gabbros at Galmoenan are isoferroplatinum, native Os, native Ir, and sulfides of the laurite–erlichmanite solid solution series (Supplementary Table 3). Low-temperature alteration of this PGM assemblage produced a secondary paragenesis dominated by tetraferroplatinum, tulameenite, hongshiite, unidentified Pt–Cu, sperrylite, and sulfarsenides of the irarsite–hollingworthite solid solution series

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(Supplementary Table 3). The main PGM phase in the magnetite-bearing clinopyroxenite of the Kachkanar (Gusevogorsk deposit) is the Pt–Fe alloy isoferroplatinum (Razin and Yurkina 1971; Fominmykh et al. 1974; Bezigov et al. 1975) containing Os–Ir alloys, ‘drop-like’ inclusions of cooperite, euhedral crystals of laurite, and erlichmanite and Ir- and Rh-rich sulfides including bowieite and kashinite [(Ir,Rh)2S3]. Other PGM identified in the Gusevogorsk deposit (associated with base-metal sulfides dispersed in the peridotite) include unidentified Pd–Pt–Fe alloys, Pd– Cu, Hg3Pd2, braggite, vysotskite, atheneite, mertieite-I [Pd11(Sb,As)4] and Pd-rich tellurides (Supplementary Table 3). Johan et al. (1989b) and Slansky et al. (1991) reported primary PGE mineralization in magnetite-bearing pyroxenites (the ‘P-unit’ of Suppel and Barron 1986) of the Fifield complexes, consisting of isoferroplatinum, tetraferroplatinum, erlichmanite, cooperite, and cuprorhodsite–malanite. The latter assemblage is locally replaced by secondary geversite, sperrylite, and stumpflite [Pt(Sb,Bi)]. Crystallization of primary PGM in peridotites and mafic rocks of CUAAC. Tistl (1994) suggested that PGM crystallization in the Alto Condoto Complex was coeval with, but independent of, chromite crystallization. He proposed that cogenetic assemblages of PGM and chromite result from mechanical trapping of PGM grains by contemporaneous crystallizing chromite. In this model, the most important factor governing PGM precipitation, particularly Pt–Fe alloys, is the crystallization of olivine from the silicate melt. This follows from the experimental works of Amossé et al. (2000), arguing that the crystallization of olivine leads to an increase of melt fO2, which lowers the solubility of Pt and Ir, but inhibits formation of PGM formed from Ru, Rh, and Pd since these PGE cannot form stable minerals at these P–T conditions. Tistl (1994) suggested that the presence of common crystallographic growth defects in the Alto Condoto PGM points to rapid growth of these minerals, potentially driven by volatile exsolution from the melt (causing a decrease in the crystallization temperature), or by changes in fO2 and/or fS2. Rapid tectonic movements, causing a decrease in pressure and superheating of the melt, could promote such changes. Thus, formation of the Alto Condoto dunite was initially characterized by early crystallization of Pt–Fe alloys, chromite and olivine. Continued crystal fractionation of chromite and olivine (principally controlled by fS2, according to Tistl 1994) drove magma evolution and the precipitation of PGE sulfides. In the model suggested for the Alto Condoto Complex by Tistl (1994), the singular Pt enrichment of the ultramafic and mafic rocks is thus not associated with especially high Pt in the parental melts or the mantle source. Johan et al. (1989b) suggested that the intergrowth of Pt–Fe alloys + erlichmanite, isoferroplatinum + cooperite, or erlichmanite + cooperite associated with the magnetitebearing pyroxenites at Fifield reflects a progressive increase of fS2 during different stages of cooling of the silicate melt. They proposed that erlichmanite and the Pt–Fe alloys would crystallize first at relatively high temperature and lowest fS2; down temperature increases in fS2 would give rise to a suite of PGE sulfides such as cooperite and cuprorhodsite–malanite. They suggested that the presence of the equilibrium OsS2 + PtS2 association implies a maximum temperature of about 860 ºC for the crystallization of PGM in the ore-forming system. In their model, crystallization of the PGM is probably related to the appearance of a reducing fluid during the final stages of evolution of the so-called P-units, which develop in the most oxidized pyroxenites of the Fifield complexes.

Chromitite-hosted PGM in CUAAC Occurrence and frequency of PGM in chromitites of CUAAC. The PGM associated with CUAAC-hosted chromitites have mainly been documented from several massifs of the Russian Platinum Belts (Fig. 17; Supplementary Table 3; Ural Mountains; Nizhny Tagil, Uktus, Volvosky, Kytlym, Kachkanar, and Koryak–Kamchatka [Galmoenan]), the Tulameen massif in British Columbia (Canada), Yubdo (Ethiopia), and Kondyor (Russian Aldan Shield). The chromitites associated with these massifs are usually small bodies (~tens of cm long and several

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cm thick) hosted in dunite. However, Augé et al. (2005a) observed larger (100 × 5 m) schlieren of disseminated chromite and similarly-sized bodies of chromitite within the ultramafic unit of the Nizhny Tagil Complex. These may be brecciated, net-textured or massive. Brecciated chromitites contain autoliths of the host dunite, suggesting that chromitite is younger (Augé et al. 2005a). Zaccarini et al. (2011) reported a composite chromitite/amphibolite clinopyroxenite vein with anomalously high PGM concentrations at Butyrin in the Kytlym massif. Chromititehosted PGM in CUAAC are usually micrometric grains (< 50 µm) positioned in one of three manners: (1) inclusions in chromite crystals, (2) at chromite–silicate grain boundaries and (3) more rarely, in the silicate matrix. A characteristic feature of CUAAC chromitites is their positive anomalies of Pt plus Ir and negative Ru and Pd anomalies on chondrite-normalized PGE diagrams, yielding diagnostic ‘M’shaped patterns (Fig. 16), as noted above. Nevertheless, the PGE patterns of CUAAC exhibit additional variability depending on locality. For example, Garuti et al. (2003) identified three distinctive PGE patterns in chromitites of the Utkus Complex, which they linked to variation in the chemical composition of chromite in the chromitites. They devised a classification scheme, in which their type I chromitites are characterized by PGE normalized-chondrite patterns with a negative slope due to enrichment of the IPGE over the PPGE. Chromite in this case is magnesiochromite with low Fe3+/(Fe3 + Fe2+) = 0.23–0.35. The principal PGM include laurite, kashinite, and cuproiridsite (CuIr2S4); secondary phases are irarsite and tolovkite (IrSbS). Their type II chromitites exhibit a PGE distribution pattern characterized by an M-like shape with marked positive anomalies of Pt and Ir; the chromite here is magnesiochromite with higher Fe3+/(Fe3++Fe2+) ratios (0.40–0.44). Type II chromitites contain abundant primary isoferroplatinum associated with native Os, erlichmanite, cuproiridsite, and cuprorhodsite, coexisting with free grains of malanite and cooperite. Isoferroplatinum is commonly replaced by tulameenite, whereas the Ir- and Rh-bearing PGM are transformed to unnamed Rh4S3 and Ir–Fe, accompanied by potarite. The third, type III chromitites contain Fe-rich chromites (Fe3+/(Fe3++Fe2+) = 0.59) and exhibit PGE patterns with an M-like shape similar to that of type II chromitites, but with higher Pt and Rh abundances. The PGM assemblage in type III chromitites consists of irarsite, tulameenite, RhSbS, PtPdCu, and PdCu alloys. The mineralogy of PGM in chromitites of the CUAAC is monotonous, typified by a predominance of Pt-rich PGM and a paucity of Os–Ir–Ru-rich PGM (Figs. 17 and 18). Alloys of the Pt–Fe–Ni–Cu system are by the far the most abundant. Isoferroplatinum is the most common alloy followed, in order of abundance, by tetraferroplatinum, tulameenite, and more rarely Fe-bearing native Pt, Fe + Ni-bearing native Pt and hongshiite, as well other unidentified alloys with variable proportions of Pt/(Fe+Ni+Cu) (Fig. 18). The Pt–Fe alloys often host inclusions of native Os or Ir; the latter can also be found as isolated inclusions in chromite. The coexistence of these alloys along with sulfides of the laurite–erlichmanite solid solution series has been described in chromitites of most CUAAC. Other rare PGM include sulfides such as cooperite, cuproiridsite, cuprorhodsite, and kashinite, and Ir–Rh– Pt sulfarsenides such as irarsite, hollingworthite and platarsite, sperrylite, geversite, and RhNiAs (possibly zaccariniite; Vymazalová et al. 2012). Secondary low-T assemblages of PGM are generally formed by metasomatic replacement of pre-existing PGM and localized remobilization of the PGE during serpentinization and regional metamorphism (e.g., Nixon et al. 1990). The Pt–Fe alloys are commonly replaced by rims of Cu-rich alloys like tulameenite and/or As- and Sb-rich minerals such as sulfarsenides of Ir–Pt–Rh (irarsite–platarsite–hollingworthite), arsenides (sperrylite) or antimonides (geversite and genkinite), and more rarely by PGE antimonides and reported occurrences of Pt-oxides. Other relatively rare secondary minerals identified in chromitites of the Galmoenan massif in the Koryak–Kamchatka Platinum belt include bismuthides like froodite (PdBi2) or sobolevskite (PdBi) and stumpflite (Supplementary Table 3).

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Figure 17. Images of Os-Ir alloy grains (a) in situ and (b) liberated by mechanical processing from chromitites of the Guli massif (adapted from Malitch et al. (2011b) Redox conditions of formation of osmiumrich minerals from the Guli massif, Russia. Geochemistry International, Vol. 49, p. 726–730, with the permission of the author and publisher).

Figure 18. Compositional variation of Pt-alloys hosted in chromitites from the concentrically zoned complexes of the Urals (Kytlym and Uktus). These plots are adapted, replotted and redrawn from Garuti et al. (2002).

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Crystallization of primary PGM during chromitite formation in CUAAC. As noted above, CUAAC chromitites may contain a primary PGM assemblage formed at high-temperatures and/ or a secondary one developed at low-temperatures during late-stage magmatic or post-magmatic events. The high-temperature PGM assemblage is usually found as inclusions in chromite. However, Augé et al. (2005a) also reported primary PGM attached to the edges of chromite grains or forming the interstitial groundmass to these grains in chromitites of the Nizhny Tagil Complex. Nixon et al. (1990) observed a similar PGM-chromite relationship in chromitites of the Tulameen Complex, suggesting that it may be a common characteristic of CUAAC chromitites in general. Malitch and Thalhammer (2002) reported the first (and only, to our knowledge) Os-isotope dataset from Os-rich minerals in CUAAC-hosted chromitite. Two grains of native Os from the chromitites of the Jurassic Kondyor massif (analyzed by N-TIMS) yielded 187Os/188Os values of 0.1248 and 0.1252. The authors interpreted these relatively low 187Os/188Os compositions as indicative of a mantle source for the PGE, suggesting the PGM have a magmatic origin. Nixon et al. (1990) also suggested that euhedral Pt–Fe alloys locked in chromite grains of the Tulameen chromitites were high-temperature segregations that formed in an early phase of crystallization from primitive mantle-derived silicate melts. Using the thermodynamic and experimental data obtained by Hill and Roeder (1974) and Amossé et al. (1990), Nixon et al. (1990) estimated an increase in fO2 associated with the crystallization of chromite would significantly drop the solubility of Pt in the surrounding melt, favoring the precipitation of Pt–Fe alloys. The latter authors estimated the co-precipitation of Pt–Fe alloys and chromite occurred within the range ~1300–1200 °C and at –6 log fO2. The latter model contrasts with that proposed by Garuti et al. (2003) who explained the abundance of Pt–Fe alloys in Uktus Complex chromitites (Fig. 18) as being related to changes in the activity of FeO in the melt, rather than fO2. Garuti and co-workers suggested, based on experimental results on the solubility of Pt in basaltic melts (e.g., Borisov and Palme 1997, 2000; Amossé et al. 1990), that the precipitation of Pt–Fe alloys may take place independently of the true variation in fO2 and may reflect an increase in FeO activity during magma fractionation. Garuti et al. (2003) suggested that the low silica activity of the melts that produce the CUAAC could favor the crystallization of olivine, resulting in an increase in the Fe2O3 and Fe3O4 activity of the melt. They concluded that under these conditions, fractionation produces increasing amounts of Fe2+ and Fe3+ in the melt that may be incorporated into the crystallizing chromite; as a result the chromite would show an increasing oxidation ratio (Fe3+/[Fe3++Fe2+]) as is observed in the Uktus chromitites. The observations of Garuti et al. (2003) are noteworthy since the effect on chromite composition of changing FeO activity in the melt is similar to that of increasing the melt fO2. Co-precipitation of Fe-rich chromite and Pt–Fe alloys in CUAAC chromitites may thus be a result of the critical undersaturation in SiO2 of the primitive melt. Another interesting observation made by Garuti et al. (2003) on the Uktus chromitites is that the precipitation of Pt– Fe alloys appears to have occurred only during the crystallization of chromite under conditions of high fO2, as evidenced by chromite with Fe3+/(Fe3++Fe2+) > 0.4. In contrast, these authors did not observe crystallization of Pt-rich phases in chromite with a low Fe-oxidation ratio (i.e., < 0.3). Chromites with low ratios would be expected to crystallize in the early stages of magmatic fractionation where the melt is characterized by lower FeO activity, so the latter scenario was interpreted as meaning that the Pt remained in the silicate liquid and other Ru–Os–Ir minerals like laurite or Ir-sulfides were the only PGM that crystallized. Augé et al. (2005a) attributed the abundance of Pt-rich PGM in the chromitites of the Nizhny Tagil and Kachkanar Complexes to mechanical processes that concentrated both chromite and PGM. On the basis of their observations, they concluded that there is no evidence to support a Pt-rich melt/fluid as being important in the formation of PGM in CUAAC. They suggested that the observed chromite breccias and networks of massive chromite in the Nizhny Tagil and Kachkanar CUAAC are indicative of the formation of chromitites via dynamic accumulation

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of chromite in cavities in magma conduits. According to Augé et al. (2005a), the PGM would be incorporated in crystallizing chromite crystals or remain attached to them due to their strong affinity for the oxide rather than the silicate melt. The effectiveness of chromite to accumulate a higher concentration of PGM would be determined by the ratio of the mass of silicate melt to the mass of chromite in contact with the ascending melt (i.e., the CR-factor, which is analogous to the sulfide mineralization R-factor, defined by Campbell and Naldrett, 1979). Where the CRfactor is high (i.e., a relatively low proportion of chromite in contact with a large volume of melt crystallizing the PGM), the efficiency of the mechanical collection of the PGM is at a maximum and the concentration of PGE (or concentration of PGM grains) will be high. In contrast, a low CR-factor produces PGE- or PGM-poor chromitites, as seen in some ophiolites. The Augé et al. (2005a) model requires the existence of melts that already contain PGM, i.e., it does not explain the crystallization of PGM; it invokes a previous mechanism to lower the solubility of the PGE in the melt and crystallize PGM that are subsequently ‘physically’ accumulated. Okrugin (2011) proposed an alternative model for the origin of PGM in CUAAC, including chromitite-hosted PGM. His model is based on studies of PGM parageneses from both chromitites and placer deposits, associated with the CUAAC of the Aldan Shield. Okrugin suggested that fractional crystallization of olivine from a picritic melt would result in the formation of a residual melt enriched in Cr, the PGE, and other metals, as well as easily fusible and volatile elements. He suggested that over the course of slow fractional crystallization, the latent heat of olivine crystallization could drive the residual melt to an immiscibility field through a superliquidus path. On reaching the immiscibility field, this residual melt would separate into two immiscible melts; a lighter SiO2-rich one and a denser Cr-rich oxide one. In a fluid-saturated parental melt, a Cr-rich oxide melt would be separated from the silicate melt based on the density contrast; upon cooling the crystallization of such a melt would lead to the formation of the chromitites. Okrugin (2011) proposed that the high PGE content of chromitites (relative to the host dunites) could be explained by the initial accumulation of PGE in the dense Cr-rich oxide melt. The preferential partitioning of the PGE into this Cr-rich oxide melt is the result of the chromite ionic-structure vs. the ionic-molecular nature of the silicate melt. Metallic nanoclusters of PGM accumulate in the Cr-rich oxide melt, stabilized by outside ‘envelopes’ of ligands of semi-metals (e.g., S, As, Sb, Te). According to Okrugin (2011), the coalescence and subsequent crystallization of such clusters leads to the formation of various PGM assemblages consisting mainly of alloys, often intimately intergrown with sulfides, arsenides, and other PGE and ligand compounds. The crystallization of the different PGM assemblages observed in the chromitites results from the progressive separation of a series of immiscible melts enriched in PGE after chromite crystallization. This is because crystallization of Cr-rich oxide melts leads to the concentration of PGE in the residual melt, since PGE are less soluble in chromite than in the melt. The characteristic Pt-enrichment observed in the chromitites of CUAAC is thus assumed to reflect the initial IPGE/PPGE ratio of the parental melt.

PGM and sulfide mineralization in CUAAC The association of PGM and sulfide mineralization in CUAAC. The association of disseminated and massive base-metal sulfides with PGM is an unusual feature in CUAAC. This association has only been described in a couple of zoned complexes in North America (Salt Chuck in Alaska and Turnagain in British Columbia), the Central Urals (Vokosvsky Complex) and from the Eastern Desert in Egypt (i.e., Gabbro Akarem and the Genina Gharbia mafic–ultramafic intrusions) (Supplementary Table 3). At the Salt Chuck complex, the PGM are associated with 15 vol.% bornite and chalcopyrite in biotite-bearing magnetite clinopyroxenite and gabbro (Loney and Himmelberg 1992), which are confined to fault zones. This sulfide mineralization was mined intermittently from 1905 to 1941 and produced about 300,000 metric tonnes of sulfide ore, estimated to have averaged 0.95 wt.% Cu, 19.91 g Au, 7.04 g Ag, and 26.1 g Pd, per tonne (Gault 1945; Holt et al. 1948). Watkinson and Melling

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(1989) identified isolated grains of PGM associated with the sulfides, consisting mainly of kotulskite (PdTe) intergrown with hessite (Ag2Te). Other assemblages consisting of temagamite (Pd3HgTe3) and isolated grains of Au-rich sperrylite + kotulskite + Pd6AsSb, temagamite + Au, and kotulskite + temagamite + merenskyite were observed. The main occurrences of massive to semi-massive Fe–Ni–Cu sulfide ores in the Turnagain complex are hosted by wehrlite and clinopyroxenite and rarely in serpentinized dunite (Nixon 1997). The base-metal sulfides are envisaged to have formed from immiscible sulfide melt that segregated from silicate magma (Clark 1975). Jackson-Brown et al. (2014) identified clinopyroxenites and hornblendites with up to 5 vol.% sulfides including chalcopyrite and pyrrhotite with minor pyrite and pentlandite, in addition to a variety of arsenides, As–Sb sulfides and PGM. The PGM include sperrylite and sudburyite [(Pd,Ni)Sb], with minor Pd-rich melonite (NiTe2), hongshiite, testibiopalladinite [PdTe(Sb,Te)], and genkinite. The PGM form small (< 40 µm across) inclusions in chalcopyrite, pyrrhotite, pentlandite, cobaltite (CoAsS), and silicates. Sperrylite and sudburyite also form veins and rims on the edges of base-metal sulfides. Zaccarini et al. (2004a) observed abundant grains of vysotskite associated with patches (50–300 µm) of chalcopyrite (± bornite, ± carrollite; CuCo2S4) rimmed by magnetite and disseminated in serpentinized olivine- and apatite-rich rocks, from the zoned complex of Volkovsky in the Central Urals. Vysotskite has been replaced along its grain boundaries by stillwaterite (Pd8As3) and unidentified Pd–As–Te (törnroosite? [Pd11As2Te2]) PGM. Zaccarini et al. (2004a) also reported large patches of kotulskite (~20 µm) intergrown with electrum and an unidentified mineral with a composition close to Pd3(As,Te). At the Genina Gharbia mafic–ultramafic intrusion, Helmy (2004) described michenerite and melonite–merenskyite, associated with hessite, altaite (PbTe), tsumoite (BiTe), sylvanite [(Ag,Au)Te2], and native Te, located mainly at sulfide–silicate grain contacts and as inclusions in altered silicates. Palladium-bearing bismuthian melonite is the most abundant PGM and forms: (1) isolated euhedral (< 30 µm) inclusions in pyrrhotite and pentlandite; (2) polyphase aggregates (30–50 µm) with hessite, michenerite, altaite, and tsumoite at the contact between sulfides and silicates; (3) grains of variable size (15–70 µm) associated with altaite and tsumoite embedded in secondary silicates (mainly quartz and epidote). Helmy (2004) described melonite associated with native Te within cracks in violarite (Fe2+Ni23+S4) and forming composite grains with hessite and altaite at the contacts between sulfides and silicates. Merenskyite was observed at the contacts between pyrrhotite and epidote and, more rarely, in hessite associated with fractured pyrrhotite. The PGM associated with Ni–Cu–Fe sulfides from Gabbro Akarem are native Pd and bismuthotellurides (Supplementary Table 3). Helmy and Mogessie (2001) described ~85 grains of merenskyite enclosed in pyrrhotite together with several grains of michenerite at chalcopyrite–silicate (serpentine and chlorite) contacts. Palladium-bearing bismuthian melonite forms small subhedral crystals (10–20 µm) associated with hessite, often along cracks within sulfides and at the contacts between the sulfides and silicates. Petrogenesis of the PGM-sulfide association in CUAAC. Gault (1945) interpreted the assemblage of PGM + sulfide at Salt Chuck (Alaska, USA) as the product of infiltrating ‘ore-bearing’ fluids along the contacts between fractured magnetite-rich clinopyroxenite and gabbro. He suggested the preferential concentration of chalcopyrite in the main faults of the massif was due to the infiltration of Cu-bearing fluids along these structures. Mertie (1969) suggested that the ubiquitous occurrence of sulfides in microfractures was evidence for the precipitation of sulfides from circulating hydrothermal solutions. The disseminated character of the sulfides and PGM, together with the predominantly igneous textures exhibited by the host rocks, suggests crystallization of these minerals from late-stage magmatic fluids during an intrusive event (Loney and Himmelberg 1992). The latter authors suggested that the

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veins and fractures observed by Gault (1945) and Mertie (1969) were formed during this late magmatic event. Loney and Himmelberg (1992) argued for a two-stage process, involving the initial (magmatic-stage) development of disseminated sulfides and PGE, largely in the gabbroic part of the zoned complex. A subsequent magmatic-hydrothermal stage remobilized and concentrated the PGE in veins and fractures in the magnetite-bearing clinopyroxenite primarily close to the contact between the magnetite-bearing clinopyroxenite and gabbro. Watkinson and Melling (1992) suggested that the formation of PGM was associated with the post-magmatic infiltration of a Cl-bearing hydrothermal fluid enriched in PGE and Cu, which replaced the magmatic sulfides. Nixon (1997) interpreted the formation of base-metal sulfides in the Turnagain intrusion as being products related to the separation of immiscible sulfide melt from a silicate melt. His interpretation is based on the fact that the sulfides are interstitial to silicates and show bleb-like textures in disseminated zones, or coalesce to form continuous networks enclosing cumulus silicate crystals. The net-textured sulfides locally occlude silicates altogether to form massive accumulations. Jackson-Brown et al. (2014) noted that the abundance of base-metal sulfides in the Turnagain intrusion differs from most Alaskan-type intrusions and Scheel (2007) suggested that the relative abundances of sulfide here (relative to other CUAAC) was a result of assimilation of crustal sulfur by the parental melt. Jackson-Brown et al. (2014) suggested that formation of the PGM was associated with the intrusion of a Cu–Pt–Pd-bearing melt phase that crystallized separately from the earlier (Ni–Co enriched) main dunite–wehrlite units. Zaccarini et al. (2004a) recognized three stages of PGM formation in the Volkovsky Complex (Central Urals, Russia) that they related to chemical changes during fractional crystallization and cooling of the ore system. According to these authors, high initial fS2 in the melt caused the crystallization of vysotskite and chalcopyrite. This was followed by continued cooling that promoted the exsolution of kotulskite from vysotskite and electrum from both vysotskite and chalcopyrite. According to Zaccarini et al. (2004a), the occurrence of arsenotellurides and kotulskite in the outer rims of the sulfide blebs indicates that the melt became enriched in Te and As after the first stage. In the final stage, Te activity in the melt decreased significantly and enrichment in As led to the formation of stillwaterite that overgrows arsenotellurides. The observation that stillwaterite often forms intergrowths with magnetite (itself associated with secondary silicates) led Zaccarini et al. (2004a) to suggest that the final stage was characterized by high aAs, coinciding with the serpentinization of the silicates under relatively oxidizing conditions. Helmy (2004) also proposed a three-stage model to account for the petrogenesis of PGM in the Genina Gharbia intrusion (Eastern Desert, Egypt). In the first stage, sulfide saturation was achieved in the silicate melt due to fractional crystallization. Base-metal sulfides were precipitated in an interstitial network around early-formed olivine and pyroxene in harzburgite and lherzolite. All of the PGE were incorporated into sulfides and sulfarsenides (mainly pentlandite and cobaltite–gersdorffite). At an advanced stage of magma fractionation, additional sulfur and other semi-metals were added to the melt from the surrounding metasediments. The base and noble metals were subsequently separated and concentrated in a highly evolved H2Orich fluid that formed in equilibrium with the parental melt to the gabbros. Sulfides, PGM, and other tellurides then crystallized from this metal-rich fluid at lower temperature. The postmagmatic stage is characterized by localized ductile deformation, which remobilized metals and concentrated them along shear zones. Helmy and Mogessie (2001) estimated a crystallization temperature for the lherzolites at Gabbro Akarem between 912–888 °C, using two-pyroxene geothermometry. They suggested that this temperature range is the upper limit for the segregation of immiscible sulfide liquid following crystallization of olivine and orthopyroxene. The identification of troilite (FeS) + smythite [(Fe,Ni)9S11 or (Fe,Ni)13S16] intergrowths at Gabbro Akarem (Sideek

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and El Goresy 1996) led Helmy and Mogessie (2001) to propose extensive equilibration from magmatic temperatures down to low temperatures (∼ 75 °C). The latter authors suggested that the initial stages of cooling of the immiscible sulfide melt were dominated by the early segregation of merenskyite now found as inclusions in pyrrhotite, whereas the final stages were dominated by the crystallization of michenerite identified at the contacts between chalcopyrite and silicates. Helmy and Mogessie (2001) also suggested that the infiltration of low-temperature hydrothermal solutions could be responsible for the mobilization of Pd and Au that contributed to the formation of merenskyite and electrum in serpentine veins that crosscut the olivine grains.

PGM IN Ni-SULFIDE DEPOSITS Komatiite-associations Komatiites are ultramafic volcanic rocks that crystallize from melts with MgO contents > 18 wt.% that are considered to be essentially free of volatiles (Arndt 2008). In general, ultramafic rocks with 10–15 wt.% MgO that are demonstrably derived from anhydrous komatiitic parental melts are referred to as komatiitic basalts. Although most komatiites are restricted to Archean and early Proterozoic greenstone belts, they have rarely been reported from rocks of the Phanerozoic Eon. Relatively young komatiites include those found on Gorgona Island in Colombia (Gansser et al. 1979; Echeverria 1980; Arndt 1994; Arndt et al. 1997) and the Triassic lavas described from the Otrhys ophiolite complex in Greece (Cameron et al. 1979; Cameron and Nisbet 1982; Paraskevoupoulos and Economou 1986; Tsikouras et al. 2008). Komatiites are typically considered to be S-undersaturated at the comparatively high temperatures of their eruption (1450–1600 °C; Lesher et al. 1984; Fiorentini et al. 2010). However, the assimilation of crustal material can substantially change their nature (e.g., S-content and degree of polymerization) and promote sulfur saturation, potentially leading to the formation of magmatic Ni-sulfide deposits (Lesher et al. 1984; Lesher 1989; Barnes 2006). Komatiite-hosted ores have been classified using a variety of criteria (see Lesher and Keays 2002 for a review). In the present contribution, we consider two principal groups for the sake of simplicity: (1) sulfide-rich (deposits in lava channels, i.e., type 1 of Lesher and Keays 2002) and (2) sulfide-poor (stratiform layers in differentiated sills or lava lakes, i.e., type 3 of Lesher and Keays 2002), in a similar manner to Fiorentini et al. (2004, 2007) and Locmelis et al. (2009, 2011, 2013). Mineralogy and petrogenesis of PGM in sulfide-poor komatiites. PGE-rich, sulfide-poor mineralization in komatiites has been described in differentiated sills and/or lava lakes formed by the strong in situ fractionation of ultramafic melts. These sills and lava lakes are generally associated with complementary cumulates. This kind of PGE mineralization, ‘type 3’ in the classification of Lesher and Keays (2002), is generally interpreted as the result of precipitation from sulfide melts that were segregated in cotectic proportions with olivine when an initially S-undersaturated komatiite melt underwent extensive closed-system fractionation. Reported examples of PGM associated with this type of mineralization include the Boston Creek Sill (Abitibi greenstone belt in Ontario, Canda; Stone et al. 1992; Fiorentini et al. 2004), the atypical stratiform mineralization of the Wiluna cumulate body (east Yilgarn craton; Fiorentini et al. 2004, 2007) and Mount Clifford (Locmelis et al. 2009; western Australia). The PGM identified in the Boston Creek Sill are merenskyite, bismuthian kotulskite, merteite I, unidentified Pd–Ag sulfide, sperrylite, and an unidentified Rh–Pt sulfarsenide. According to Stone et al. (1992), the PGM occur as grains < 15 µm in diameter associated with Cu–Fe sulfides, native Au, electrum, a suite of Ag minerals, and secondary silicates derived from the alteration of Fe–Ti oxides in pyroxenite. Stone et al. (1992) suggested that the mineralogy, textures, and spatial associations of the PGM indicate the importance of secondary processes in their formation. The anhedral, sutured, and elongate shapes of the PGM at grain boundaries with secondary silicates were interpreted as reflecting remobilization of the PGE during growth of the secondary silicates.

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Stone et al. (1992) proposed that Pd and Pt were remobilized and concentrated in the pyroxenites by hydrothermal fluids during the metasomatic transformation of the primary cumulate, during either cooling of the flow or later greenschist-facies metamorphism. Fiorentini et al. (2004) suggested that inclusions of Ir–Os (± Pt) were present in chromitites hosted in komatiitic basalt flows from the greenstone belts of Abitibi (Theo’s, Fred’s, and Boston Creek flows) and Agnew-Wiluna (Wiluna). Though not observed directly, the presence of these inclusions was inferred from peaks for Ir, Pt, and Os in time-integrated depth profiles generated during laser ablation ICP-MS analyses of chromite grains. Fiorentini et al. (2007) subsequently re-examined a suite of chromite-free pyroxenites and melagabbros from the upper mafic section overlying the peridotitic cumulates at Wiluna, using an SEM equipped with an automated image analyzer optimized for PGM detection. In this study, they identified > 70 PGM grains, none of which formed part of the same paragenesis as sulfide, i.e., the PGM were only found in sulfidepoor/free intervals. The most frequent PGM they identified were small (1–3 µm) alloy grains, including 24 grains of unidentified Pt–Pd–Cu alloys, and 10 grains of isoferroplatinum, fully embedded in clinopyroxene crystals. In addition, they identified 31 grains of sperrylite in zones of oikocrystic pyroxene in pyroxenite and 12 grains of unidentified Pd-bearing bismutotelluride in small veinlets at the edges of clinopyroxene crystals in melagabbro. According to Fiorentini et al. (2007), the fact that all of the PGM in oikocrystic pyroxenites are bleb-shaped and occur as clusters of very closely spaced grains hosted within clinopyroxene indicates that the crystallization of the PGM took place coevally with the onset of pyroxene crystallization. Because the amount of sulfide at Wiluna is very small, they hypothesized that PGE saturation was driven purely by fractional crystallization of silicate and oxides in a S-undersaturated melt. They suggested that clinopyroxene catalyzed the nucleation of PGM, similar to the role of chromite in chromite-saturated melts. Locmelis et al. (2009) identified 66 grains of PGM in the komatiitic (dunite) cumulates at Mount Clifford, approximately 150 km south of Agnew-Wiluna (Wiluna). According to Locmelis et al. (2009), all of the PGM are hosted in serpentine and rarely intergrown with sulfides (pentlandite or millerite), magnetite, and nickel arsenides that formed during the alteration. The PGM are small (< 5 µm) and occur as clusters with very heterogeneous internal textures. The minute size of the grains coupled with their chemical heterogeneity makes their accurate identification challenging. Locmelis and co-workers classified these PGM into three compositional types: (1) Pd-antimonides (Pt-free and Pt-bearing), (2) Pt-dominated PGM and (3) Pt-bearing Ni-antimonides. According to Locmelis et al. (2009), the association of a PGErich zone with a reversal in the Mg/Fe ratio of the host cumulates, as well as offsets in peak concentrations of Pt, Pd, and Cu-S peaks for whole-rock compositions indicates a primary magmatic origin for the PGM mineralization, rather than a secondary origin associated with hydrothermal alteration. They suggest a model in which the formation of the different PGM was the result of a combination of processes, including early sulfide segregation at low R-factor values producing Ni- and IPGE-rich olivine, partial re-dissolution and removal of most of this sulfide component during a major replenishment event and subsequent overprinting by a mobile Pt–Pd–Cu-rich fluid derived from highly differentiated trapped liquid. Mineralogy and petrogenesis of PGM in sulfide-rich komatiites. Proterozoic komatiite– basalt sulfide-rich PGE deposits, such as those of the Raglan Ni–Cu–(PGE) deposit (Quebec, Canada), exhibit the highest PGE grades (cf. Seabrook et al. 2004). Two well-studied examples of PGM-bearing komatiite-related sulfide deposits are the ∼ 2.7 Ga Kambalda Ni–Cu sulfide deposit of western Australia and ∼ 2.9 Ga O’Toole Ni deposit in Brazil. These styles of mineralization fall within the ‘type 1’ classification of Lesher and Keays (2002), as outlined above, and are discussed further below. At Kambalda, the PGM are associated with lenses of massive and disseminated sulfide formed after the immiscible segregation of sulfide melts from evolving silicate melts. The PGM associated with these orebodies are < 20 µm in size and are Pt–Pd-rich. Sperrylite

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and moncheite are most abundant, followed by the Pd-rich PGM, including: sudburyite, merenskyite, stibiopalladinite, palladoarsenide (Pd2As), michenerite and testibiopalladinite. Other rare PGM include unidentified minerals of Pd combined with As or Sb. Sperrylite is very abundant in bodies of chalcopyrite-bearing massive sulfide. In contrast, moncheite and the other Pd-rich PGM are observed in veins in the massive and matrix ores, within stringers of sulfides in the footwall rocks, or associated with late-stage hydrothermal veins (Hudson and Donaldson 1984). The association of sperrylite with primary chalcopyrite suggests that it was formed during the magmatic event that segregated the massive sulfide ore (Hudson 1986). In contrast, the association of the other PGM with late-stage veins suggests a post-magmatic origin for the PGM, perhaps linked to the infiltration of hydrothermal fluids during shearing and deformation associated with metamorphism. Lesher and Keays (1984) interpreted the footwall stringers, which are enriched in Pd and Cu, as a product of hydrothermal remobilization. Hudson and Donaldson (1984) suggested that hydrothermal fluids carrying Te, As, and Sb would have infiltrated the massive and net-textured sulfide ore, releasing PGE from the sulfides and producing discrete telluride-, arsenide-, and antimonide-dominated PGM. The O’Toole deposit in the Morro do Ferro Greenstone Belt (Brazil) is a Cu–Ni–Co sulfide lens made up of 65% pyrrhotite, 30% pentlandite, and 5% chalcopyrite with accessory cobaltite–gersdorffite and sphalerite (Marchetto 1990). According to Marchetto (1990), cobaltite and gersdorffite may contain small (generally < 10 µm) inclusions of PGM; the most frequent (i.e., 203/372 grains) is a mineral with a composition varying between kotulskite and melonite, followed by irarsite (123 grains) and sperrylite (20 grains). Other PGM include omeiite [(Os,Ru)As2], osarsite, and a series of unidentified PGM: OsAs5, OsRhAsS, RuTeAs, OsRuAs, and a complex Os–Re–As–Te–Fe–Y–Rh mineral. Marchetto (1990) observed that some portions of oxidized material from the weathered top of the O’Toole deposit also contain PGM. Here, sperrylite is the most common, followed by irarsite and kotulskite–melonite and a single grain of native Pt. Marchetto (1990) suggested that the PGM crystallized early in the first stage of cumulate formation, after the crystallization of magnetite but before the onset of sulfide and arsenide crystallization. This explains why irarsite, sperrylite, and kotulskite–melonite occur frequently as small inclusions in the cores of cobaltite–gersdorffite grains. It was also suggested that since pyrrhotite locally replaces cobaltite– gersdorffite, pyrrhotite likely also crystallized later than the PGM and their host cobaltite–gersdorffite. After deposition and crystallization, the ores underwent regional metamorphism under upper greenschist and low-temperature amphibolite facies conditions. Together with the subsequent weathering of the sequence, this metamorphism may have modified the relative abundances of the PGM (Marchetto 1990; De Almeida et al. 2007). It should be noted that PGE mineralization associated with arsenide enrichment has also been reported from some komatiite localities. For example, the Rosie Ni Prospect in the Duketon greenstone belt (Yilgarn craton, western Australia) contains a suite of sulfides and sulfarsenides together with sperrylite, melonite, and bismuthotellurides (Godel et al. 2012). In addition, Hanley (2007) documented Pt- and Pd-rich arsenides at the Dundonald Beach South locality (Abitibi Sub-province, Ontario, Canada). In both of the latter studies, early segregation of a PGE-rich arsenide melt phase was proposed. By contrast, Prichard et al. (2013a) proposed a hydrothermal origin for PGE mineralization associated with arsenides in the Spotted Quoll Nickel ore deposit (Forrestania greenstone belt, western Australia).

Magmatic Ni–(± Cu–± PGE)–Sulfide Deposits in non-komatiitic rocks Non-komatiitic sulfide ore deposits with significant concentrations of Ni (± Cu, ± PGE) occur in a variety of mafic and ultramafic magmatic associations (Naldrett and Duke 1980; Naldrett 2004; Eckstrand and Hulbert 2007). The parental melts are sourced from the upper mantle and their sulfide component is derived from either the initial S content of the magma and/or from the crustal rocks through which the magma passes. In either case, ore formation is envisaged to proceed via accumulation of an immiscible sulfide fraction, before crystallization

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Figure 19. CI-chondrite (Naldrett and Duke 1980) normalized patterns of samples from drill cores R695 and R713, representative of the upper part of the Keivitsansarvi deposit (Kevitsa; after Gervilla and Kojonen 2002). Gray lines correspond to samples containing arsenides and sulfarsenides.

and differentiation of the sulfide liquid. There is some overlap with LMI environments (e.g., Sudbury and Muskox, Canada; Duluth, USA; Kevitsa, Finland, Fig. 19). In general, however, LMI Ni-sulfide deposits are not recognized as being amongst the world’s major PGE resources; a feature that might be related to dilutional effects on the ore forming process (i.e., low R-factor; Naldrett 1989, 2004). There are relatively few non-komatiite-associated Ni-sulfide ore deposits with significant PGE-enrichments; these are typically associated with mafic–ultramafic sills, stocks or ’tube-like’ intrusions. In fact, a wide range in PGE-contents is observed in non-komatiite-associated Ni-sulfide ore deposits, with highly enriched deposits such as that at Noril’sk-Talnakh (Russia) at one end of the spectrum and relatively PGE-poor deposits such as Jinchuan (China) at the other end. Several examples are described in more detail below, presented in no particular order. The Noril’sk-Talnakh Ni-sulfide deposit. The most economically significant Ni-sulfide deposit that is unrelated to komatiite magmatism (or associated with an LMI) and that also contains appreciable PGE-mineralization is the Ni–Cu sulfide-enriched suite of intrusions at Noril’sk-Talnakh in western Siberia (Russia). The ore deposits are associated with the Siberian flood basalt province and, in particular, with a suite of elongate sill-like mafic intrusions (referred to as chonoliths) that underlie the 3.5 km thick lava sequence. The sills comprise a sequence of olivine-bearing dolerite–gabbro units that can be several 100s-m thick with lengths of several km. Noril’sk is exceptionally enriched in the PPGE and is the world’s second leading producer of Pd (after the UG2 chromitite) as well as producing significant quantities of Pt. The nature and distribution of the PGM in the Noril’sk, Talnakh and Oktyabr’sk deposits are similar (Genkin 1968; Genkin et al. 1981; Genkin and Evstigneeva 1986; Vymazalová et al. 2009). The PGM have not commonly been documented from the disseminated ores of Noril’sk, but isoferroplatinum and cooperite have been reported as being typical (Genkin and Evstigneeva 1986; Vymazalová et al. 2009). More detail exists on the PGM assemblages in the massive sulfide ores that occur as sheets and lenses at the bases of the sills. In general, there is an increase in PGE content with an increase in the concentration of Cu-Fe sulfides (chalcopyrite, talnakhite [Cu9(Fe,Ni)8S16], mooihoekite (Cu9Fe9S16), and putoranite [Cu9(Fe,Ni)S16]), but not with pyrrhotite concentration. The PGM in the massive sulfide ores generally exhibit irregular textural features, occurring

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either as individual grains or clusters of grains of varying size. They are characterized by a considerable range in chemical composition; common minerals include tetraferroplatinum and a diverse assemblage of Pd-dominant PGM (e.g., atokite [(Pd,Pt)3Sn], cabriite (Pd2SnCu), froodite, insizwaite [Pt(Bi,Sb)2], isomertieite (Pd11Sb2As2), majakite (PdNiAs), maslovite, michenerite, palarstanide [Pd8(Sn,As)3], palladoarsenide, paolovite, sobolevskite, stannopalladinite [(Pd,Cu)3Sn2], pašavaite (Pd3Pb2Te2), stibiopalladinite, and taimyrite [(Pd,Cu,Pt)3Sn]) along with an unusually rich variety of Pd–Pb minerals (plumbopalladinite (Pd3Pb2), polarite [Pd,(Bi,Pb)], urvantsevite [Pd(Bi,Pb)2], and zvyagintsevite). A key feature of the massive orehosted PGM at Noril’sk is the wide development of ternary compounds of Pd, Sn, and Cu. Palladium- and Pt-sulfides, bismuthotellurides, and tellurides (braggite, cooperite, kharaelakhite [Pt,Cu,Pb,Fe,Ni)9S8], kotulskite, merenskyite, telargpalite [(Pd,Ag)3Te] and vysotskite) are more typical of vein-hosted and breccia ores that form at sill upper contacts. Genkin and Evstigneeva (1986; their Table 1) provide a detailed list of Noril’sk PGM. These authors envisaged a progressive (three-stage) process as being responsible for the PGM parageneses at Noril’sk. The earliest stage of crystallization, directly from the sulfide melt fraction, is typified by the earlyformed PGM, isoferroplatinum and cooperite, in the disseminated ores. Genkin and Evstigneeva (1986) suggested that the PGM in the massive sulfides were concentrated by a residual Cu– volatile-rich liquid that evolved after separation of the sulfide melt fraction. Incorporation of the PGM into the sulfides was envisaged to be a relatively late-stage metasomatic process as this residual liquid permeated the interstitial spaces between sulfides and microfractures in sulfide grains. The presence of the Cl-bearing sulfide djerfisherite [(K,Na)6(Fe,Cu,Ni)25S26Cl] and a Pd–Bi chloride in the mineral assemblage was used to develop this model. The existence of compositionally zoned intergrowths and clusters of PGM in the massive and vein-hosted ores at Noril’sk suggests that metasomatism and modification of primary PGM assemblages occurred well below the solidus, possibly related to the presence of hydrothermal chloride solutions (Genkin and Evstigneeva 1986). However, the model proposed by Naldrett (2004) and Czamanske et al. (1992) is generally considered to supersede the older model of Genkin and Evstigneeva (1986). Naldrett (2004; see also Barnes and Ripley 2016, this volume) highlighted the widespread evidence for assimilation of country rock and proposed that thermochemical erosion of coal measures provided the S that triggered sulfide saturation at Noril’sk. Separation and differentiation of the sulfides, via concentration in hydrodynamic traps, formed the large ore deposits. The Aguablanca Ni–Cu–(PGE) sulfide deposit. The exotic Aguablanca ore deposit is an economic Ni–Cu sulfide orebody (6.6 Mt at 0.6% Ni and 0.4% Cu) associated with a sub-vertically oriented magmatic breccia, located in the northern portion of the Aguablanca calc-alkaline mafic intrusion (SW Spain; Piña et al. 2008, 2012). Ages from U–Pb dating of zircon indicate that the Aguablanca stock was emplaced at ca. 340 Ma ago when an Andeantype magmatic arc developed during the Hercycinian orogeny in the Ossa Morena Zone (Piña et al. 2008 and references therein). This tectonic setting is very unusual for this type of ore body, but structural and gravity data show that the deposit occurs as a funnel-like body of mineralized breccia, adjacent to the Cherneca ductile shear zone (a Variscan sinistral transpressional structure). The orientation of the breccia zone, possibly representing a feeder zone to the stock, corresponds to that expected for tensional fractures formed within the strain field of the Cherneca shear zone. Piña et al. (2010) distinguished two different events as being important for the origin and emplacement of the Aguablanca deposit, (1) initial ore-forming processes associated with magma emplacement in the crust, assimilation of crustal S, and segregation/gravitational settling of sulfide melt (a scenario similar to most plutonic Ni–Cu sulfide ores), and (2) later-stage emplacement of the Ni–Cu sulfide-bearing rocks by multiple melt injections. The melt injections are envisaged as being controlled by successive opening of tensional fractures related to the Cherneca shear zone. Three types

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of ores were ultimately produced: semi-massive, disseminated, and veined chalcopyriterich ore. Piña et al. (2008) interpreted the semi-massive ore, comprising pyrrhotite and pentlandite with minor amounts of chalcopyrite, as a cumulate of mss enriched in Os–Ir– Ru–Rh and poor in Pd–Pt–Au–Cu. In contrast, the chalcopyrite veins rich in Cu–Pd–Au– (Pt) crystallized from a residual Cu-rich melt formed after the crystallization of the mss. The disseminated ore consists of equal proportions of pyrrhotite, pentlandite and chalcopyrite and represents in situ crystallization of the original sulfide melt. PGM are present in all ore-types, although they are most abundant (71%) in the semimassive and the chalcopyrite-veined (21%) ores, relative to the disseminated (8%) ore (Ortega et al. 2004; Piña et al. 2004). Most of the PGM are rounded and lath-like inclusions in pyrrhotite and pentlandite, although they can be also found along sulfide–silicate and sulfide–sulfide grain boundaries. Only a small proportion of the PGM occur as inclusions in matrix silicates. The PGM assemblage includes, in decreasing order of abundance, merenskyite, melonite, michenerite, moncheite, and sperrylite (Fig. 20a–d). The mode of occurrence and the composition of the PGM suggest that they are magmatic rather than hydrothermal. The positive correlation between the total PGE and S contents in whole-rock samples suggests that the PGE were initially collected by the immiscible sulfide liquid and later exsolved from the mss upon cooling. The fractionation of the IPGE and PPGE between the different ore types, and therefore the formation of distinct PGM, is likely directly controlled by the partition coefficients between the PGE and the sulfide phase.

Figure 20. Images of PGM included in sulfides of the Ni ± (Cu ± PGE) sulfide deposits of Aguablanca in southwestern Spain (image courtesy of Dr Rubén Piña).

The Jinchuan intrusion, northern China. Prichard et al. (2013b) described ~140 PGM grains from fresh and altered samples of olivine–sulfide liquid cumulate from the Jinchuan intrusion, North China Craton. Jinchuan is ~832 Ma (Zhang et al. 2010) and is estimated to

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be the third largest Ni–Cu(–PGE) deposit on Earth. The intrusion is sheet-like in geometry, ∼ 200 m thick and is structurally subdivided into the west, west-central, and eastern subchambers by a series of northeast trending strike-slip faults. It intrudes Paleoproterozoic migmatites, gneisses, and marbles of the Longshoushan Group. Prichard et al. (2013b) studied samples of peridotite (lherzolites and dunites). They found that although a sample of relatively fresh dunite contained the highest abundances of Pt, the types of minerals hosting Pt were unclear. In particular, Prichard et al. (2013b) reported concentrations of ∼ 3 ppm Pt in the dunite, but only 7 grains of irarsite, which they considered insufficient to account for the whole-rock abundance. One PdBi selenide was also documented from this sample. In contrast, they located ~140 PGM grains in more altered samples, including froodite (n = 44), michenerite (n = 24), and padmaite (PdBiSe; n = 24). Other documented phases include 33 PGM of the irarsite–hollingworthite–platarsite solid solution series and minor (< 11% of the total PGM observed) Pt- and Pd-tellurides. Prichard et al. (2013b) concluded that Pd initially partitioned into base-metal sulfides, whereas early crystallization of irarsite and sperrylite may have occurred if As was present (explaining the paucity of Ir and Pt in basemetal sulfides). A second phase of metasomatism involving oxidizing fluids was invoked to explain the Se-bearing PGM. To explain the metasomatism, Prichard et al. (2013b) highlighted a general worldwide association between Pd-selenides and carbonates. They pointed toward the carbonate country rocks that host the Jinchuan intrusion and suggested that the interaction of low pH oxidizing fluids with the carbonates may have been the driver for precipitation of Pd and Se at Jinchuan. The Duluth Complex, Minnesota, USA. The Duluth Complex is one of the largest known layered intrusions on Earth, occupying an area of > 4,700 km2. It was emplaced close to the unconformity between Archean supracrustal rocks and Proterozoic metasediments in the north and comagmatic Mesoproterozoic intrusions and lava flows in the south, at ~1.1 Ga (Hoaglund 2010). Mogessie et al. (1991) described PGM from drill core samples of the basal troctolites of Duluth, including sperrylite, taimyrite, froodite, michenerite, and moncheite. The PGM occur within a zone of Cu–Ni–PGE sulfide mineralization at or close to the base of the Duluth Complex, and are dominantly associated with serpentinized olivine and secondary magnetite, or with hydrous silicate phases (e.g., prehnite, actinolite, hornblende, chlorite), all interpreted as alteration products. Mogessie et al. (1991) suggested that a late hydrothermal magmatic event was responsible for mobilization and re-deposition of the PGM, with possible sourcing of the volatiles from assimilated metasedimentary country rocks. They noted relatively high Cl contents of some of the associated silicates (e.g., serpentine, biotite, apatite) and suggested that transport of the PGM in Cl-complexes may have been a viable mechanism of ore formation. The Sudbury Igneous Complex, Canada. The Sudbury Igneous Complex has been an important source of global Pt and Pd and contains four different styles of mineralization (see also Barnes and Ripley 2016, this volume). Broadly, these include contact ores, offset ores, deep chalcopyrite-rich veins (occur beneath contact ores), and PGE-rich disseminated mineralization around the latter veins. The Sudbury body is an atypical igneous intrusion in that it is believed to have formed from a melt sheet produced after a boloid impact, at ~1.85 Ga. It has been suggested by a number of workers that the Sudbury PGE occur as PGM, rather than in solid solution in base-metal sulfides (Cabri and Laflamme 1976; Cabri 1981b, 1988; Cabri et al. 1984; Li et al. 1993; Farrow and Lightfoot 2002; Huminicki et al. 2005). Ford et al. (2011) carried out a grade recovery exercise utilizing samples from Sudbury’s Coleman and Creighton mines as natural examples. The Coleman and Creighton samples are of ‘sublayer’ material lying immediately above the footwall (country rock) lithologies, on the North and South Range of the intrusion, respectively. Samples from the Coleman Mine yielded 2866 grains, which were fashioned into 26 mounts. The samples from the Creighton Mine yielded a total of 440 PGM grains in 40 thin sections. Both samples exhibit markedly

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different PGM populations, attributed by Ford et al. (2011) and Farrow and Lightfoot (2002) to the interaction of the parental magmas with significantly different (sedimentary/ volcanic vs. gneissic lithologies) country rocks. In the sample from the Coleman mine, six PGM account for 88% of the PGM and precious metals by total area, including Pt–Sn alloy (20%), maslovite (18%), froodite (17%), michenerite (15%), Au–Ag alloy (11%) and moncheite (7%). The Creighton sample exhibits less diversity, with three principal mineral phases accounting for ~90% of the total surface area of the PGM and precious metals, including michenerite (39%), sperrylite (32%), and Pt–Sn alloy (20%). The distribution of Pt from the Creighton sample is dominated by sperrylite, whereas Pt in the sample from Coleman is controlled by bismuthotellurides and Pt–Sn tellurides. The distribution of Pd at Creighton is dominated by michenerite (78% of Pd) whereas both michenerite (37%) and froodite (39%) dominate the Pd budget at Coleman. The results of the study by Ford et al. (2011) are generally in agreement with an earlier study by Farrow and Lightfoot (2002) which found that the Sudbury PGM population is dominated by tellurides, bismuthotellurides and arsenides, with a notable absence of PGE-sulfides or alloys. However, the samples studied by Dare et al. (2010) suggested that PGE sulfarsenides (86%) dominated the PGM at Creighton, with sperrylite (9%), michenerite (5%) and electrum (0.1%) constituting only a minor proportion of the overall budget. This observation highlights that the PGM mineralogy can be highly heterogeneous at the hand specimen scale. Dare et al. (2010) reported that the PGE sulfarsenides are characteristically compositionally zoned, with irarsite cores, hollingworthite mantles, and a PGE-rich Ni cobaltite rim. The sulfarsenides and sperrylite thus concentrate Ir, Rh, Pt ± Os, and Ru, such that accompanying mss accounts for < 10% of the bulk rock Ir, Rh and Pt, and 50–90% of Os, Ru, and Pd. Dare et al. proposed that the PGE-sulfarsenides and sperrylite at Creighton crystallized from a high-temperature (1,200–900 °C) sulfide melt and were subsequently surrounded by the PGE-depleted mss cumulate. Thus, the observed mineralogy is predominantly a magmatic assemblage, with limited low-temperature modification. However, Dare et al. (2010) did suggest a degree of Pd mobility at lower temperatures, linking recrystallization of the base-metal sulfides to michenerite formation at < 540 °C. At higher temperatures and an earlier stage of evolution, pentlandite may have been an important host for Pd (Cabri 1981b, 1988; Cabri et al. 1984).

EXAMPLES OF UNCONVENTIONAL PGM OCCURRENCES Kimberlite- and Cu-porphyry-hosted PGM Kimberlites. Kimberlites are the solidified remnants of volatile-rich ultrabasic potassic magmas that typically develop in cylindrical pipes and contain olivine phenocrysts, megacrysts, and pyroclasts (Sparks et al. 2006). Xenolithic materials in kimberlites can include a variety of mantle lithologies, including peridotites, pyroxenites, and eclogites. Their origin close to the interface between the cratonic continental lithosphere and underlying convecting asthenospheric mantle means that they can provide insights into the structure and composition of the deep mantle. An example of a PGM assemblage in a kimberlite setting is described by Stone and Fleet (1990). The intrusion occurs in Fayette County (Pennsylvania, USA). Two forsteritic olivine megacrysts contain composite (chalcopyrite–pentlandite) sulfide inclusions, two of which each host a single grain of isoferroplatinum. Although Stone and Fleet (1990) discussed the formation of the PGM within the kimberlite system, it also seems plausible that the olivine megacrysts are xenolithic and that these PGM have a mantle (SCLM) affinity. Power et al. (2004) report PGE-mineralization associated with two dioritic sheet intrusions at Talnotry, southwest Scotland. The style of magmatism is broadly similar to kimberlite occurrences in that these sheets are lamprophyric. The intrusions are likely lower-mid Paleozoic in age. Power et al. (2004) recognized three distinct mineralogical assemblages in the Talnotry

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ore deposit, including a pyrrhotite–chalcopyrite assemblage, a niccolite–gersdorffite assemblage, and a chalcopyrite–gersdorffite assemblage. Overall, the PGM assemblage is diverse, with abundant sperrylite, irarsite, and electrum, as well as minor merenskyite, michenerite, and froodite. In the pyrrhotite–chalcopyrite assemblage, > 100 PGM were reported in 19 thin sections. Almost 90% of the grains are sperrylite, irarsite or electrum (50%, 26%, and 13%, respectively; electrum was included as a PGM in this particular study). The remainder are Pd-bearing PGM, including merenskyite, michenerite, and froodite. A total of 63 PGM were observed in 17 thin sections in the niccolite–gersdorffite assemblage. Electrum is the dominant mineral (63%), with Pd–Bi–Te phases (21%), and sperrylite (13%) also significant. The chalcopyrite–gersdorffite assemblage contains sperrylite and electrum, but due to the diffuse and cross-cutting nature of this assemblage, the PGM are more difficult to isolate and identify. Power et al. (2004) proposed a paragenetic sequence where early sperrylite and gersdorffite crystallized first from metal-rich sulfide liquid. Exsolution of Pt and Ir occurred from the gersdorffite, whilst Rh remained in solid solution. Crystallization of mss from the sulfide liquid formed the pyrrhotite–chalcopyrite assemblage, accompanied by separation of the remaining IPGE in the solid fraction, whilst As, Pd, and some Pt remained in the sulfide liquid. Crystallization of the NiAs liquid to form the niccolite–gersdorffite assemblage was associated with removal of the remaining Pt and Pd from the sulfide liquid. A Cu- and Au-rich iss is the final crystallization product, and locally reacts with and cross-cuts the pre-existing ore assemblage. Porphyry-Cu (± Au) deposits. Porphyry-Cu deposits can be significantly enriched in Au, but their potential for PGE enrichment has been less well studied (cf. Eliopoulos et al. 2014). Hanley (2005) suggests that moderately oxidized hypersaline fluids are capable of dissolving ppm quantities of Pt and Au, such that the arc magmas typically associated with porphyryCu deposits may have the potential to be PGE ore deposits. Examples of Pt + Pd enrichment associated with porphyry-Cu ores occur at Copper Mountain and Galore Creek (British Columbia), Allard Stock (Colarado; Werle et al. 1984), the Elatsite porphyry deposit (Chelopech, Bulgaria; Tarkian et al. 2003; Augé et al. 2005b), and the Skouries porphyry deposit (Greece; Economou-Elipoulos et al. 2000). However, detailed accounts of the PGM assemblages of many of these localities have not been reported. The Skouries Cu–Au deposit is ∼ 18 Ma in age and hosted in the Vertiskos Formation of the Serbo-Macedonian massif. The dominant PGM phase at Skouries is merenskyite, which occurs as intergrowths with hessite, electrum and Cu-sulfides (bornite and chalcopyrite). The Elatsite porphyry-Cu deposit is related to multiple monzonitic and monzodioritic intrusions of Upper Cretaceous age (92.3 ± 1.4 Ma), along a portion of the socalled ‘Tethyan Eurasian Metallogenic Belt’. Augé et al. (2005b) describe a PGM assemblage of merenskyite, moncheite, palladoarsenide, and a Pd–Ag–Te–Bi phase associated with base-metal sulfides including bornite and chalcopyrite. The latter authors attribute the PGM assemblage to initially high concentrations of PGE in the magma, subsequently concentrated in hydrothermal fluids, rather than early separation in high-temperature sulfides. As noted by Eliopoulos et al. (2014), an interesting feature of PGM occurrences in porphyry-Cu deposits is that they underline the potentially important roles of aqueous vapors and brines for PGE concentration, these being capable of scavenging the metals in evolved hydrothermal systems. The Bon Accord Ni-deposit, South Africa. The ∼ 3.5 Ga Bon Accord Ni-deposit was situated in the lowermost stratigraphic units of the Barberton greenstone belt of the Kaapvaal craton (South Africa). It was completely quarried out for smelting, never having comprised more than ∼ 6.3 m2 of material (Tredoux et al. 1989). The deposit is located in the Tjakastad subgroup of the Onverwacht Group and contained within serpentinized mafic and ultramafic rocks. It comprised a unique suite of Ni-rich minerals, including Ni-rich spinel (trevorite; NiFe3+2O4) and Ni-rich silicates: liebenbergite (olivine; Ni2SiO4), nimite (chlorite; [(Ni,Mg,Al)6[(OH)8(Si,Al)4O10]) and népouite (serpentine; [(Ni,Mg)3[(OH)4Si2O5]]). There is also an abundance of accessory phases, such as Ni-rich sulfides, arsenides and antimonides, bunsenite (NiO), bonaccordite (Ni2FeBO5)

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and gaspéite ([Ni,Mg]CO3), some of which are likely unique to Bon Accord (Tredoux et al., in press). There is some controversy over the origin of the deposit, with two competing hypotheses emerging in the 1970’s and 1980’s. The first suggested that Bon Accord had an extraterrestrial provenance (an iron meteorite; De Waal 1978) whilst the second argued that the ore body represents an oxidized fragment of Fe–Ni alloy originating from the Earth’s core (Tredoux et al. 1989). Other alternatives such as oxidation of either (a) podiform chromitite or (b) awaruite (Ni–Fe alloy stabilized during the highly reducing conditions associated with serpentinization) were dismissed (Tredoux et al. 1989). Although formation of the deposit as an oxidized volcanic massive sulfide (VMS) deposit or as a modified Ni–sulfide deposit was also ruled out by the latter authors, these possibilities have recently been revisited (O’Driscoll et al. 2014b). Zaccarini et al. (2014) and Tredoux et al. (in press) have recently reported the presence of PGM in Bon Accord material. This is not surprising, since Tredoux et al. (1989) reported ppm levels of the PGE. Zaccarini et al. (2014) report a variety of Pd–Sb Pd–Sb, Pd–Sb–As, Pd–Cu–Sb, Pt–Sb, Pt–As–S, Ru–As–S, and Ru–S phases, as well as sperrylite and phases in the solid solution series sobolevskite–kotulskite. Electrum was also reported. Tredoux et al. (in press) report the presence of native Ru alloys (< 1 µm in diameter) in their samples. Although heavy mineral separation techniques were used by Zaccarini et al. to isolate the PGM, they were able to provide some textural context for the PGM, which appear to be preferentially sited close to the edges of Ni-arsenide and Ni-antimonide grains, in turn typically encased in trevorite grains. Zaccarini et al. (2014) suggest that the observed assemblage was formed during a hydrothermal or low temperature metamorphic episode, possibly during exhumation of the body, but this plausibly lends support to the origin of Bon Accord as an altered Ni-sulfide deposit too. It is difficult to speculate much further than this on the petrogenesis of the PGM assemblage in the Bon Accord deposit. However, it is worth highlighting the importance of future study on what is one of the oldest (reported to date) PGM assemblages in the (terrestrial) geological record.

OUTLOOK AND FUTURE WORK This chapter has described the occurrence and petrogenesis of the PGM in numerous natural examples, with an emphasis on settings where PGE-enrichment is known to occur. An important point arising from this review is that whilst many PGM may initially form under high-temperature (magmatic; 900–1200 °C) conditions, modification of such assemblages can occur down to much lower temperatures (e.g., ∼ 300 °C). A good example of this is the Platinova reef of the Skaergaard intrusion, where the final PGM assemblage observed is the product of a range of processes, including low-temperature equilibration. It is worth concluding this chapter by mentioning some of the outstanding problems in the field of PGM petrogenesis, as well as highlighting areas for future work.

Assessing the mineralogical and textural complexity of PGM assemblages It is evident that even well-characterized PGM assemblages exhibit features that are not well understood. For example, the Merensky Reef of the Bushveld contains one of the most intensely studied PGM assemblages on Earth, the observations on which raise as many issues as they resolve with respect to formation of the reef. The distribution of the PGM in the Merensky Reef varies from one locality to the next, and from pothole-style reef to normal reef, despite the relative (lithological) similarity of the host rock cumulates over 100’s of km of strike (e.g., Table 2). Paradoxically, the overall grade of the PGE remains relatively consistent. Similarly, each of the Rum Layered Suite chromitites, which are typically ∼ 2 mm thick, has a distinctive PGM assemblage that is relatively consistent over 100’s m of strike. However, the PGM assemblages of the chromitites can vary considerably up or down stratigraphy. For example, the Unit 11/12 chromitite is dominated by arsenides (i.e., sperrylite), whereas the Unit 7/8 chromitite is dominated by Pt–Fe alloys and PGE sulfides. Despite this, the PGE

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abundances of the chromitites vary by one order of magnitude or less (O’Driscoll et al. 2009a). These observations suggest that PGE grades may not be wholly dependent on the PGM mineralogy. Perhaps it is the high-temperature magmatic system that fixes the PGE budget of a reef, and the mineralogy is simply a consequence of the whole-rock PGE geochemistry. Another complexity that has arisen in studies of PGM mineralogy is the presence of xenocrystic PGM in ophiolite chromitite. For example, González-Jiménez et al. (2014a) reported 187Os/188Os heterogeneities within single PGM grains, some of which yielded model ages that are older than the host chromitite. Thus it is clear, in upper mantle examples at least, that recycling of PGM may occur, complicating the interpretation of the observed assemblage. Establishing whether xenocrystic PGM form a statistically significant proportion of PGM in all ophiolite chromitites, and whether they exist in PGM assemblages in other settings, is therefore an important target for future study. There is a clear, if not ubiquitous, link between the formation of chromitite seams and concentration of PGM in LMI (e.g., Fig. 4), the upper mantle and CUAAC. A critical factor here appears to be suppression of silicate phases on the magma liquidus, enabling the precipitation of non-silicates (such as chromite) only. An observation that seems to hold for many chromitite occurrences, irrespective of setting, is that IPGE-rich PGM (i.e., laurite) are preferentially hosted in inclusions in chromite, whereas PPGE-rich phases occur interstitially. An important advance in our understanding of this relationship was the work of Finnigan et al. (2008), which experimentally documented the nucleation of IPGE-rich PGM in boundary layers around growing chromite crystals. Given the importance of chromitite in studies of upper mantle Os isotope heterogeneity (e.g., Walker et al. 2002), placing further constraints on the distribution of the PGE in chromitite is a very useful goal. In general, this review has shown that many published models favor the high-temperature separation and fractionation of sulfide melt, the subsequent crystallization of which may allow the PGE (together with As, Te, and Bi) to be concentrated in the residual melt fraction. However, early crystallization of PGM from the sulfide melt (e.g., Dare et al. 2010), or desulfurization of base-metal sulfides which fractionated with the PGE in solid solution (e.g., Fonseca et al. 2012; Bowles et al. 2013) have also been suggested. Looking forward, progress in our understanding of PGM petrogenesis will be made by consideration of the PGM in conjunction with their coexisting sulfide and silicate parageneses, rather than as an isolated mineral group.

Constraints on quantifying the distribution and grain size of PGM A significant outstanding challenge to future investigations of the PGM is their small grain size. Indeed, the lower grain size limit for these minerals is quite poorly constrained. A good example of this problem is the Bon Accord deposit. Although Tredoux et al. (1989) reported ppm levels of all of the PGE, PGM have proven difficult to find and where present do not explain the observed whole-rock abundances (Zaccarini et al. 2014). In addition, the relevance of fully understanding the textural and compositional diversity and distribution of PGM present in a sample should not be underestimated. For example, Godel (2013) showed that small widely distributed PGM tend to be over-sampled relative to coarser-grained sparsely distributed grains. The latter may contain an abundance of some or all of the PGE, leading to problems in the interpretation of the whole-rock geochemistry. Godel (2013) highlighted the utility of X-ray micro-CT scanning in addressing these problems (see also Godel et al. 2010). The advantage of this technique is that the composition and distribution of the PGM can be studied in 3D, although sub-µm grain sizes are still challenging to analyze. For example, König et al. (2015) carried out X-ray micro-CT scanning of PGM in a harzburgite from the Lherz peridotite (in material originally studied by Luguet et al. 2007). They reported IPGE-rich PGM inclusions in olivine that they interpreted as residual after mantle melting (sulfide exhaustion). These PGM are relatively abundant, compared to

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the interstitial PGM, which were considered to be metasomatic. König et al. (2015) quoted a lower limit of 2.5 µm for the resolution of the micro-CT scanning technique, but were able to document grain sizes in the range 3–5 µm. Advances in microbeam techniques such as LA-ICP-MS have highlighted the wealth of information that may be extracted from individual PGM and base-metal sulfides, for example where isotopic information (e.g., 187Os/ 188Os and 186Os/188Os) can be derived (Coggon et al. 2011b; González-Jiménez et al. 2012a). Further progress will undoubtedly be made through increasing the resolution of existing approaches, such as in situ elemental mapping, and the application of new instrumentation. For example, the NanoSIMS offers the potential to investigate other isotope systems (i.e., S) in PGM and coexisting sulfide and arsenide phases, opening up avenues for investigation of the petrogenetic histories of these minerals by assessing isotopic heterogeneity in domain sizes of 10s–100s nm. In addition, field-emission SEMs are capable of analyzing samples at higher resolutions (10s–100s nm) than conventional SEMs and synchrotron-based XRF mapping and high-resolution transmission electron microscopy (HRTEM) also offer future potential in PGM investigations.

Advancing our understanding of the link between PGM assemblage and PGE geochemistry In general, the whole-rock PGE budgets of natural samples from different environments vary. For example, ophiolite peridotites and especially chromitites are characterized by PGE patterns on chondrite-normalized diagrams that have negative slopes, signifying relative enrichments in the IPGE relative to the PPGE. By contrast, the distributions of the PGE from layered intrusion lithologies, including many chromitites, typically have patterns with a positive slope on similar diagrams. Exceptions do occur, such as the Cliff chromitites in the Shetland Ophiolite Complex (Prichard et al. 1986; Prichard and Lord 1993), which have anomalously high PPGE abundances (> 250 ppm Pt + Pd; O’Driscoll et al. 2012) and strongly positive patterns on chondrite-normalized diagrams. The high abundances of the PPGE here can be linked to the abundant sperrylite at this locality. The typical ‘M-shaped’ PGE patterns of CUAAC-associated rocks reflect relatively positive anomalies of Pt and Ir, which likely indicate a mineralogical control on the fractionation of the PGE. However, as highlighted for the Merensky Reef above, the extent to which the PGE patterns are a function of the mineralogy or perhaps the mineralogy is the result of the whole-rock chemistry (allowing for down temperature processes to affect the mineralogy but not change the bulk chemistry drastically) for such samples is an important line of investigation for future studies. One important conclusion that can be drawn from the majority of the studies described in this chapter is that, where present, PGM account for the whole-rock budget of most of the PGE. One key exception is Pd, and an important focus for additional study should be the ubiquity of Pd occurrences in solid solution in base-metal sulfide, vs. the occurrences of Pd-rich PGM. Junge et al. (2015) recently reported the presence of µm- and nm-sized PGM in pentlandite from the UG2 chromitite and the Platreef of the Bushveld Complex. They observed no orientation (crystallographic) relationship between some of the PGM nanoparticles and the sulfide host, leading them to suggest that the PGM were discrete phases trapped during pentlandite growth. However, they also suggested that Rh and Ir were present in the sulfide liquid. Palladium was observed as nanoparticles and also homogeneously distributed throughout pentlandite grains, suggesting both modes of occurrence are possible. Brenan et al. (2012) have shown experimentally that Ru may also partition into chromite, although this may not be significant for PGE-enriched chromitites where chromite has relatively low ferric iron contents. Additional experimental investigation will be invaluable in further elucidating our knowledge of PGE solubility in silicate and sulfide melts (cf. Mungall and Brenan 2014; Brenan et al. 2016, this volume). Although not the focus of this chapter and not an area where a large amount of research has been carried out, the basic solubilities of the PGE in aqueous solutions in the presence of As-bearing ligands (or Bi and Te) also has the potential to have implications for PGE mineralization in nominally hightemperature deposits (cf. Hanley 2005).

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BO’D acknowledges research support from a NERC New Investigator grant NE/ J00457X/1 during the early stages of writing this paper. Support for this study has also been provided by the FONDECYT #11140005 and ‘Millenium Nucleus for Metal Tracing Along Subduction NC130065 to JMGJ. The following are thanked for their permission to use various images and illustrative material (in the case of published studies, the first author is listed): JFW Bowles, B Godel, G MacPherson, KN Malitch, T Oberthür, R Piña, D Schwander, RJ Voordouw. The authors are also grateful to Louis Cabri and Fernando Gervilla for discussion and feedback on early manuscript versions. Detailed reviews by Andrew McDonald and Steve Barnes as well as editorial handling and comments by Jason Harvey, James Day, and Ian Swainson helped improve the clarity and focus of the manuscript.

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PGM in placers associated with ophiolite complexes Platinum-group minerals have been reported in many placer deposits derived from ophiolites (Supplementary Table 2). These detrital PGM have generally been sampled in chromite-rich beach sands associated with river terraces or coastline fans, formed as a result of mechanical erosion of chromite-bearing peridotites (e.g., Meibom et al. 2002 and references therein). The vast majority of ophiolite-derived PGM-bearing placers are uneconomic but some of them were mined using artisanal methods for their associated gold in the Middle Ages (e.g., Krstić and Tarkian 1997 and references therein). Platinum-group mineral bearing placers are associated with ophiolites in Papua New Guinea (Harris and Cabri 1991; Weiser and Bachmann 1999), Adamsfield in Tasmania (Hattori and Hart 1991; Brandon et al. 1998, 2006; Pearson et al. 2007), Meratus-Bobaris in Borneo (Harris and Cabri 1991; Hattori et al. 2004; Coggon et al. 2011), Samar in the Philippines (Nakagawa and Franco 1997), Hokkaido in Japan (Hattori and Hart 1991; Nakawaga and Franco 1997; Hirata et al. 1998), Veluce in Yugoslavia (Krstić and Tarkian 1997), the Rhodope Complex (Tsintsov and Damayanov 1994; Tsintsov 2000, 2003, 2004), Sagua de Tánamo in eastern Cuba (Díaz-Martínez et al. 1998), southwestern Oregon and northern California (Walker et al. 1997; Bird et al. 1999; Meibom and Frei 2002; Meibom et al. 2004; Walker et al. 2005; Pearson et al. 2007), and in the ophiolites of Karaginsky in the Kamchatka peninsula (Tolstykh et al. 2009). Additionally, placer PGM have been found in laterites developed on the ophiolites of Samar and Dinagat islands (Franco et al. 1993; Nakawaga and Franco 1997) and the Pirogues River in New Caledonia (Augé and Legendre 1994). Ophiolite placer PGM occurrences. Approximately 80% of PGM in placers associated with ophiolites are Os–Ru–Ir alloys (Supplementary Table 2), comprising many of the possible mineral species. Other significant phases, in order of abundance, are laurite, sulfarsenides of Ir–Pt–Rh (irarsite, platarsite, and hollingworthite), arsenides (sperrylite) and Pt–Fe–Cu alloys (isoferroplatinum, tetraferroplatinum, and tulameenite). Other accessory PGM in these types of placer deposits include Pt-sulfides (cooperite, cuproiridsite), Pt-tellurides (moncheite) and rare chengdeite (Ir3Fe). The Os–Ir–Ru alloys show a variety of grain sizes (from a few µm up to a few mm in the largest dimension) and morphologies (from rounded to angular). The Os– Ir–Ru alloys may be homogeneous, zoned or porous. The intergrowth or exsolution patterns of the different IPGE may result in complex internal structures, which have been used to distinguish these types of alloys. For example, Bird et al. (1999) and Walker et al. (2005) differentiated the Os–Ir–Ru alloys in the southwestern Oregon placers as being portions of either lamellae or host matrix (see Figs. 1 and 2 of Walker et al. 2005). Inclusions of other PGM, base-metal sulfides or silicates in the Os–Ir–Ru alloys are not uncommon, but are much less frequent than observed in Pt–Fe alloys of CUAAC. They typically consist of other microPGM alloys that are co-genetic, exsolution products or silicates. In addition to the alloy phases, other PGM in ophiolite-derived placers may show a range of morphologies and sizes. Hattori et al. (2004) reported different grain morphologies in placer-hosted laurite from Borneo; ‘…including euhedral grains with conchoidal fractures and pits, and spherical grains with no crystal faces, probably because of abrasion…’. Placer laurites often contain inclusions of silicates (amphibole, epidote, clinopyroxene, serpentine, olivine, anorthite), BMS (base-metal sulfide; pyrite, pyrrhotite, pentlandite and chalcopyrite) and composite inclusions of silicate + sulfide + alloy (Coggon et al. 2011). Similarly, Tsintsov (2004) recognized two types of placer sperrylite based on ‘…the degree of mechanical processing during the exogenic transport: very strongly or relatively weakly processed…’. According to the latter author, sperrylite that undergoes abrasion during transport exhibits smooth grain surfaces; these grains partially preserve crystal faces but the edges and apices between them are considerably smoothed. In contrast, sperrylite that was

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transported over short distances has undergone less mechanical abrasion, resulting in better preservation of crystal faces but with pitted surfaces and/or mechanical defects formed by breakages during transport. In addition to mechanical alteration of placer PGM, chemical alteration is also observed. Augé and Legendre (1994) reported abundant PGM oxides in the placers of the Pirogues River (New Caledonia), including Pt–Fe oxides, Ir–Fe–Pt–Rh oxides (Ir–Fe–Rh, Fe–Rh–Pt, Pt–Ir–Fe–Rh) and Ru–Mn–Fe oxides. These oxide grains show large infilled cracks, suggesting volume loss by replacement of a precursor mineral as well as bands concentrically arranged around partly desulfurized PGM (e.g., laurite). Augé and Legendre (1994) interpreted these microstructures as evidence that the remobilization of PGE in the supergene environment may facilitate growth of secondary PGM. More recently Hattori et al. (2010) analyzed these oxides using X-ray absorption spectroscopy and confirmed that these minerals are not true PGE-oxides, but likely comprise a sub-microscopic intergrowth of native PGE with Fe-oxides. Krstić and Tarkian (1997) also reported the presence of RuO2 in placers associated with the Veluce ophiolite (Serbia). Another example of remobilization of the PGE in placers was provided by Díaz-Martínez et al. (1998) who described composite PGM nuggets in the Sagua de Tánamo River with at least three generations of laurite and Ir-arsenide. These particles consist of an early generation of Os-rich laurite (magmatic) replaced by a network of veins of secondary Os-free laurite; the contact between these two types of laurites is marked by a third variety having a composition intermediate between both. Pt–Re–Os and S isotope systematics of ophiolite-derived placer PGM. The examination of the Re–Os–Pt isotopic compositions of detrital ophiolite PGM has greatly contributed to establishing fundamental constraints on the Os isotopic composition and heterogeneity of the convecting upper mantle. A number of studies have reported Os isotopic analyses of detrital Os–Ir–Ru alloys and laurites from placer deposits derived from ophiolites. Hattori and Hart (1991) observed significant dispersion of the 187Os/186Os ratios (1.022 ± 0.006 to 1.050 ± 0.006; 187 Os/188Os = 0.1229 ± 0.0007 to 0.1263 ± 0.0007) in five Os-rich iridium grains from two massifs (Teshio and Onnabetsu) on Hokkaido, Japan. One osmium grain analyzed by Hirata et al. (1998) in the Horonobe ophiolitic massif of Hokkaido Island yielded 187Os/188Os = 0.12611 ± 0.00028). The latter authors noted that there were no significant variations in the 187Os/188Os isotopic ratios within the grain but important changes in the 183W/188Os ratio (0.002–0.006), which they interpreted as due to zoning or multi-stage growth. An analysis of one osmium grain in the same study yielded 187Os/186Os = 1.066 ± 0.006 (187Os/188Os = 0.1282 ± 0.0007). In the placers associated with ophiolites on Kalimantan (Borneo), the Os–Ir–Ru alloys (including iridium, osmium and ruthenium) yield 187Os/186Os ratios between 1.041 ± 0.006 (187Os/188Os = 0.1252 ± 0.0007) and 1.084 ± 0.006 (187Os/188Os = 0.1304 ± 0.0007), less radiogenic than laurite (1.044 ± 0.10 to 1.088 ± 0.10; Hattori and Hart 1991; 187Os/188Os = 0.1256 ± 0.0120 to 0.1304 ± 0.0120). More recently, Coggon et al. (2011) analyzed the Pt–Re–Os isotopes of 260 PGM grains, including laurite, Pt–Fe alloys and one sperrylite from the Meratus Mountains (Borneo). Laurite yields 187 Os/188Os between 0.13445 ± 0.000014 and 0.122117 ± 0.000035 (n = 81); a much more restricted range than that shown by the Pt–Fe alloys (0.125063 ± 0.000042 to 0.140674 ± 0.000056; n = 178). Collectively, the PGM yield a Pt–Os isochron age of 197.8 ± 8.1 Ma (2s), which fits well with published age constraints for the ophiolite in question. Hattori and Hart (1991) analyzed the Os isotope compositions of five grains of osmium from the placers of Adamsfield (Tasmania) that yielded 187Os/186Os isotopic ratios between 1.013 ± 0.006 and 1.028 ± 0.006 (187Os/188Os = 0.1219 ± 0.0007 to 0.1237 ± 0.0007). Walker et al. (1997) and Brandon et al. (1998) carried out repeat analyses of an iridium grain from 19 Mile Creek (Tasmania) and obtained 186 Os/188Os = 0.1198346–0.119841 and 187Os/188Os = 0.123801–0.123824. Pearson et al. (2007) analyzed 72 osmium grains from different localities worldwide that yielded a wide range of 187 Os/188Os ratios between 0.11965 ± 0.00003 and 0.12234 ± 0.00002.

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Several detrital Os–Ir alloys from different placers (southwestern Oregon and northern California) have been examined for Re–Os isotopes by different authors. For example, Hirata et al. (1998) carried out five spot analyses on one grain of Ir from Lower River (California), which yielded variable 186Os/188Os (0.12047 ± 0.00012 to 0.12060 ± 0.00002) and 187Os/188Os (0.12127 ± 0.00007 to 0.121322 ± 0.00007). Walker et al. (1997) and Brandon et al. (1998) analyzed the Os isotopic composition of Os-rich Ir grains derived from different ophiolitic massifs of California and Oregon and revealed relatively homogeneous 186Os/188Os (0.119829– 0.1198361) but heterogeneous 187Os/188Os (0.119940–0.12329) isotopic ratios. The Ir and Os grains analyzed by Bird et al. (1999) from the same locality yield roughly similar 186Os/188Os (0.1195563–0.1198489) but more radiogenic 187Os/188Os (0.124787–0.129307) isotopic ratios. More recently, Meibom and Frei (2002), Meibom et al. (2002, 2004), Walker et al. (2005) and Pearson et al. (2007) have used N-TIMS and LA-ICP-MS to provide a more complete picture of the Os isotopic composition of the placer iridium and osmium grains derived from tectonic peridotites in northwestern California and southwest Oregon. A total of 739 Os–Ir alloy grains reveal large heterogeneities of 187Os/188Os (0.10953 ± 0.000003 to 0.18703 ± 0.0003) and yield TRD model ages as old as 2.7 Ga (Pearson et al. 2007). Malitch et al. (2011) also used N-TIMS and LA-ICP-MS to analyze Re–Os isotopes on 19 detrital Ru–Os–Ir alloy grains from the Kunar massif in the Chelyuskin ophiolite (northeastern Taimyr, Russia). The 187Os/188Os ratios of the studied alloys ranged from 0.1094 ± 0.0004 to 0.1241 ± 0.0004, with corresponding model ages (relative to the present-day chondritic uniform reservoir, 0.12863 ± 0.00046; Chen et al. 1998) calculated between 0.64 and 2.6 Ga. Hattori et al. (2004) analyzed the sulfur isotopes of 14 placer laurites from Tambanio (Borneo). They noted that ‘…the values of all grains show a narrow spread in d34S values (+1.16 ± 0.36 ‰) and minor enrichment in 34S, but they are very close to the meteorite standard value 0. The values are independent of the morphology and the composition of the grains…’. The origin of ophiolite-derived placer PGM. The origin of PGM found in placers associated with ophiolites is subject to debate, and two principal models exist as follows: (1) they simply represent magmatic grains that were liberated from their primary bedrock; (2) they are grains formed in the supergene setting as a result of remobilization of the PGE under appropriate conditions. The placer PGM studied by Hattori et al. (2004) and Coggon et al. (2011) in Borneo are a good example for illustrating the first hypothesis. The placer laurite grains described by these authors contain euhedral and droplet-like inclusions of BMS (made up of aggregates of chalcopyrite + bornite + pentlandite + heazlewoodite) identical to those found in primary laurites hosted in the chromitites of the Sagua de Tánamo ophiolite (eastern Cuba; GonzálezJiménez et al. 2012). The nature of the BMS and the shape of the inclusions suggest that these inclusions crystallized contemporaneously with laurite at high (magmatic)-temperatures. This interpretation is consistent with the fact that these placer laurites also yield S isotopic ratios typical of mantle sulfides, as opposed to the fractionated and variable S isotope signatures expected from grains formed in the supergene environment (Hattori et al. 2004). These sulfides, like the coexisting Pt–Fe alloys in the placers, also contain a suite of silicates typically precipitated at high T from a magma, clearly indicating that these grains are primary in origin. Coggon et al. (2011) interpreted the similar 187Os/188Os ratios in these placer and in situ chromitite-hosted PGM to be additional evidence for the primary origin of the placer PGM. Their placer PGM yielded a relatively precise 186Os/188Os vs. 190Pt/188Os isochron age (197.8 ± 8.1 Ma (2s); MSWD = 0.90), fitting well with the published age of the ophiolite. They concluded that the Pt–Os isotopic system is unaffected in single PGM during surficial processes and can be used to track the origin of the detrital PGM. These results confirm the previous observation of Malitch et al. (2003), who interpreted the similar ‘unradiogenic’ 187 Os/188Os ratios in placer PGM in the Chelyuskin ophiolite and in-situ PGM in chromitites

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of the Kraubath ophiolite, as evidence that the Re–Os system ‘…remains unchanged from the time of formation of the PGM until now, despite later thermal events, which occurred in the vicinity of ophiolite-type complexes…’. The second hypothesis for the origin of placer PGM is illustrated by so-called PGM oxides from the laterites of the Pirogues River (Augé and Legendre 1994). The PGM oxides replace pre-existing magmatic sulfides and Pt–Fe alloys, providing clear evidence that PGM can grow in the supergene environment (e.g., Aiglsperger et al. 2014). Platinum–Re–Os isotopic data for these PGM do not exist, so evaluating how robust the isotopic systems are under these specific conditions of alteration is difficult. However, it might be expected that PGM precipitation from (or interaction with) surficial solutions may have significantly modified 187 Os/188Os ratios, as these solutions typically contain both crustal 187Re and common 187Os. The recent results of González-Jiménez et al. (2012) have shown that metamorphic fluids are able to disturb the Re–Os isotopic system within single PGM grains and more importantly, the secondary PGM that have precipitated from these fluids yield 187Os/188Os ratios within the range of typical mantle values. If this situation applies to secondary PGM in placers, then only detailed petrographic study will reveal the true origin of the PGM grains. Interestingly, alluvial ‘lamellae’-type Os–Ir alloy grains from southwestern Oregon are richer in Fe and more radiogenic than ‘matrix’-type alloys (Walker et al. 2005), suggesting the possible contribution of common Os by secondary oxidizing (i.e., magnetite-bearing) solutions. The very radiogenic compositions of some PGM found in placers associated with the Freetown Complex (Bowles et al. 2000) are also worth considering in this regard (see main text).

PGM mineralization in CUAAC placer deposits The CUAAC are usually associated with placer deposits that contain PGM (Weiser, 2002). Some of the most important are associated with the massifs of the Russian Platinum Belts in the Urals (Nizhny Tagil and Uktus) and Koryak-Kamchatka (Galmoenan), the margins of the Aldan Shield (Inagli, Guli, Zolotaya and Fadeevka), Tulameen (British Columbia), Alto Condoto (Colombia), Fifield (Australia) and Red Mountain (southern Goodnews Bay, southwestern Alaska) (Supplementary Table 3). Placer deposit PGM similar to those found in the aforementioned CUAAC have also been reported in river systems draining poorly exposed mafic and ultramafic rocks of the Yukon (i.e., Florence and Burwash Creeks) and British Columbia (island-arc terrane of Quesnellia) territories in Canada, the Esmeraldas Province (Ecuador), Atonambao-Manamposty (Madagascar), Kompiam (Papua New Guinea), the Serpentinite Hill Complex (Dundas Trough, western Tasmania) and with the laterite of the Yubdo Deposit in Ethiopia (Supplementary Table 3). Occurrences of CUAAC placer-hosted PGM. Weiser (2002) presented a list of PGM associated with selected placer deposits worldwide. He showed that PGM in placers associated with CUAAC are typically in the size range of several tens of µm to a few mm, with a few exceptional grains reaching up to several kilograms (e.g., 11.641 kg for a Pt–Fe alloy from Colombia and 5 kg for a Pt–Fe alloy found in Nizhny Tagil in the Urals). In these types of placer deposit, the PGM are rarely single homogeneous grains but they commonly form heterogeneous nuggets made up of extensive PGM intergrowths with smaller inclusions and/ or exsolution of other PGM, silicates, base-metal sulfides or oxides. Some exotic nuggets may contain large euhedral inclusions of chromite or magnetite with diameters of up to a few tens of µm, totally surrounded by interstitial platinum (e.g., Tulameen, Nixon et al. 1990; Nizhny Tagil, Augé et al. 2005). Most placer deposits associated with CUAAC have a PGM assemblage dominated by large grains of isoferroplatinum or Pt characterized by rounded shapes and/or well-preserved crystal faces, which often contain bleb-like inclusions or crystallographically oriented exsolution lamellae of Os, Ir, and (rarely) Ru (Supplementary Table 3). Where present, such

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exsolution textures are interpreted as evidence for a high temperature origin of the PGM. Other smaller PGM trapped by the Pt-alloy at high temperatures are euhedral grains of the laurite– erlichmanite solid solution series and a variety of sulfides of Ir–Pt–Rh ± base metals (bowieite, xingzhongite, kashinite, and thiospinels such as cuproiridsite, malanite, and cuprorhodsite), and Pt–Pd (cooperite–braggite). The sulfides are accompanied by sulfarsenides (platarsite, hollingworthite and irarsite), arsenides (cherepanovite and vincentite), keithconnite and tolovkite. Other Pt–Fe alloys such as tetraferroplatinum, tulameenite, ferroanplatinum or hongshiite may also comprise the main part of the nugget or occur as intergrowths with isoferroplatinum. The placer deposits associated with the Atonambao-Manamposty CUAAC in Madagascar contain a PGM assemblage dominated by a diversity of unidentified Pd, Pt, Rh, Ir, and Os sulfides that coexist with keithconnite, vincentite, and unnamed Pt3Cu. A further example of a placer deposit with a peculiar PGM assemblage has been reported from the Guli Massif (Aldan Shield), which contains abundant PGM of Os, Ir, and Ru, including large grains of osmium (Malitch and Kostoyanov 1999; Malitch and Badanina 2011; Malitch et al. 2011; Merkle et al. 2012; Supplementary Table 3). In addition to minerals interpreted as high-temperature phases, placer nuggets may contain PGM assemblages formed at low-temperature. The high-temperature PGM assemblage comprising placer nuggets is commonly subjected to lower temperature modification during metamorphism, serpentinization and alteration. The resultant secondary PGM typically form rims around, or fill fractures penetrating into, the primary PGM. In the CUAAC placers that drain the Urals and Koryak–Kamchatka platinum belts, the secondary PGM assemblage is dominated by Pt–Fe–Cu alloys (i.e. tulameenite, Pt-rich copper and unnamed Pt–Fe–Cu compounds), sulfarsenides and/or Ir–Rh–Os such as irarsite, hollingworthite, and osarsite. Other rare secondary PGM found in the placers of Nizhny Tagil include Pd and Pd2S7. Overall, this assemblage is similar to that observed in placers associated with the Tulameen Complex, comprising native platinum and Pt–Fe–Cu alloys (undefined Pt–Cu and Pt–Cu), antimonides (genkinite, geversite), sulfantimonides (tolokvite and undefined RhSbS) hollingworthite, sperrylite, and kotulskite. Secondary PGM found in the placers of Goodnews Bay (Alaska) are also Pt–Fe–Cu alloys (tetraferroplatinum and tulameenite) and sulfarsenides (irarsite, platarsite, and osarsite). The only secondary PGM identified in the Atonambao-Manamposty CUAAC is an Ir oxide. Shcheka and Lehmann (2007) have reported superfine (3–5 µm) discontinuous Au films on Pt–Fe alloys grains often overgrown by cooperite rims in the placer nuggets of Fadeevka (eastern Russia); these authors interpreted that the Au rims formed via selective Pt–Fe leaching during low-temperature alteration and/or weathering of the PGM alloy. Large Pt–Fe or Os–Ir alloy nuggets from placers typically contain numerous inclusions of silicates, which may have been formed before (imposing their crystal morphology on the surrounding alloys) or after (by exsolution during cooling) their host alloy (Johan 2002). In some cases the silicate inclusions are monomineralic but more commonly they are polyphase, preserving an approximation of the composition of the original mineralizing fluid/melt. Primary silicates trapped at high temperature by the alloy include both anhydrous (olivine, pyroxene, plagioclase, K-feldspar, quartz, silicate-glass) and hydrous (amphibole, phlogopite, biotite, vermiculite) phases. Frequently observed secondary silicates in placer PGM nuggets are talc and chlorite (Johan et al. 2000; Johan 2002; Malitch and Thalhammer 2002; Tolstykh et al. 2004, 2005; Okrugin 2011). Other inclusions hosted in the alloy nuggets are oxides (chromite, spinel, magnetite, titanite, ilmenite, corundum), diamond and apatite (cf. Weiser 2002 and references therein). Re–Os-isotopes in PGM from CUAAC-associated placer deposits. In their pioneering work, Hattori and Cabri (1992) used an ion microprobe to analyze the Os isotopic compositions of Os–Ir alloys and laurite–erlichmanite grains included in larger Pt–Fe alloys from a suite of placers, including those associated with the Tulameen (British Columbia), Chocó (Colombia),

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Niznhy Tagil (Russia), Yubdo (Ethiopia) and Goodnews Bay (Alaska) CUAAC. The nuggets from Yubdo yield very homogeneous 187Os/186Os ratios that range between 1.019 ± 0.010 and 1.034 ± 0.008 (187Os/188Os = 0.1226 ± 0.0012 to 0.1244 ± 0.0010). Hattori and Cabri (1992) also note that the PGM from the Niznhy Tagil have an average 187Os/186Os value of 1.029 ± 0.006 (187Os/188Os = 0.1238 ± 0.0007). The latter value is the same, within error, as the average 187 Os/186Os value of 1.035 ± 0.006 (187Os/188Os = 0.1245 ± 0.0007) measured for a PGM nugget from the Omutnaya River that drains the Omutnaya Dunite Massif, considered to be part of the same ultramafic belt as the Niznhy Tagil Massif and situated 180 km to the south. More recently, Malitch and Kostoyanov (1999) and Malitch and Thalhammer (2002) used N-TIMS to analyze the 187Os/188Os ratios of Os–Ir alloys recovered from placer deposits associated with the Guli, Kondyor and Inagli CUAAC. All of the alloys are characterized by low Re (< 0.05 wt.%) and homogeneous 187Os/188Os; ranging between 0.1248 ± 0.0003 and 0.1252 ± 0.0003 at Kondyor, between 0.1249 ± 0.0003 and 0.1250 ± 0.0003 at Inagli and between 0.1246 ± 0.0003 and 0.1260 ± 0.0003 at Guli. In a subsequent study, Malitch and Badanina (2011) used LAMC-ICP-MS to analyze another set of Os–Ir alloy nuggets from Guli and obtained 187Os/188Os values between 0.12433 ± 0.00010 and 0.12472 ± 0.00034. The latter authors also analyzed several laurite grains from the Guli placer deposits that have 187Os/188Os compositions that range between 0.12432 ± 0.00029 and 0.12463 ± 0.00009. Origins of PGM in placers associated with CUAAC. As noted by Weiser (2002) ‘…the complete history of the platinum-bearing placer deposits and their origin is not always known with certainty’. Evidently this statement is still valid, given the lively debate on the origin of PGM found in placers associated with CUAAC. Hattori and Cabri (1992) reviewed the ideas on the origin of PGM in placers and stated that by the mid-1960s, the prevailing view was that these minerals were originally present in ultramafic rocks and subsequently released to the placers by weathering and transportation processes. This hypothesis was largely based on studies of PGM from placer deposits of the Urals (Betekhtin 1961), Tulameen (Rice 1947) and Yubdo (Molly 1959). One of the well-accepted theories at the time was that large PGM placer nuggets were formed ‘by element agglutination under a phase of hydration (low temperature) of the ultrabasic rocks’ (Ottemann and Augustithis 1967). These ideas were later used to propose the hypothesis of ‘chemical accretion’ that was defended by several other authors (e.g., Cousins 1973; Cousins and Kinloch 1976; Stumpfl 1974, Cabri and Genkin 1991). The observation that PGM might crystallize in lateritic soils (e.g., Bowles 1986; Barker and Lamal 1989) led Cabri and Genkin (1991) to interpret the rounded shape and colloform habits exhibited by some PGM placer nuggets as evidence of the growth of PGM in a sedimentary environment. In their evaluation of the potential source of PGM in the placers of the Tulameen Complex, Nixon et al. (1990) concluded ‘…the mineralogical and compositional discrepancies between the PGM in chromitites and placers are perhaps not as striking as certain differences in grain size and texture…’. Hattori (2002) observed that the Os isotopic data of placer PGM are similar to primary mantle-derived PGM, and refuted their formation during sedimentary processes at low temperature. It was maintained that the isotopic data are not consistent with models invoking the dissolution and precipitation of the PGE in sediments and residual soils, as proposed by Bowles (1986). A primary origin for placer-hosted PGM, as proposed by Cabri and Genkin (1991), Cabri et al. (1996) and Nixon et al. (1990) was advocated instead. A significant body of data published subsequently demonstrate fairly convincingly that PGM found in lode mineralization and in placers share a common heritage (e.g., Razin 1976; Nixon et al. 1990; Malitch and Lopatin 1997; Malitch and Kostoyanov 1999; Garuti et al. 2002). However, it is worth noting that the grains analyzed by Hattori and co-workers were Os-rich PGM included within larger Pt–Fe alloys; significantly different to those PGM exhibiting low temperature growth habits interpreted to have been precipitated in the sedimentary environment.

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Garuti et al. (2002) studied the PGM assemblages of chromitites from Kytlym and Uktus and compared them with PGM from other Ural CUAAC placers, proposing that ‘…the paragenetic relationships of the Pt-alloys that we have observed in the chromitites of Kytlym and Uktus are practically the same as those reported from the alluvial nuggets of the Urals…’. Garuti et al. (2002) also stressed the difference in grain size between the PGM (< 100 µm) locked in chromitite and the alluvial nuggets reported from the Urals; the latter may be > 2 mm in diameter. Although a few placer-hosted Pt–Fe alloys associated with CUAAC have been reported to be relatively heavy (up to ~11 kg; Weiser 2002), Garuti et al. (2002) ruled out the hypothesis that these alloys precipitated or grew in the surficial environment by either physical or chemical agglutination (e.g., Ottemann and Augustithis 1967; Cabri and Genkin 1991). According to Garuti et al. (2002), the differences in PGM grain sizes are simply an artefact produced by several concomitant factors, including (1) inability of panning techniques to recover very small particles < 20–40 µm and (2) inherent difficulty liberating small (1–35 µm) grains from their chromite host during erosion of the bedrock or crushing during laboratory processing. Garuti et al. (2002) also noted that PGM with grain sizes from several tens of µm to a few mm occur in chromite interstices or associated with interstitial silicates in massive chromitites; such grains are comparable to those recovered as nuggets in placers. This is consistent with the observation of Augé et al. (2005), who described euhedral chromite grains totally surrounded by interstitial platinum in a PGM concentrate obtained from placer samples of Nizhny Tagil. Garuti et al. (2002) concluded that large PGM in lode chromitites could be more frequent than indicated by the study of polished sections and are therefore the best source material candidates for placer nuggets. The above arguments notwithstanding, many placer-hosted PGM show coatings of other PGM ± lower-temperature silicates, indicating that a solely residual origin for placer PGM (as grains liberated from ultramafic source rocks and later concentrated by mechanical processes) is not completely sufficient to explain all natural occurrences (e.g., Shcheka and Lehmann 2007).

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Molly EW (1959) Platinum deposits in Ethiopia. Econ Geol 54:467–477 Nakawaga M, Franco H (1997) Placer Os–Ir–Ru alloys and sulfides: indicators of sulfur fugacity in an ophiolite? Can Mineral 35:1441–1452 Nixon GT, Cabri LJ, Laflamme JHG (1990) Platinum-group element mineralization in lode and placer deposits associated with the Tulameen Alaskan-type complex, British Columbia. Can Mineral 28:503–535 Okrugin AV (2011) Origin of platinum-group minerals in mafic and ultramafic rocks: from dispersed elements to nuggets. Can Mineral 49:1397–1412 Ottemann J, Augustithis SS (1967) Geochemistry and origin of ‘platinum-nuggets’ in lateritic covers from ultrabasic rocks and birbirites of W. Ethiopia. Miner Deposita 1:269–277 Pearson DG, Parman SW, Nowell GM (2007) A link between large mantle melting events and continent growth seen in osmium isotopes. Nature 449:202–205 Razin LV (1976) Geologic and genetic features of forsterite dunites and their platinum-group mineralization. Econ Geol 71:1371–1376 Rice HMA (1947) Geology and mineral deposits of the Princeton map-area, British Columbia. Geol Surv Can Mem 243 Shcheka GG, Lehmann B (2007) Gold overprint of PGE alloy: an example from the Fadeevka Au–PGE placer, Russian Far East. Miner Petrol 89:275–282 Stumpfl EF (1974) The genesis of platinum deposits: further thoughts. Mineral Sci Eng 6:120–141 Tolstykh ND, Sidorov EG, Kozlov AP (2004) Platinum-group minerals in lode and placer deposits associated with the Ural–Alaskan-type Gal’Moenan Complex, Koryak–Kamchatka Platinum Belt, Russia. Can Mineral 42:619–630 Tolstykh N, Sidorov EG, Krivenko AP (2005) Platinum-group element placers associated with Ural– Alaska type complexes. In: Exploration for Platinum-Group Element Deposits. Mungall, JE (ed) Mineral Assoc Can Short Course 35:113–143 Tolstykh ND, Sidorov E, Kozlov A (2009) Platinum-group minerals from the Olkhovaya-1 placers related to the Karaginsky ophiolite complex, Kamchatskiy Mys Peninsula, Russia. Can Mineral 47:1057–1074 Tsintsov Z (2000) Platinum-group minerals in sediments from Gotse Delchev graben, SW Bulgaria. Compt Rend Acd Bulg Sci 53:73–76 Tsintsov Z (2003) Platinum-Group Minerals (PGM) from the alluvial sediments of Samokov region, West Bulgaria. Rev Bulg Geol Society 64 Tsintsov Z (2004) Sperrylite from alluvial placers of Vurbitsa River, SE Rhodopes. Bulgarian Geological Society, Annual Scientific Conference Geology, 16–17.12.2004, p 92–94 Tsintsov Z, Damyanov Z (1994) Sperrylite from Struma River alluvial placers, Blagoevgrad graben, SE Bulgaria. N Jb Miner Mh 11:518–528 Walker RJ, Morgan JW, Beary ES, Smoliar MI, Czamanske GK, Horan MF (1997) Applications of the 190Pt–186Os isotope system to geochemistry and cosmochemistry. Geochim Cosmochim Acta 61:4799–4807 Walker RJ, Brandon AD, Bird JM, Piccoli PM, McDonough WF, Ash RD (2005) 187Os–186Os systematics of Os–Ir–Ru alloy grains from southwestern Oregon. Eart Planet Sc Lett 230:211–226 Weiser TW, Bachmann HG (1999) Platinum-group minerals from the Aikora rivers area, Papua New Guinea. Can Mineral 37:1131–1145 Weiser TW (2002) Platinum-group minerals (PGM) in placer deposits. In: Cabri LJ (ed) The Geology, Geochemistry, Mineralogy and Mineral Beneficiation of Platinum-Group Elements. Can Inst Min Metall Petrol Spec Publ Vol 54, p 721–756

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Reviews in Mineralogy & Geochemistry Vol. 81 pp. 579-649, 2016 Copyright © Mineralogical Society of America

Mantle Sulfides and their Role in Re–Os and Pb Isotope Geochronology Jason Harvey Institute of Geophysics and Tectonics School of Earth and Environment University of Leeds Leeds, LS2 9JT United Kingdom

Jessica M. Warren Geological and Environmental Sciences Stanford School of Earth Science Stanford University Stanford, CA 94305-2115 USA

Steven B. Shirey Department of Terrestrial Magnetism Carnegie Institution of Washington Washington DC 20015-1305 USA INTRODUCTION Mantle sulfides (Fe–Ni–Cu-rich base metal sulfides or BMS; Fig. 1) play a crucial role in the distribution of Re, Os, and Pb in mantle rocks and are thus fundamental to obtaining absolute ages by direct geochronology using the Re–Os and Pb–Pb isotope systems on mantle samples. Mantle samples exist as hundreds of exposures of peridotites, pyroxenites and diamonds, either brought to the surface as accidental xenoliths and xenocrysts during kimberlitic or alkali basaltic volcanism (for comprehensive reviews, see Pearson et al. 2014; Aulbach et al. 2016, this volume; Luguet and Reisberg 2016, this volume), or as orogenic, ophiolitic and abyssal peridotite obducted at convergent margins and drilled / dredged from oceanic basins (e.g., Bodinier and Godard 2014; Becker and Dale 2016, this volume). This chapter reviews the occurrence of BMS in mantle samples and the role that they play in controlling the Re–Os and Pb isotope systematics of the mantle. Included in this review is a discussion of the role BMS plays in recording the multiple depletion / enrichment / metasomatic events that the mantle has undergone and the preservation of chemical heterogeneities that are inherently created by these processes. Along with discussions of the utility of Re–Os and Pb isotope measurements, this review will also consider the potential pitfalls and some of the surprises that can arise when analyzing these BMS micro-phases. Specifically excluded from this review is the extensive literature on Re–Os and Pb for the geochronology of sulfide systems in magmatic ores. This study is another field entirely from the study of sulfides in their native mantle hosts because of the complicated magmatic concentration processes occurring at crustal levels. 1529-6466/16/0081-0010$10.00

http://dx.doi.org/10.2138/rmg.2016.81.10

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Figure 1. Backscattered electron and chemical maps of typical mantle BMS grains. (a) Enclosed; (b) interstitial BMS, both from Mt Gambier peridotites, SE Australia (Alard et al. 2002); BSE, backscattered electron; MSS, monosulfide solid solution; Pn, pentlandite; Cp-Icb, chalcopyrite–isocubanite. Grayscale indicates the relative abundance of a given element. Reproduced with permission of Elsevier BV from Alard O, Griffin WL, Pearson NJ, Lorand J-P, O’Reilly SY (2002). Earth and Planetary Science Letters 203:651–663.

This review represents the first time that BMS in mantle rocks and diamonds are discussed together for the purposes of age determination. The interested reader should also consult other articles that have discussed these topics individually (e.g., Shirey and Walker 1998; Burton et al. 1999, 2012; Pearson and Shirey 1999; Luguet et al. 2001, 2003, 2004, 2007, 2008; Richardson et al. 2001, 2004, 2009; Shirey et al. 2001, 2002, 2004a,b, 2013; Alard et al. 2002, 2005; Aulbach et al. 2004a,b, 2009a,b,c, 2011; Harvey et al. 2006, 2010, 2011; Pearson and Wittig 2008; Gurney et al. 2010; Shirey and Richardson 2011; Warren and Shirey 2012).

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Figure 1 (Cont’d). Backscattered electron and chemical maps of typical mantle BMS grains. (c) intergranular BMS from Montferrier, French Massif Central (Alard et al. 2002); (d) cpx-enclosed BMS from Montboissier, French Massif Central (Alard et al. 2002). BSE, backscattered electron; MSS, monosulfide solid solution; Pn, pentlandite; Cp-Icb, chalcopyrite–isocubanite. Grayscale indicate the relative abundance of a given element. Reproduced with permission of Elsevier BV from Alard O, Griffin WL, Pearson NJ, Lorand J-P, O’Reilly SY (2002) Earth and Planetary Science Letters 203:651–663.

BACKGROUND Through the 1980s, much of the information obtained regarding mantle composition and its inherent heterogeneity was derived through indirect evidence, i.e., the study of basaltic volcanism and the interpretation of its isotopic signatures and trace element compositions (e.g., Zindler and Hart 1986). This method of studying inaccessible regions of the mantle relies heavily on the fundamental assumption that isotope ratios of a source peridotite are faithfully transferred to the resultant melt. However, isotopic studies of mantle rocks have shown that melts average out some of the variability and extreme depletions present in the mantle, especially when the melting region is larger than the scale of heterogeneity (e.g., Saal et al. 1998; Cipriani et al,. 2004; Harvey et al. 2006; Liu et al. 2008; Maclennan 2008; Warren et al. 2009; Stracke et al. 2011; Day 2013; Lassiter et al. 2014; Gannoun et al. 2016, this volume).

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The focus on basalts is not only because of a lack of mantle rocks exposed at the Earth’s surface. Problems arise in analyzing mantle samples because of the mobility during melting of the incompatible trace elements upon which early geochronological measurements relied (e.g., Rb–Sr, Sm–Nd, U–Pb, Th–Pb, Lu–Hf isotope systems). Mantle rocks exposed at the Earth’s surface typically undergo melting (i) as a result of adiabatic decompression, such as beneath a mid ocean ridge, continental rift, or in a mantle plume, and/or (ii) by lowering the solidus with CO2 or water such as when fluid fluxes into the mantle wedge at convergent margins, or carbonated peridotite generates carbonatite and kimberlite. In peridotites, melting is frequently overprinted by melt–rock interaction at a variety of scales and intensities (Warren 2015). Observations from seismic investigations, convection models, and basalt geochemistry also indicate that the mantle is heterogeneous due to prior episodes of melt depletion and enrichment (e.g., Stixrude and Lithgow-Bertelloni 2012; Dalton et al. 2014). The overprint of recent melt addition makes it difficult to “see through” metasomatic processes and determine the nature and, critically, the timing of melt-depletion in the mantle using lithophile elements. In contrast, the highly siderophile elements (HSE), which comprise the platinum group elements (PGE: Ir, Os, Ru, Rh, Pt, Pd) and Re, are compatible at a bulk-rock scale during melting (Re excepted) and have increasingly been used to study mantle composition and processes. Compared to the Earth’s core, HSE in the mantle are highly depleted (around 10 mg·g−1 versus 30 ng·g−1, respectively; Palme and O’Neill 2003; Lorand et al. 2008). However, these concentrations in the mantle are still unexpectedly high compared to those predicted by core–mantle separation models (e.g., Borisov et al. 1994). The most likely explanation for the elevated HSE concentrations in the Earth’s mantle is the “late veneer” hypothesis - an influx of meteorites impacting the Earth after core–mantle differentiation (Kimura et al. 1974; Chou 1978; Holzheid et al. 2000; Palme and O’Neill 2003; Lorand et al. 2008; Walker 2009). Two of the HSE, rhenium and osmium, comprise the Re–Os isotope system. Osmium remains in the mantle residue during bulk peridotite melting, while Re behaves incompatibly, so basaltic melts contain relatively little Os. Consequently, metasomatic and refertilization processes involving silicate melts can have little overall effect on the Os isotope ratio of residual mantle, which is fixed by its high Os content. As such, the Re–Os system in ultramafic rocks is generally regarded as the most favorable isotopic system for recovering model ages of mantle melting events (different types of model ages based upon Re–Os isotope systematics are discussed later in this chapter), even though they may contain inherent uncertainties of up to 300 Ma caused by heterogeneity or poor characterization of reference reservoirs such as the mantle (Carlson 2005; Rudnick and Walker 2009). As a result, over the past 26 years Os isotopes have been used routinely to date the melting events that an individual portion of the mantle has experienced, i.e., the transformation of fertile asthenospheric material into a depleted, buoyant lithosphere (e.g., Walker et al. 1989; Snow and Reisberg 1995; Pearson et al. 1995a; Reisberg and Lorand 1995; see also Luguet and Reisberg 2016, this volume, and Aulbach et al. 2016, this volume, for comprehensive reviews). Partitioning experiments (see Brenan et al. 2016, this volume, for details) have consistently demonstrated that PGE, including Os, on the whole behave as a group of strongly chalcophile and siderophile elements in mantle peridotites, being hosted—in the case of moderately melt-depleted ultramafic lithologies—entirely in accessory Fe–Ni–Cu BMS. As will be discussed in more detail later in this chapter, BMS therefore exert the main control on Re–Os isotope systematics in most ultramafic samples. The high KdOssulfide/silicate of ~ 105–106 therefore ensures that bulk-rock 187 Os/188Os is tracked by the behavior and abundance of BMS in the mantle. However, in highly depleted and/or altered peridotites, HSE are often hosted in refractory platinum-group minerals (PGM) and alloys (e.g., Luguet et al. 2007; Lorand et al. 2013) that have formed from BMS as a consequence of reactions involving melts or fluids. Thus, in strongly melt-depleted and/or altered peridotites, the influence of PGM on Os mass balance should not be ignored. Platinum-group minerals are discussed in detail in O’Driscoll and González-Jiménez (2016, this volume).

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In the mantle, BMS grains (Fig. 1) are usually present as trace or ultra-trace phases (≪ 0.1 modal %; e.g., Luguet et al. 2001). They have been observed as intergranular grains and as inclusions in silicate minerals in spinel lherzolite and spinel harzburgite xenoliths (Fig. 2; White 1966; Vakhrutshev and Prokoptsev 1972; MacRae 1979; Pasteris 1981; Amundsen 1987), garnet peridotite (e.g., Bishop et al. 1975; Meyer and Tsai 1979) and pyroxenite xenoliths (e.g., O’Reilly and Griffin 1987; Wilshire et al. 1988). Similar BMS grains are also found in abyssal peridotites (Fig. 3; e.g., Luguet et al. 2001, 2003; Alard et al. 2005; Harvey et al. 2006; Warren and Shirey 2012; Marchesi et al. 2013), orogenic peridotities (e.g., Reisberg et al. 1991; Luguet et al. 2007; Lorand et al. 2008; van Acken et al. 2010a), and ophiolites (e.g., Luguet et al. 2004). In diamonds they are recognized as a common group of mineral inclusions, although 99% of gem-quality macroscopic diamonds are inclusion-free (Stachel and Harris 2008). In the last 15 years, the accurate and precise measurement of Re–Os isotope systematics in individual BMS grains has become possible (e.g., Pearson et al. 1998; Burton et al. 1999; Richardson et al. 2001; Shirey et al. 2001; Griffin et al. 2002). The utility of Re–Os measurements for individual BMS grains has become increasingly apparent in light of growing evidence for

Figure 2. (a) Back scattered electron image of Type-1 BMS enclosed within host olivine grain from a lherzolite xenolith (Harvey et al. 2010). Several tiny BMS grains to the lower left of the main sulfide may be evidence for decrepitation resulting from rapid decompression as host xenolith is brought to the surface. Scale bar 100 µm. (b) Reflected light image of a trail of sulfide inclusions and silicate melt inclusions, precipitated from a C–O–S–H-rich fluid, spanning a fractured olivine grain that was subsequently annealed. Scale bar 100 µm. (c) Transmitted light image of a Type-1 enclosed BMS in a clinopyroxene megacryst, with clear evidence of decrepitation evidenced by trails of minute BMS grains leading away from the main sulfide. Field of view 400 µm. (d) Transmitted light image of a clinopyroxene megacryst containing “exploded” BMS grain surrounded by a cloud of minute BMS droplets which escaped from the main sulfide through decrepitation and decompression (Andersen et al. 1987). Field of view 300 µm. (a) and (b) reproduced from Harvey J, Gannoun A, Burton KW, Schiano P, Rogers NW, Alard O (2010) Geochimica et Cosmochimica Acta 74:293–320, (c) and (d) from Andersen T, Griffin WL, O’Reilly SY (1987) Lithos 20:279–294 with premission of Elsevier BV.

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(a)

(b)

(c)

(d)

Figure 3. Backscattered electron images of clinopyroxene-spinel-metasomatic BMS associations in abyssal peridotite. (a) Large metasomatic BMS grain (light gray) interlocked with magmatic clinopyroxene (cpx) surrounding a large orthopyroxene porphyroclast (opx) from the Kane Fracture Zone, Mid-Atlantic Ridge. (b) Clinopyroxene (cpx), with strongly arcuate and curvilinear grain margins, intergrown with residual orthopyroxene. A secondary origin for the clinopyroxene is suggested by the non-equilibrium dihedral angles between clinopyroxene grains, the common association of clinopyroxene with secondary BMS grains (S, circled), and the interstitial location of clinopyroxene. From the 15° 20’ N Fracture Zone, Mid-Atlantic Ridge. Scale bar 200 µm. (c) Base metal sulfide (arrow 1) along the margins of coarse clinopyroxene, with stringer of metasomatic BMS in close association with spinel, from the 15° 20’ N Fracture Zone, Mid-Atlantic Ridge. Scale bar 500 µm. (d) Intergranular BMS at boundary between olivine and pyroxene, and as inclusions in pyroxene from the SW Indian Ridge. (a) reproduced from Luguet A, Lorand J-P, Seyler M (2003) Geochimica et Cosmochimica Acta 67:1553–1570 and (d) from Seyler M, Lorand J-P, Toplis MJ, Godard G (2004) Geology 34:301–304 with permission of Springer-Verlag.

the disturbance of bulk-rock HSE systematics by C–O–H-bearing fluids (e.g., Alard et al. 2011) and sulfur-saturated silicate melts (e.g., Harvey et al. 2015, and references therein). Distinct generations of PGM and BMS grains have different Re–Os isotopic signatures, thus leading to the observation that the mantle has a heterogeneous isotopic composition at various length scales (Rehkämper et al. 1999; Griffin et al. 2002; Alard et al. 2005; Beyer et al. 2006; Frei et al. 2006; Harvey et al. 2006). Moreover, in the last few years, combined Pb–Pb and Re–Os isotope studies of peridotite-hosted sulfides have shed new light on mantle geochronology (Burton et al. 2012; Warren and Shirey 2012). Whole-rock Re–Os analyses of mantle-derived peridotites almost inevitably reflect the mixing of different generations of BMS because they have experienced multiple generations of

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processes. The movement of Os during melting, recrystallization, or metasomatism may occur independently of Re and there can be separate addition of Re (e.g., Alard et al. 2002; Griffin et al. 2002, 2004), making the interpretation of bulk-rock Re–Os isotope data ambiguous in some samples, in the context of the timing of melt depletion (see reviews in Reisberg et al. 2004; Carlson 2005; Rudnick and Walker 2009). For example, in the study of BMS inclusions in the Western Gneiss region of Norway, Beyer et al. (2004) identified at least two generations of BMS with model ages that ranged from 4.0 to < 0 Ga. Similarly, in the study of Marchesi et al. (2010) of BMS grains from the Ronda peridotite, Spain, multiple generations of BMS were detected at the thin section scale, implying that the whole-rock Os isotope composition and model ages of these samples are controlled by the modal abundance and Os contents of distinct generations of BMS. Such variability at the mineral scale overprints the inherent heterogeneity known to exist in mantle reservoirs fostering concerns that the mantle is adequately homogeneous in terms of 187Os/188Os to be described by large-scale geochemical reservoir systematics (Meibom et al. 2002). Although unradiogenic 187Os/188Os in peridotitic residues often can be adequately explained by ancient melt depletion, using bulk-rock Re–Os isotope systematics indiscriminately can have its drawbacks (Reisberg et al. 2004; Rudnick and Walker 2009). Like the standard caveats in other radiogenic isotopic systems, the Re–Os geochronology of bulk-rock peridotites has two built-in weaknesses: (i) mixing ages can be generated that are not geologically meaningful model ages, and (ii) potential parent (Re)–daughter (Os) mobility either directly or in sulfide liquids at deep-mantle temperatures lead to problems in interpreting Re–Os isotopic composition of sulfides as an absolute age. The utility of the Re–Os isotope system as a geochronometer thus lies in the ability of BMS to retain an Os isotope ratio that may have been generated billions of years ago. Despite the low closure temperature for Re–Os (e.g., ≤ 300 °C for pyrrhotite; Brenan et al. 2000), the low diffusivity of Os in silicate materials (Behrens et al. 1990; Cherniak 1995; Ganguly et al. 1998a,b; Burton et al. 1999) means that once Os is contained in BMS, it is difficult to equilibrate the Os isotope ratio of a BMS grain with surrounding silicates. For example, for a BMS inclusion in olivine, if the KdOs sulfide/olivine is 1, then at 1200 °C DFe–Mg in olivine is about 10−17 m2s−1 (Chakraborty 1997) and the diffusion distance in 106 years is about 1.7 m (using the approximation x2 ~ Dt). However, using a Kd of 106 for sulfide/olivine; i.e., a more realistic partition coefficient as derived from experimental data and measurement of natural samples, the diffusional flux drops by 4 orders of magnitude about 10−21 m2 s−1, and the diffusion distance is reduced to less than 0.5 mm. This makes the retention of ancient 187Os/188Os in a BMS grain trapped within a silicate much more plausible, irrespective of re-equilibration of silicates with metasomatic events subsequent to BMS grain isolation.

ANALYTICAL METHODS AND PRACTICAL ASPECTS OF SAMPLE PREPARATION Over the last 15 years or so, two main methods have been developed for the analysis of Re–Os isotopes in BMS grains; (i) in situ laser ablation multi-collector inductively coupled plasma mass spectrometry (LA MC ICPMS) and (ii) sulfide extraction, chemical purification, and thermal ionization mass spectrometry (TIMS). Lead isotopes have also been measured using the TIMS technique, as well as by in situ secondary ion mass spectrometry (SIMS). There are distinct advantages and disadvantages associated with each analytical method, particularly with respect to sample preparation and interpretation of the information obtained. Early work on Pb, Re, Os, and other PGE concentrations in BMS grains was performed by in situ proton microprobe analysis coupled with a MC ICPMS (Bulanova et al. 1996; Guo et al. 1999). Due to limited availability and because detection limits were higher and standard reproducibility for this method was less precise, in situ determination of Re and PGE

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concentrations are now routinely performed using LA ICPMS (e.g., Alard et al. 2000, 2002 2005; Lorand and Alard 2001; Luguet et al. 2001; Pearson et al. 2002). For Os isotope ratios, in situ LA MC ICPMS typically employs a Nd:YAG 266-nm or 213-nm laser microprobe coupled to the mass spectrometer (e.g., Pearson et al. 2002; Nowell et al. 2008). The new generation of Excimer lasers are also eminently suitable for this purpose and are routinely employed for PGE analysis in metals, BMS grains and PGM (e.g., van Acken et al. 2012). The spatial resolution achievable by LA MC ICPMS and SIMS, with analytical spot sizes that are typically as small as 50 µm, means that multiple analyses can be made on large (> 100 µm) BMS grains, allowing within-grain heterogeneity to be assessed and any contributions from exsolved phases to be examined. Also, it is possible to raster across large BMS grains to search for variability across a larger area. For smaller grains, the entire grain can be analyzed by increasing the beam diameter such that the whole grain and even part of the surrounding host phase are incorporated in the analysis, reducing the likelihood of skewing of data by omitting sub-grain exsolution products. As silicate or oxide phases are unlikely to contain significant amounts of Re or Os, letting the laser overlap small amounts of surrounding silicates in order to capture the entire volume of the sulfide is unlikely to result in any significant interference on the analysis of the sulfide itself. The ability to sample grains of a wide range of sizes is only limited by the amount of Re and / or Os contained in the grain and the resulting analytical precision that will be achieved. Moreover, after ablation, samples can be polished down further to reveal more material for subsequent ablations, as long as the sulfide is large enough. Hence, a benefit of in situ analysis is that internal textures of BMS grains can be assessed prior to ablation and that profiles through individual grains are possible. For the analysis of Re–Os in BMS grains by LA MC ICPMS, a mass bias correction is made by bleeding a dry aerosol of Ir into the gas line between the ablation cell and the ICPMS. Although the values obtained by plasma machines may drift over the course of a day, this can easily be corrected for by analyzing a sulfide standard at regular intervals during the analytical session. Typically, instrument induced drift is limited to 1–2 % over a 24-hour period (Pearson et al. 2002). However, there are several limitations that have been highlighted regarding the analysis of BMS grains using LA MC ICPMS. The overlap of 187Re on 187Os needs to be corrected for by measuring 185Re and using the natural ratio of 187Re/185Re of 1.6742. This isobaric interference can only be corrected for accurately when the 187Re/188Os is below a certain level. There has been some debate as to exactly what the limit should be for precise 187Os/188Os determination. The early study of Pearson et al. (2002) defined the limit of 187Re/188Os as < 1.2 but this was later refined by Nowell et al. (2008), who suggested that the much lower 187 Re/188Os value of < 0.5 is more appropriate to ensure an accurate correction of the isobaric overlap. Pragmatically, this 187Re/188Os upper limit rules out the analysis of eclogitic sulfides that have been in equilibrium with basaltic and more evolved liquids—leaving peridotitic sulfides to be tractable. LA MC ICPMS only allows the determination of mass ratios, and not Re and Os concentrations, as internal normalization is not possible. Semi-quantitative Os contents in BMS grains can be obtained indirectly by comparing the signal intensity of an unknown to that obtained on the standard that is analyzed throughout the analytical session. There are, however, several potential limitations on the accuracy of these data. For example, it is extremely difficult to manufacture completely homogeneous sulfide standards for LA MC ICPMS. Nanoclusters of PGE alloy may form when the bead is not quenched instantly, resulting in the possibility of tiny, but significant, heterogeneities on an analysis-to-analysis basis. Ablation conditions may also vary. This method cannot make use of an internal standard and optimum signals may not always be obtained when analyzing particularly small unknowns. In BMS grains of sufficient size, it may be possible to ablate a second hole, next to that ablated

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for isotopic analysis, for concentration determinations. However, this is only advisable when a large enough area of homogeneous sulfide is available for both measurements. Where in situ methods have a distinct advantages over TIMS analyses is the speed of analysis. A typical analysis of a BMS grain, including laser warm up time, the ablation of the sulfide itself and then wash out to return to the instrumental baseline often takes less than 15 minutes. Sample preparation time is, in general, low compared to the amount of chemical purification necessary for a TIMS analysis, and is limited to the amount of time necessary to cut and polish a thick (ca. 250 µm) section, inspect the section under microscope, and to locate the BMS grains. Significant sample preparation time is also saved as a single polished section may contain several 10s of grains that would otherwise have to be prepared separately for analysis by TIMS. However, care must be taken in the preparation of sulfides for laser ablation analysis as it is possible to inadvertently pluck some phases from the BMS grains. For example, removal of exsolved chalcopyrite will skew measured Re/Os ratios to lower values, as chalcopyrite may preferentially partition Re (Richardson et al. 2001). Nonetheless, several hundred sulfides can be analyzed in the same time it takes to prepare and analyze a single BMS grain by TIMS. Textural and mineralogical information relating to a BMS grain and its surroundings is preserved using in situ measurements. This can be particularly useful in immediately identifying BMS that are enclosed versus interstitial relative to the silicate phases. However, the work of Alard et al. (2002) demonstrated that the simplistic distinction between “enclosed” versus “interstitial” as corresponding to “primary” versus “metasomatic” BMS grains is not appropriate in many cases. The large datasets obtainable by LA MC ICPMS often have to be heavily filtered for sufficiently low Re/Os ratios and for high enough Os contents to give precise isotopic analyses, meaning that, in some extreme examples, up to 90% of a dataset needs to be cut because the 187Re/187Os isobaric interferences cannot be resolved. If bulk-rock Re–Os ratios are also available, then it is advisable to assess the remaining BMS data points in the context of the data on the bulk rock. One of the benefits of being able to rapidly generate large datasets is that probability density plots of Os isotope ratios and model ages (see section on the utility of sulfide Re–Os and Pb isotope geochronology below) can be generated. The large number of analyses required to produce meaningful probability density plots often, but not always, precludes this type of study using TIMS analyses alone (cf. Pearson and Wittig 2008). The TIMS method is conceptually much simpler than LA MC ICPMS—simply physically remove the sulfide grain by picking or drilling it out from surrounding silicate minerals. There also are some distinct advantages of analyzing individual BMS grains by TIMS. The level of precision and accuracy attainable for Re–Os–Pb isotope analysis is substantially higher, allowing precise measurements at the picogram level for Os and Re and at the tens of picogram level for Pb. This level of sensitivity makes TIMS the more desirable method for the analysis of Re–Os isotope systematics of diamond-hosted BMS, which tend to be very small. By adding enriched spike solutions (185Re 190Os 205Pb) to samples during their dissolution, precise elemental abundances can be obtained in addition to isotope ratios. Without prior knowledge of the paragenesis of a sulfide grain, it can be difficult to accurately spike a sulfide and only an estimate can be made in advance. In practice, the accuracy of a TIMS measurement compensates for under- or over-spiking, except when either excess is catastrophically large. Weighing errors can artificially influence the measurement, particularly for smaller sulfide grains that may have a mass of only a few micrograms, but the use of mixed spike solutions minimizes uncertainty in the analysis. Sample preparation is by far the greatest drawback of TIMS analysis. Extraction of the sulfide from the host rock, dissolution and purification are time consuming. If the sulfide is being extracted from a polished block, then this represents an additional step in the sample

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preparation compared to LA MC ICPMS. As with the polishing of samples for laser ablation, the problem exists of incomplete extraction of grains. When a sulfide is extracted from a host diamond or a thick section, the hole can be inspected to check that no sulfide phase remains. However, if sulfides are handpicked from a peridotite crushate or magnetically separated from an aggregate, then this check is not possible and the possibility exists of fractionating Re from Os in an individual BMS grain and producing an artificially low Re/Os ratio (e.g., Harvey et al. 2010, 2011). The advantage of hand picking or magnetically separating sulfide grains is that no material is lost compared to polishing down through a sulfide to identify a large enough grain to analyse. However, individual grains could become broken and grains can be fractured and entire fragments of a sulfide lost. Handpicking of sulfides from a crushate may also result in the preferential selection of interstitial grains that are easier to liberate from a bulk rock crushate. For the analysis of diamond-hosted sulfides (Fig. 4), the distinction between an eclogitic or peridotitic origin for the sulfide is important because of the large Os concentration differences

(a)

(b)

rosette fracture

(c)

sulfide

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Figure 4. Eclogitic (a) and (b) and peridotitic (c) and (d) diamond-hosted BMS inclusions. (a) and (b) Back-scattered electron images of eclogitic BMS inclusions liberated from Jwaneng diamonds, Kaapvaal craton (scale bar in µm). Pyrrhotite (po), chalcopyrite (cp), and pentlandite (pn). (c) and (d) Visible light images of peridotitic BMS inclusions (~150 μm diam) surrounded by rosette fracture systems in a diamond recovered from Panda, Slave craton, Canada. Reproduced from Aulbach S, Stachel T, Creaser RA, Heaman LM, Shirey SB, Muehlenbachs K, Eichenberg D, Harris JW (2009) Lithos 112:747–757 with the permission of Elsevier BV and Gurney JJ, Helmstaedt HH, Richardson SH, Shirey SB (2010) Economic Geology 105:689–712, under the Fair Use provision.

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that can be up to 1000 times less in the eclogitic sulfides. This distinction is typically made by placing the sulfide on a conducting carbon sticky disk in a scanning electron microscope and analyzing the grain for Fe, Ni, Cu, Co, and S. Exsolution features and grain surfaces can also be examined to determine if multiple BMS phases are present (Fig. 4a-b) and to check if the grain has been fractured. This technique could also be applied to BMS grains separated from bulk rock by magnetic separation in order to determine BMS paragenesis, but the textural context that is retained using in situ methods is still lost through the disaggregation of the bulk rock. Two recent studies (Burton et al. 2012; Warren and Shirey 2012) have presented Pb isotope and concentration data for individual BMS grains from abyssal peridotites, measured using TIMS analytical methods. Burton et al. (2012) measured Pb compositions in sulfides from the same section of peridotite drill core that was previously analyzed by Harvey et al. (2006) for Re–Os compositions. Warren and Shirey (2012) measured Pb, Re, and Os isotopic compositions and concentrations in the same grain by extracting all three elements during the chemical purification procedure and employing spikes. Lead isotopic compositions have also been measured in situ by SIMS, using either a Sensitive HighResolution Ion Microprobe (Eldridge et al. 1991; Rudnick et al. 1993b) or a Cameca IMS 1280 ion microprobe (Blusztajn et al. 2014). The main drawbacks of this technique are that 204Pb concentrations are too low to be detected and Pb concentration data have large uncertainties due to the lack of low concentration sulfide standards.

BASE-METAL SULFIDE OCCURRENCE, MAJOR ELEMENT GEOCHEMISTRY AND PETROLOGY For nearly half a century, BMS have been identified in a number of mantle-derived materials, including: peridotite and pyroxenite xenoliths entrained in basaltic and kimberlitic rocks (Fig. 1; e.g., Meyer and Brookins 1971; Desborough and Czamanske 1973; Frick 1973; Vakhrushev and Sobolev 1973; Bishop et al. 1975; De Waal and Calk 1975; Meyer and Boctor 1975; Mitchell and Keays 1981; Lorand and Conquéré 1983; Dromgoole and Pasteris 1987; Fleet and Stone 1990; Szabó and Bodnar 1995), disaggregated clinopyroxene, garnet and diamonds preserved as megacrysts (Figs. 1 and 4; e.g., Peterson and Francis 1977; Gurney et al. 1984; Andersen et al. 1987; Fleet and Stone 1990; Rudnick et al. 1993b; Deines and Harris 1995; Bulanova et al. 1996); and tectonically exposed ultramafic sequences derived from the upper, mostly oceanic, mantle (Fig. 3; e.g., Lorand 1985, 1987; Luguet et al. 2004; Alard et al. 2005). The interest in sub-mm BMS in ultramafic material stems from early, broad scale investigations into the behavior of BMS under magmatic conditions and sulfide liquid immiscibility in mafic and ultramafic rocks (Figs. 5 and 6; Bell et al. 1964; Kullerud et al. 1969; Naldrett 1969; Skinner and Peck 1969; Craig 1973; Misra and Fleet 1973; Haughton et al 1974; Usselman 1975; Buchanan and Nolan 1979; Fleet and Pan 1994; Karup-Møller and Makovicky 1995). More recently, experimental investigations into the behavior of BMS and, in particular, the HSE and strongly chalcophile elements that they contain, have provided a more comprehensive overview over a wide range of mantle P–T–X conditions (e.g., Bockrath et al. 2004; Ballhaus et al 2006; Holzheid 2010; and, in particular, see the review of experimental methods of Brenan et al. 2016, this volume). The most impressive, and economically interesting examples of the redistribution of BMS through magmatic processes are the large Fe–Ni–Cu-sulfide ore deposits such as Sudbury in Canada, Norilsk in Russia, and the Duluth Complex in the USA. Highly siderophile and chalcophile element-bearing ore deposits are reviewed in this volume by Barnes and Ripley (2016). The major sulfide phases observed in mantle rocks are generally pentlandite ((Fe,Ni)9S8), chalcopyrite (CuFeS2) ± pyrrhotite (Fe1-xS) and, in isolated instances,

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ISS + S (liquid)

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ISS + py

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Figure 5. Phase relations in the central portion of the Cu–Fe–S system at 600 °C. Numbered circles represent the stoichiometric compositions for the minerals (1) pyrite, (2) troilite, (3) chalcopyrite, (4) talnakhite, (5) mooihoekite, (6) haycockite, (7) cubanite and (8) bornite. All phases and phase assemblages coexist with vapour. Modified after Cabri (1973).

S

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(Fe, Ni)S2 + L (Fe, Ni)S2 + (Ni, Fe)S2 +L

MSS + (Fe, Ni)S2

(Fe, Ni)S2 + (Ni, Fe)S2 + MSS MSS + (Ni, Fe)S2 (Fe, Ni)9S8 + MSS (Ni,Fe)33+xS2

(Fe, Ni)9S8 + MSS

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S + (Fe, Ni)9S8 + (Ni, Fe)3+xS2 (Fe, Ni)9S8

α+γ + FeS

Fe

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(Fe1-xS + (Ni, Fe)3+xS2

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Ni

Figure 6. 600 °C isothermal section of the condensed Fe–Ni–S system (modified after Naldrett et al. 1969). All phases are in equilibrium with vapor.

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bornite (Cu5FeS4), shown on phase diagrams for the systems Ni-Fe-S (Fig. 6) and CuFe-S (Fig. 5). These phases are interpreted as representing the low-temperature (< 300 °C) subsolidus assemblages formed first from monosulfide solid solution (MSS = Fe1–x,Ni1–x S; Fig. 6) and later from intermediate solid solution (ISS) of the Cu–Fe–S system (Fig. 5), and heazlewoodite (Ni3S2) solid solution (HzSS) of the Ni–Fe–S system (e.g., Cabri 1973; Lorand and Conquéré 1983; Dromgoole and Pasteris 1987; Szabó and Bodnar 1995; Guo et al. 1999; Fleet 2006; Lorand and Grégoire 2006). Bornite, if present, likely crystallizes from bornite solid solution (BoSS), which is shifted toward a higher Cu content (and higher metal/sulfur ratio) than chalcopyrite. At the ambient, high pressure and high temperature conditions of the mantle, BMS forms either a homogeneous sulfide melt (known as a matte if it is high abundance) or is only partially molten (Bockrath et al. 2004). Experimental work suggests that MSS is the main stable BMS phase in the uppermost convecting mantle (Bockrath et al. 2004). Independent of their tectonic origin, the BMS found in peridotites, i.e., MSS/pyrrhotite–pentlandite–chalcopyrite, are the result of fractional crystallization from a high temperature sulfide melt followed by sub-solidus re-equilibration. At 1192 ºC, MSS with a composition close to FeS crystallizes, leaving a Cu–Ni-rich residual sulfide liquid at 1192 °C. As temperature decreases towards 1000 ºC, the MSS field extends toward Cu and Ni-richer compositions. With continued cooling (between 800 and 900 ºC, the ISS and HzSS fractionally crystallize from the Cu–Ni-rich residual sulfide liquid. This is often manifested as an outer rim on BMS grains or sometimes as veinlets escaping intergranular BMS. Subsequently, at ca. 600 ºC, HzSS and the MSS may react and crystallize pentlandite, while chalcopyrite crystallizes from the ISS shortly afterwards at ca. 560 ºC. Below 300 ºC further exsolution into a Ni-poor MSS, Ni-rich MSS (> 30 wt% Ni) and pentlandite takes place. However, when MSS–pentlandite–chalcopyrite melts, it does so incongruently, producing a Ni–Cu-rich sulfide melt, leaving behind a Ni–Cu-poor sulfide. Peridotitic BMS represent ca. 0.1 vol% of the upper mantle, which corresponds to a S concentration of 250 ± 50 µg·g−1 (Lorand 1990; O’Neill 1991; Palme and O’Neill 2003), and therefore, typically, sulfides are exhausted from the mantle after 12–30% partial melting (Luguet et al. 2003; Lorand and Grégoire 2006; Alt et al. 2007). The extent of sulfide exhaustion depends on local sulfur abundance, pressure of melting (Mavrogenes and O’Neill 1999) and melt composition. These reactions with melting result in the formation of PGM that become the major carriers of the PGE, as discussed in detail in this volume by O’Driscoll and González-Jiménez (2016). During melting, Ir and Os are strongly partitioned into the MSS, whereas Cu, Pt and Pd are partitioned into the coexisting melt (Brenan et al. 2016 this volume). Monosulfide solid solution at 900–1100 ºC can accommodate > 104 µg·g−1 of Os, Ir, Ru, and Rh (Li et al. 1996; Alard et al. 2000; Brenan 2002; Mungall et al. 2005; Ballhaus et al. 2006), while pentlandite, chalcopyrite, and pyrrhotite can accommodate from 102–103 µg·g−1 of each PGE (e.g., Peach et al. 1990; Fleet et al. 1996; Fonseca et al. 2011). Alteration of primary BMS occurs by supergene weathering (Fig. 7), in the case of continentally derived non-cratonic peridotites and by hydrothermal alteration / seawater interaction, in the case of oceanic peridotites. These processes typically increase the number of sulfide minerals observed in ultramafic assemblages. For example, during serpentinization at low temperature (< 250 ºC) and reducing conditions, BMS can experience desulfurization or, conversely, metal enrichment, to S-poor BMS such as heazlewoodite (Ni3S2), jaipurite (CoS), violarite (FeNi2S4), millerite (NiS, covellite (CuS), digenite (Cu9S5) and ultimately native metals or alloys such as copper or awaruite (Ni2Fe–Ni3Fe; Lorand and Conquéré 1983; Lorand 1987; Luguet and Lorand 1998; Alt and Shanks 2003; Bach et al. 2004; Marchesi et al. 2013; Schwarzenbach et al. 2014; Foustoukos et al. 2015). Continental supergene weathering also results in a loss of sulfur but, under oxidizing weathering conditions, it tends to produce Fe oxy-hydroxides at the expense of primary sulfides (Fig. 7; Dromgoole and Pasteris 1987; Lorand et al. 2003a).

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Figure 7. Progressively weathered BMS grains recovered from a Kilbourne Hole (NM, USA) peridotite xenolith (Harvey, unpublished data). Supergene weathering, resulting from interaction of interstitial BMS with meteoric water become increasingly weathered to Fe-oxyhydroxides (Lorand et al. 2003a) from (a) to (e), as sulfide oxidizes to sulfate, leading to loss of sulfur and oxidation of Os.

Peridotite-hosted sulfides Mantle-derived BMS can be divided into those that are peridotite-hosted, pyroxenitehosted, and diamond-hosted. Most primary BMS grains found in mantle peridotites have a diameter of 10–50 µm (Figs. 2 and 3; e.g., Guo et al. 1999), although in exceptional cases this may extend to 200 µm for sulfides enclosed in silicate grains (e.g., Beyer et al. 2004) and up to 500 µm for sulfides in intergranular interstices (e.g., Gonzáles-Jiménez et al. 2014). Sulfide commonly resides as a discrete inclusion within a silicate grain, but rarely within spinel (Ferraris

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and Lorand 2008). Alard et al. (2000, 2002, 2011) and Lorand and Alard (2001) define Type-1 sulfides as sulfides enclosed in olivine, orthopyroxene and garnet that have MSS compositions with minor amounts of pentlandite and chalcopyrite. The Type-2 BMS of Alard et al. (2002) have compositions dominated by pentlandite and chalcopyrite, with little or no pyrrhotite or MSS. These BMS tend to be intergranular and restricted to grain boundaries, fractures and pockets interstitial to the silicate minerals, as well as metasomatic BMS encapsulated within secondary (i.e., melt-crystallized) clinopyroxene. Type-2 sulfides may be derived by crystallization from a melt or liquid that was trapped during partial melting (Frick 1973; Mitchell and Keays 1981; Hamlyn and Keays 1986; Dromgoole and Pasteris 1987; Szabó and Bodnar 1995; Fig. 1) or introduced after melt depletion as a metasomatic melt or volatile-rich fluid (e.g., Alard et al. 2002, 2011; Luguet et al. 2008; Harvey et al. 2011, 2015; Lorand and Luguet 2016, this volume). These sulfides are the most common BMS population in peridotites, whereas Type-1 sulfides are rare and generally restricted to peridotite xenoliths (e.g., Alard et al. 2002). The original Type-1 and Type-2 definitions of Alard et al. (2002) encompass not only the geochemical, but also the petrological and textural contexts of BMS grains. Historically, the sub-division into Type-1 versus Type-2 BMS has served the BMS community well, yet the last decade of investigations has highlighted increasing difficulties with dividing mantle BMS into only two categories. In some instances, a third population of BMS may be present where limited mixing of Types 1 and 2 has occurred, resulting in a hybrid sulfide population that shares characteristics of both of its precursors (e.g., Griffin et al. 2002; Wang et al. 2009; Harvey et al. 2010). As will be discussed in the following sections, simple discrimination between “primary” versus “metasomatic” and “enclosed” versus “interstitial” can be difficult. In off-craton xenoliths, sulfide precipitation has also been attributed to sulfidation reactions between S-rich fluids and olivine, or metals dissolved in highly alkaline, volatile-enriched melts. Sulfides related to sulfidation have been identified in xenolith suites from the Kerguelen archipelago (Lorand et al. 2004), the Western United States (Lee 2002), and the Massif Central, France (Lorand et al. 2003b) and Languedoc, France (Alard et al. 2011). A first order approximation of the nature of the BMS hosted in a particular peridotite, i.e., the balance of BMS composition resulting from melt depletion versus metasomatism, can be made based upon sulfide modal abundance. Sulfur, like Al, is moderately incompatible during partial melting (e.g., Burnham et al. 1998). A sulfide mode of 0.1 % is the approximate theoretical maximum that could be expected from primitive mantle, i.e., a portion of mantle that has not undergone any melt extraction. This is based upon a fertile mantle S abundance of 250 ± 50 µg·g−1 S (Lorand 1990; O’Neill 1991; Palme and O’Neill 2003) and assumes a metal:sulfur ratio of 2:1. Modal abundances of BMS can be determined by careful point counting (e.g., Luguet et al. 2001, 2003, 2004), which is often made difficult by the low abundance and heterogeneous distribution of BMS in peridotite. The preparation of several thin sections is often required in order to ensure a realistic estimate of BMS abundance has been achieved. Although this is not an infallible means for detecting peridotites that have experienced extensive metasomatism, sulfide modal abundances in excess of 0.1 % are unlikely to be the result of melt depletion alone and may signify melt addition. For example, Aulbach et al. (2004a) reported unusually abundant BMS in garnet spinel lherzolites from the Slave craton, which could only be attributed to a secondary influx of sulfur accompanying a silicate melt. This conclusion was also supported by silicate melt patches in nearby spinel lherzolites, half of which comprised aggregates of polygonal pentlandite. Wang et al. (2009) also reported high BMS modal abundances in Taiwanese spinel lherzolite xenoliths as being the result of extensive metasomatism. The modal BMS contents of their lherzolites from Penghu Island range from 0.1–1 modal % sulfide, far in excess of that preserved in peridotite that has only experienced melt depletion. Peridotite-hosted BMS are predominantly derived from a MSS precursor which was in equilibrium with a sulfide melt (e.g., Craig 1973). Its occurrence is commonly interpreted

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as representing a residuum after partial melting (Helz 1977; Lorand 1987). Lherzolitehosted BMS are typically polyphase assemblages of pentlandite ± chalcopyrite ± Ni-poor MSS ± cubanite ± heazlewoodite ± millerite ± bornite. The experiments of Kullerud et al. (1969) in the Cu–Fe–Ni–S system demonstrated that Fe-rich MSS, with a composition close to FeS, crystallizes at 1192 ºC, i.e., at the temperatures at which most non-cratonic peridotite xenoliths equilibrated ( 7.5 wt% and up to 41 wt%, clearly in excess of the Cu solubility limit in residual MSS, indicating that these BMS grains are not residual MSS but, in fact, ISS. The cut-off of 7.5 wt% Cu serves as a convenient geochemical indicator for the distinction between ISS and MSS. In this instance, it is probable that the ISS exsolved from Cu-rich sulfide melts (e.g., Lorand et al. 2003b; Wang et al. 2009). Ni-rich (> 46 wt% Ni) BMS in these same samples exceed the solubility limit of Ni in residual MSS at 1000 ºC. One possibility suggested by Wang et al. (2009) is that these high-Ni BMS crystallized from sulfide melts at up to 1000 ºC (Ni < 54 wt%). The interpretation of BMS produced metasomatically through sulfidation reactions (Lorand et al. 2004; Delpech et al. 2012) is based on their morphology (high dihedral angle and sharp-straight BMS grain contours), textural relationships (associated with carbonates), and mineralogy (low Cu/Cu + Ni (< 0.1) and low-Ni, pyrrhotite-dominated assemblages; Lorand et al. 2013). These BMS are always derived from small fractions of alkali melts that have experienced extensive reaction with the host peridotite in the uppermost SCLM (Menzies and Hawkesworth 1987; Bedini et al. 1997; Alard et al. 2011). Such melts have the potential to be particularly effective metasomatic agents as they are enriched in C–O–H–S and remain liquid, and therefore mobile, to low temperatures (Fig. 8; e.g., Menzies and Dupuy 1991; Moine et al. 2004). In addition to the compositional heterogeneity preserved in BMS as a result of magmatic or metasomatic processes, many intergranular BMS grains preserve varying intensities of dendritic veinlets of Fe-oxyhydroxides (Fig. 7), or can be partially or completely replaced by magnetite. Both of these features are commonly interpreted as resulting from the oxidation of sulfide at sub-

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(b)

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Figure 8. C–O–S–H-rich melt inclusions in peridotite xenoliths from Mont Briançon, French Massif Central (modified after Harvey et al. 2010). (a) Plane polarized light image of a plane of melt inclusions trapped within an annealed olivine grain. Scale bar 20 µm. (b) CO2-fluid inclusion (top) with co-genetic, silicate melt inclusion beneath. The two phases are joined by a narrow neck ca.1 µm wide. Scale bar 10 µm (c) Co-genetic silicate and BMS grains preserved as immiscible liquids within the same inclusion. Scale bar 10 µm (d) CO2fluid inclusion with co-genetic sulfide bleb in olivine. Scale bar 5 µm. Reproduced from Harvey J, Gannoun A, Burton KW, Schiano P, Rogers NW, Alard O (2010) Geochimica et Cosmochimica Acta 74:293–320

magmatic temperatures (e.g., Luguet and Lorand 1998). Sulfide grains in mantle xenoliths may have been partially oxidized by fluids released from the host melt during degassing associated with eruption (e.g., Ryabchikov et al. 1995). Supergene weathering is another cause of oxidation of BMS to oxy-hydroxides (Fig. 7), which results in the selective removal of S in its more soluble sulfate form (Dreibus et al. 1995; Lorand et al. 2003a). This also provides a mechanism for at least partial oxidation of Os and its mobilization, resulting in sub-chondritic Os/Ir in many peridotite xenoliths (Handler and Bennett 1999; Lorand et al. 2003a; Pearson et al. 2004). Such oxidation may even extend to those BMS grains that resided at depth armored within a silicate host grain. For example, Lorand (1990) reported the loss of sulfur and Fe en route to the surface due to the fracturing of host olivine grains during decompression. Finally, metasomatic BMS grains that contain high amounts (> 4 wt%) of Si have been attributed to silicate melt inclusions (Aulbach et al. 2004b). Some sulfides also have unusually high-O abundances (up to 16 wt%), which has also been ascribed to oxidation and alteration of the sulfides (Aulbach et al. 2004b).

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Despite a narrow range of starting compositions, the large range of processes that can effect peridotite-hosted BMS results in a great potential for morphological and compositional heterogeneity. This variability is also expressed in the HSE and Re–Os isotope systematics of BMS, which may be a function of primary composition, cooling history, metasomatism and mixing of different populations of sulfides. Alard et al. (2002) described peridotite-hosted BMS as being either Type-1 (MMS enclosed within a silicate host), or Type-2 (dominated by Cu-rich pentlandite, commonly interstitial, but also occurring enclosed within metasomatic minerals such as secondary clinopyroxene) and resulting from either precipitation from a sulfur-saturated silicate melt or the introduction of a metasomatic sulfide melt. It is within this context that the Re–Os and Pb isotope systematics of peridotite-hosted BMS are now reviewed, using primarily the nomenclature of Alard et al. (2002). However, this nomenclature is modified as necessary in the examples that follow, where this division fails to adequately describe the petrogenesis of the BMS in question. The origin and nature of “Type-1” sulfides. Many mantle peridotites contain at least two generations of BMS (Alard et al. 2000; Fig. 1). As discussed in the previous section, Type-1 BMS are mixtures of MSS (with minor pentlandite and chalcopyrite), whereas Type-2 BMS comprise pentlandite + chalcopyrite ± pyrrhotite that exsolved from the higher temperature MSS or due to reaction between MSS and HzSS. The two sulfide types have very different Re and Os contents, and consequently Re/Os and 187Os/188Os, which in turn have implications for their utility in Re–Os geochronology. The underlying assumption regarding Type-1 BMS is that they formed during a melt depletion event. This produced a sulfide melt which was lost from the original peridotite and a residual Ni-rich MSS (Type-1 BMS), which was retained within the residual peridotite (e.g., Holzheid 2010 and references therein). Enclosed BMS can exhibit a wide range of shapes that range from polygonal cross-sections with straight contacts to spheroids, but all have habits consistent with entrapment. In harzburgites and dunites, the increase in olivine modal abundance that occurred during melt depletion (70–90 %) could have trapped pre-existing BMS (Lorand 1987). Osmium concentrations of Type-1 BMS typically range from 1–10s of µg·g−1 in noncratonic peridotites (e.g., Burton et al. 1999; Alard et al. 2000; Wang et al. 2009; Harvey et al. 2010, 2011) but this extends to 1000s of µg·g−1 in cratonic peridotites (e.g., Griffin et al. 2002; Aulbach et al. 2016, this volume), as summarized in Fig. 9; see also Supplementary Information. Type-1 BMS are Os-rich because their mineralogy is dominated by MSS. This is a function of the greater number of octahedral sites in Ni-rich MSS (Mackovicky et al. 1986; Cabri 1992; Ballhaus and Sylvester 2000), which are capable of accommodating Os, Ir, and Ru. Osmium isotope ratios (Fig. 10; see also Supplementary Information.) in undisturbed Type-1 BMS are typically less radiogenic than the upper limit estimated for Primitive Upper Mantle (PUM, 187Os/188Os ≤ 0.1296 ± 8, Meisel et al. 2001) or mean chondrite meterotites (187Os/188Os ≤ 0.127, Shirey and Walker 1998). Rhenium concentrations, on the other hand, tend to be much lower (< 1 µg·g−1 in many instances; Fig. 9) with typical 187Re/188Os for Type-1 BMS being equal to or lower than the upper value for chondrite meterotites (187Re/188Oschon ≤ 0.40186; Shirey and Walker 1998). The host phase for Type-1 BMS is most frequently olivine (e.g., Griffin et al. 2002; Aulbach et al. 2004a; Harvey et al. 2006), but occasionally orthopyroxene or garnet (e.g., Beyer et al. 2004). Base metal sulfide included in clinopyroxene with unradiogenic Os isotope ratios (e.g., Burton et al. 1999; Harvey et al. 2011) are likely to be Type-1 BMS, whereas BMS in clinopyroxene with radiogenic Os ratios (187Os/188Os > 0.1270, e.g., Warren and Shirey 2012) are probably Type-2 BMS. Base metal sulfides contained in clinopyroxene megacrysts and in pyroxenitic samples have very different Re/ Os and 187Os/188Os systematics, which are a consequence of a different petrogenesis. These are discussed in the following section along with the potential pitfalls of associating unradiogenic Os isotope ratios exclusively with Type-1 BMS enclosed within primary silicates.

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Os µg g-1

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peridotite / macrocrysts serpentinized peridotite diamond inclusions eclogite / pyroxenite

Figure 9. Rhenium and osmium concentrations in BMS grains hosted in abyssal peridotites, continentally derived peridotites and diamonds. See text and supplementary data for data sources.

10 10-1 10-3 10-3

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103 peridotite / macrocrysts serpentinized peridotite diamond inclusions eclogite / pyroxenite

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Figure 10. 187Re/188Os and 187Os/188Os in the same BMS grains from Figure 9. See text and supplementary data for data sources.

187

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The identification of unequivocally wholly enclosed BMS can be problematic. Minute (1000 µg·g−1 S), thus melting at depths where the S solubility of basaltic melt is relatively low. This melt would therefore be S-saturated, with the potential for precipitation of eclogitederived BMS when the pyroxenites crystallize. This would result in pyroxenite-hosted BMS with similar Re–Os isotope characteristics to MORB itself, notwithstanding any interaction with peridotite-hosted BMS during pyroxenite formation.

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The complexity of progressive melt infiltration, increasing melt/peridotite ratios, pyroxenite formation and sulfide compositional evolution was investigated in two suites of ultramafic rocks from the South China craton by Wang et al. (2009). They demonstrated that changes in HSE distribution and Re–Os isotope systematics of peridotite- and pyroxenite-hosted BMS could be accounted for by the progressive percolation of melts. In porphyroclastic and partly metasomatized peridotites, Wang et al. (2009) reported subchondritic to suprachondritic 187 Re/188Os and 187Os/188Os in BMS grains, which likely reflected fluid/melt interaction at low melt/rock ratios with residual MSS and/or crystallization of fractionated sulfide melts. This resulted in elevated Cu, Ru, Rh and Pd, and high Re/Os. At intermediate melt/rock ratios, peridotites displayed equigranular textures, were extensively metasomatized and contained sulfur-rich Ni–(Co)-rich MSS, with subchondritic to chondritic 187Os/188Os and subchondritic 187 Re/188Os. At the highest degrees of melt/rock interaction, i.e., where spinel-pyroxenite formed, pyroxenite-hosted BMS possessed the lowest HSE contents, had subchondritic to suprachondritic 187Os/188Os and subchondritic 187Re/188Os, and were interpreted as having precipitated from sulfide melts that segregated from basaltic melts under S-saturated conditions. With the exception of the subchondritic 187Re/188Os, this process is similar to that observed during the formation of dunites at high melt / rock ratios in the Troodos ophiolite by Büchl et al. (2002, 2004) and in Shetland, Scotland (O’Driscoll et al. 2012). In the Totalp Massif, Switzerland, in garnet–spinel-clinopyroxenites and spinelwebsterites contained within melt-infiltrated Jurassic ocean-floor peridotites, van Acken et al. (2008, 2010a,b) identified pyroxenite-hosted pentlandite and godlevskite ((Ni,Fe)8S9) as being the major hosts of HSE. In the host lherzolites, up to 20% of the observed BMS grains had been derived from a mafic melt, resulting in substantial enrichments of Re, Pd and Pt, whereas Ir, Os, and Ru remained largely unaffected. Heterogeneous HSE abundances within individual pyroxenite-hosted BMS were attributed to subsolidus processes, such as exsolution, not resolvable using LA ICPMS, due to small sub-grain sizes. Large grain-to-grain variations were likely caused by multiple events of melt–rock interaction and sulfide precipitation, which demonstrate the potentially complex Re–Os isotope systematics that may occur in pyroxenites and that bulk-rock Re–Os isotope systematics are unlikely to preserve geochronological information pertaining to a single incidence of melt–rock interaction. van Acken et al. (2010a) reported cm-scale heterogeneity amongst pyroxenite-hosted BMS, corresponding to both residual-peridotite and melt-like HSE signatures, suggesting that the prior Re–Os signature of the original peridotite may not have been completely erased. Over time, the enrichment in Pt and Re and depletion in Os of the Totalp pyroxenites would be expected to develop radiogenic 187 Os and 186Os, compared to ambient mantle. In a recent study of Horní Bory, Bohemian Massif, Ackerman et al. (2013) demonstrated that the relationship between bulk-rock Fe-wehrlites, wehrlite-clinopyroxenites, websterites, clinopyroxenite and olivine clinopyroxenite, and the BMS grains they contain are also convoluted and result from the interaction between a melt-depleted peridotite residue and a sulfur-undersaturated melt generated in a subduction setting. Here, the bulk-rock pyroxenites are depleted in Os, Ir, and Ru, enriched in Pd, Pt, and S, and have radiogenic 187Os/188Os. They interpret the origin of these pyroxenites as melt–rock reaction between peridotite and subduction-related melt at high melt/rock ratios, which resulted in primary BMS grains being replaced by radiogenic (high 187Os/188Os) but HSE-poor BMS. Bulk-rock Os abundances are highly variable (0.08–9.3 ng·g−1) but generally low (< 1 ng·g−1), while Re abundances are less variable (1–4 ng·g−1), corresponding to high 187Re/188Os (up to 198) and radiogenic 187Os/188Os (0.139–1.23). However, most of the pyroxenite-hosted BMS have exceptionally low Os and Re, typically below detection limits (< 0.01 µg·g−1), with the result that Ackerman et al. (2013) calculated that the PGE mass balance was not wholly dependent upon the BMS present. Only 28% of the pyroxenite HSE budget was hosted by BMS, with the remainder contained in

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micron-scale PGM. Despite the small-scale heterogeneity, wehrlite, dunite and pyroxenite rocks yielded a Re–Os errorchon age of 334 ± 19 Ma (MSWD = 1318), which is consistent with the age of other high temperature/pressure rocks from the Bohemian Massif, and likely represents the time of melt–rock reaction and subduction. However, it should be remembered that this “age” is derived from a complex mixture of PGM, secondary BMS, and pyroxenite melt, all assuming that the Re–Os characteristics of the replaced peridotite were completely erased. The three case studies just described represent the current state of the art with regard to what has been achieved with pyroxenite-hosted BMS. While the geochronology of these rocks must be interpreted with care, detailed grain-scale analyses have the potential to provide unique constraints on the timing and nature of melt–rock reaction events.

Diamond-hosted sulfides Around 1% of macrocrystalline, gem-quality diamonds contain inclusions large enough to identify with a hand lens or optical microscope. Of such inclusions, BMS grains are not uncommon (Harris 1992; Deines and Harris 1995). The association between diamond occurrences and cratonic mantle has been known since the 1960s (e.g., Clifford 1966; Gurney and Switzer 1973; Boyd and Gurney 1986; Janse 1992). As a result of this association, coupled with (i) cold geotherms derived from geothermobarometry on peridotite xenoliths that pass through the diamond stability field, (ii) the actual presence of diamonds in xenoliths, and (iii) inclusions in diamonds that yield Archean ages, the hypothesis that diamonds reside in ancient lithospheric keels beneath continental interiors has gained wide acceptance (e.g., Richardson et al. 1984; Boyd et al. 1985; Boyd and Gurney 1986; Haggerty 1986). Compared to their low abundance in the mantle in general, BMS inclusions in diamonds from some kimberlites are unusually common (Gurney 1989), making them of particular interest to those who are interested in diamond paragenesis (Meyer 1987) and absolute ages of diamond formation (Pearson and Shirey 1999; Gurney et al. 2010; Shirey and Shigley 2013; Shirey et al. 2013). Once enclosed in their chemically inert host, inclusions in diamonds are shielded from interaction with transiting melts and fluids, meaning that their chemical composition and the age of diamond formation can be preserved for > 3 Ga (Richardson et al. 1984, 1993, 2001; Smith et al. 1991). Just like the peridotite-hosted BMS discussed above, diamond-hosted sulfide mineralogy is dominated by pyrrhotite (Fe1-xS) and/ or pentlandite ((Fe,Ni)9S8), with lesser amounts of Cu- or Co-rich sulfides such as chalcopyrite (Deines and Harris 1995; Fig. 4). Diamond-hosted BMS have been interpreted as precipitating from a C-O-S-H-rich fluid associated with Archean mantle melting (Deines and Harris 1995) and suggested to be syngenetic with the host diamonds (Bulanova et al. 1996). However, recent work (e.g., Howarth et al. 2014a,b) has shown that diamonds can also be found in metasomatic veins in eclogite and bear no particular spatial association with BMS grains. Furthermore, some diamonds host multiple BMS inclusions that are not in sulfur isotope equilibrium (e.g., Thomassot et al. 2009). In these cases the sulfides must be protogenetic (growing before the diamond). Protogenetic versus syngenetic distinctions are mineralogical in nature and may not compromise the use of BMS for age determination if the age difference between protogenetic inclusion growth and host diamond growth is within the resolution of the age dating method. Nonetheless, the similarity of BMS inclusions to peridotite-hosted BMS is demonstrated by their identical experimentally determined paragenesis i.e., derivation from sulfide/silicate immiscibility in the FeS–FeO–Fe3O4–SiO2 system at around 1100 °C (Maclean 1969). Base metal sulfide inclusions are distinguished from silicate and oxide inclusions by the presence of large rosette fractures surrounding each sulfide (Fig. 4). The rosette fractures result from rapid decompression during transport of the host diamond to the surface, commonly in a kimberlitic eruption. Silicate inclusions in diamond can be subdivided into eclogitic (E-type diamonds) and peridotitic (P-type diamonds; see Meyer 1987; Stachel and Harris 2008), More

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recent work has added lherzolitic and wehrlitic parageneses to the classification scheme. Diamond-hosted BMS can be also be categorized by the presence or absence of peridotitic or eclogitic silicate and / or oxide mineral inclusions in the same diamond (e.g., Bulanova 1995), but the co-occurence of BMS and silicate inclusions in the same diamond is so rare that it is hardly a workable classification method for BMS inclusions. Sulfide composition can itself be used, but this classification can only be made after the inclusion is broken out of the diamond as the reflectivity and color differences between different BMS are hard to observe through the diamond. Major element discriminators for peridotitic versus eclogitic BMS inclusions, such as Ni wt%, were used early in the 1980s in Siberian diamonds (Yefimova et al. 1983). However, with increasing numbers of analyses, the distinction between these two groups of Siberian diamonds became less clear (Deines and Harris 1995; Bulanova et al. 1996). Some studies (e.g., Smit et al. 2010) have observed a pyroxenitic BMS paragenesis compositionally in between eclogitic and peridotitic, which perhaps may explain the less clear distinction for some BMS inclusions. Subsequent studies have found that a combination of Ni and Os abundances is a more robust discriminator (Pearson et al. 1998). Osmium abundances range by a factor of 103 between eclogitic and peridotitic BMS inclusions and provide a more sensitive means of discrimination. Although 12% Ni content was the original identifier for peridotitic BMS, defined by Yefimova et al. (1983) and further refined by Deines and Harris (1995) and Pearson et al. (1999a,b), this scheme of classification does not work as well for BMS inclusions in Siberian diamonds (Sobolev 1974; Gurney 1989). As discussed in the following section, eclogitic (as opposed to peridotitic) BMS were originally defined as containing < 8 wt% Ni (Yefimova et al. 1983), but the two-fold classification was found not to be appropriate for all sulfide inclusions. For example, BMS inclusions containing 11–18 wt % Ni in several Yakutian diamonds (Bulanova et al. 1996) left the provenance of Siberian diamonds in doubt. Consequently, a third sulfide-based paragenesis was proposed for BMS inclusions with intermediate Ni contents between 8 and 12%; a pyroxenitic paragenesis (Deines and Harris 1995; Pearson et al. 1998; Smit et al. 2010). As with peridotite-hosted BMS, the relatively high Re and Os concentrations in diamondhosted BMS mean that the Re–Os isotope system is ideal for the analysis of individual inclusions (Pearson et al. 1998; Pearson and Shirey 1999; Richardson et al. 2001). Single grain analysis of BMS has an advantage over the use of lithophile element-based isotope systems, e.g., Sm–Nd, in the analysis of silicate inclusions, as in these instances several inclusions may need to be pooled in order to provide enough material for an adequately precise measurement. However, the advent of a new generation of stable, high-ohmic input resistors on TIMS instruments is likely to change this story in the near future and allow single silicate grains to be analyzed. The strength of single BMS grain Re–Os measurements comes from the observation that not only can a single BMS grain give a model age for its formation, but also multiple nontouching sulfides with a range of Re/Os elemental ratios, closed to diffusive exchange between each other by the intervening inert diamond, have the potential to yield an isochron age if they grew during the same generation of diamond growth. Combined with the Sm–Nd system in lherzolitic and eclogitic garnet and clinopyroxene, for example (e.g., Richardson et al. 1990) or perhaps Pb–Pb if there is enough in BMS, it may be possible to generate complimentary age results by using two or more independent chronometers on silicate and BMS inclusions. Peridotitic sulfides. Peridotitic diamonds are identified mainly by the peridotite mineral inclusions that they contain. These include diopside, olivine, enstatite, chromite and Crpyrope, (Meyer and Boyd 1972; Gurney and Switzer 1973; Harris and Gurney 1979; Fig. 4). Peridotitic diamond paragenesis can also be subdivided into harzburgitic or lherzolitic on the basis of the melt depletion signature of the inclusions. For example, this may be determined by olivine and / or orthopyroxene composition, the presence or absence of diopside or the CaO and Cr2O3 content of included garnet (Sobolev et al. 1973; Harris and Gurney 1979). As discussed in the previous section, unequivocal diamond paragenesis becomes more problematic in the absence of silicate inclusions.

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Peridotitic BMS are likely to be residual after low to moderate degrees of partial melting (< 25%; Keays 1995), where the basalts produced will almost certainly be sulfur saturated at the source of melting. In contrast, at the higher degrees of partial melting (>25%) necessary for the production of picrites or komatiites, the exhaustion of sulfur at the source of melting is much more likely to occur, leaving little or no residual sulfur for the formation of BMS that could ultimately become enclosed in diamond (e.g., Keays 1995; Mavrogenes and O’Neill 1999). Therefore, highly depleted portions of the SCLM that are associated with komatiite formation seem to be unlikely hosts for peridotitic BMS inclusions (e.g., Aulbach et al. 2009a). Nickel concentrations of diamond-hosted BMS grains provide some insights into diamond paragenesis, in the absence of co-genetic silicate inclusions. A minimum, cut-off value of 12 wt% Ni for BMS was suggested to be appropriate for peridotitic diamond formation (Yefimova et al. 1983; Deines and Harris 1995; Pearson et al. 1999a,b). More recently, Aulbach et al. (2009a) used (Ni + Co)/Fe ratios for BMS, rather than absolute abundances, in order to determine diamond paragenesis. Rather than analyzing the sulfide major element abundances using in situ methods—and the possible difficulties in reconstructing sulfide major element geochemistry from a number of phases that exsolved from a high temperature MSS (cf. Griffin et al. 2002)—Richardson et al. (2001), Westerlund et al. (2006), and Aulbach et al. (2009b) obtained the sulfide major element abundances from the eluent produced during the purification procedures for Re–Os isotope analysis. As the major metals are removed in a single aliquot, their relative abundance can be determined by mixing the eluent with appropriate gravimetric standards and then analyzing by ICPMS. Peridotitic BMS share geochemical and isotopic similarities with Type-1 peridotite-hosted BMS, and in some respects can be thought of as extreme versions of those BMS found in noncratonic SCLM. Osmium concentrations tend to be high (up to µg·g−1 levels) compared to BMS derived from the other diamond parageneses (Figs. 9 and 11). Rhenium concentrations are low,

101

Re c lti

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sa

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ba

s de pl et = ed 0. S 01 C to LM 0. Re 03 -O

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-O

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Archean ( 3.5 Ga), late Archean (< 3.5 Ga) and Proterozoic. Fertile mantle abundances (PUM; Morgan 1986) are for whole-rock peridotites. Field labeled “abyssal peridotite” encloses data for BMS in present-day abyssal peridotites (from Harvey et al. 2006). Modified after Shirey and Richardson (2011).

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and with high Os concentrations lead to extremely unradiogenic 187Os/188Os ratios in BMS older than 2 Gyr (see below). Although the Re–Os isotope systematics of cratonic diamonds will be discussed in greater detail below (see section titled “The utility of Re–Os and Pb isotope geochronology”), typical Os abundances range from 10s–100s of µg·g−1 in peridotitic BMS from the Siberian craton, Russia (Griffin et al. 2002), and the Slave Craton, Canada (Westerlund et al. 2006; Aulbach et al. 2011, respectively). Exceptionally, several thousand µg·g−1 and even up to 2 × 104 µg·g−1 have been reported in one Siberian BMS grain (Griffin et al. 2002). The abundance of Re in Slave Craton peridotitc BMS grains tend to be < 1 µg·g−1 (e.g., Aulbach et al. 2011) and, while studies using LA MC ICPMS to obtain Os isotope ratios do not report Re abundances because of the isobaric interference on 187Os (e.g., Griffin et al. 2002, 2004; Beyer et al. 2004; González-Jiménez et al. 2014), 187Re/188Os ratios in peridotitic BMS are very low, i.e., significantly less than the mean for chondrites (187Re/188Os = 0.40186; Shirey and Walker 1998; see also Day et al. 2016, this volume), and often much lower, e.g., 187Re/188Os = 0.002 in sulfide LT98/10-5 from the Kaapvaal craton (Griffin et al. 2004) and DA36 peridotitic sulfide from Diavik, Slave craton, Canada (Aulbach et al. 2011). In general, undisturbed peridotitic BMS grains possess 187Os/188Os ratios that are less radiogenic than mean chondrite and PUM (i.e., 187 Os/188Os ≤ 0.1270 and 0.1296; Shirey and Walker 1998 and Meisel et al. 2001, respectively). A consequence of these sulfides evolving for several billion years, trapped as inclusions in diamond, and hence in a low-Re environment, is that peridotitic BMS tend to have very unradiogenic Os isotope ratios; typically 187Os/188Os < 0.110 in sulfides with the lowest 187Re/188Os. Eclogitic sulfides. Unlike the high degree of melt depletion associated with komatiite and picrite formation, subduction of oceanic crust provides more favorable conditions for growth of diamonds with BMS inclusions. Not only may there be an adequate supply of carbon, but also sulfur, as oceanic crust formation occurs under S-saturated conditions. In addition, the low Ni content of most eclogitic BMS leads to low Ni/Fe ratios (0.1–0.2), which are too low to be in equilibrium with mantle olivine and suggest a basaltic precursor (Deines and Harris 1995; Richardson et al 2001). Although a subduction origin for the carbon from which some diamonds form may be controversial (cf. Cartigny et al. 2001), a subduction-related origin for diamond formation and for the supply of material to Archean SCLM roots is independently supported by (i) both S and Pb isotope systematics of BMS grains in eclogitic diamonds (Eldridge et al. 1991; Rudnick et al. 1993b; Farquhar et al. 2002; Thomassot et al. 2009), (ii) by diamondbearing eclogite xenoliths whose oxygen isotopic systematics are not consistent with a mantle origin (e.g., Jacob 1994; Shirey et al. 2001, 2004a, 2008), and (iii) by the correlation between isotopically light C in diamond and the isotopically heavy O in eclogitic silicate inclusions (Ickert et al. 2013, 2015; Schulze et al. 2013). The S-saturated nature of the eclogitic protolith is less controversial and is based on the occurrence of BMS globules in MORB (e.g., Roy-Barman et al. 1998) and the absence of evidence for wholesale melting of the eclogitic protolith that would have removed the sulfur during subduction. A consequence of the mafic nature of the subducted material is that ultimately this will experience eclogite facies metamorphism, form eclogite and hence supply material that may become encapsulated within diamonds as eclogitic inclusions. Diamonds are classified as eclogitic when they contain pyrope–almandine garnet ± coesite ± kyanite ± omphacite or sulfide with low Ni contents. In contrast to the high Ni content of peridotitic BMS, eclogitic BMS grains are defined as containing predominantly Fe-sulfide, with less than 8 wt% Ni. Eclogitic BMS mineralogy is dominated by a low-Ni, pyrrhotite–pentlandite– chalcopyrite assemblage derived from cooling of MSS and ISS precursors. Pyrrhotite is usually the dominant phase, but the proportions of chalcopyrite and pentlandite may vary considerably within an individual sulfide. Chalcopyrite is concentrated on exterior surfaces of inclusions (Richardson et al. 2001), as a result of exsolution during kimberlite emplacement. Because Re is preferentially concentrated in chalcopyrite relative to pyrrhotite and/or pentlandite (Richardson et al. 2001; Brenan 2002; Brenan et al. 2016, this volume), incomplete recovery of a diamond

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enclosed inclusion may preferentially skew the Re–Os elemental ratio to lower or higher values than the true ratio. If no significant exsolution of chalcopyrite occurred, for example if the sulfide was molten during extrusion into the surrounding fractures, then this is unlikely to be a problem. However, the observation that chalcopyrite is preferentially concentrated around the edges of a wide range of sulfides suggests that great care should be taken to exclude exsolution as a potential source of bias from Re–Os geochronological measurements. This exsolution problem and the high Re content make these grains unsuitable for LA MC ICPMS. Rhenium–osmium isotope systematics of eclogitic BMS differ from those of peridotitic BMS. Osmium concentrations are often several orders of magnitude lower than peridotitic BMS, e.g., < 1.5 µg·g−1 (e.g., Pearson et al. 1998; Pearson and Shirey 1999; Richardson et al. 2001, 2004; Aulbach et al. 2009a), while rhenium contents are only moderately lower in eclogitic BMS than peridotitic BMS (≤ 2.2 µg·g−1; Aulbach et al. 2009a), as shown in Figure 9. Consequently Re/Os are also higher (eclogitic mostly > 1 versus peridotitic < 1; Pearson et al. 1998) and 187Re/188Os are also high (often > 100; e.g., Pearson et al. 1998, Aulbach et al. 2009a) and sometimes exceptionally high (> 1500; Richardson et al. 2004). Again, like pyroxenitic BMS, 187Os/188Os of eclogitic samples are all suprachondritic and often very high in old BMS grains, which contrasts with the dominantly unradiogenic Os isotope signature of undisturbed peridotitic BMS. This is predominantly because of the high 187Re/188Os and, especially in the case of diamond-hosted eclogitic BMS recovered from cratons, time integrated decay of 187 Re to 187Os. Interestingly, in a recent study of eclogitic BMS inclusions recovered from Yakutian diamonds in the Siberian craton (Wiggers de Vries et al. 2013), Re–Os isotope dating demonstrated that individual diamonds may have protracted timescales of growth, with vastly different ages for cores and rims. For example, eclogitic BMS recovered from individual growth zones in diamonds from Mir and 23rd Party Congress kimberlite pipes were demonstrated to have crystallized over intervals lasting < 200 Ma, but separated by up to ca. 1 Ga.

Re–Os–Pb MASS BALANCE IN ULTRAMAFIC SAMPLES Osmium mass balance The preferential partitioning of HSE into sulfide versus silicate was demonstrated experimentally in the early 1990s (e.g., Peach et al. 1990) and was well-known in ore-forming sulfide systems (see Shirey and Walker 1998 for an early review). But the investigation of HSE concentrations in natural sulfides in non-ore systems has lagged behind by several years. It was not until the investigation of Os abundances in handpicked mineral separates from a picrite from Mauna Loa, Hawaii, an ankaramite from Pico, Azores, and a Kilbourne Hole lherzolite (Hart and Ravizza 1996) that an Os mass balance was estimated for sulfides in equilibrium with silicate melt. This early study demonstrated that in the Kilbourne Hole peridotite, a mineral separate of BMS grains contained 3 orders of magnitude more Os than the bulk rock, and 4–5 orders of magnitude more Os than silicate and spinel mineral separates. Moreover, leaching an aliquot of the powdered lherzolite (in 2.5 M HCl for 1 hour) removed almost 75% of the Os (and all of the S) compared to an unleached aliquot, suggesting that the majority of the Os was hosted by sulfide. Leaching of a finely powdered bulk-rock lherzolite clearly does not provide direct constraints on sulfide Os compositions in the context of their Type-1 or Type-2 assemblage, but can serve as a rough guide to where the phases may be in the rock. In this study, the distribution of Os amongst the other constituent phases of the lherzolite was determined, resulting in calculated KdOssulfide/silicate (≤ 105) that were comparable to those determined experimentally. Working on a different Kilbourne Hole lherzolite, Burton et al. (1999) analyzed the Re–Os isotopic and elemental composition of inclusion-free silicates, spinel and BMS. By careful handpicking and sulfide selection, they were able to separate both enclosed and interstitial BMS and analyze them separately. In addition, they separated and analyzed clinopyroxene and

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orthopyroxene fractions that were visibly contaminated by sulfide inclusions. The inclusionfree silicate phases contained very small amounts of Os (104–136 pg·g−1), whereas spinel contained an order of magnitude more Os (1.2 ng·g−1). However, Type-1 and Type-2 BMS contained 3.9 and 3.7 µg·g−1 Os, respectively, unlike several later studies (Alard et al. 2002; Harvey et al. 2010, 2011) where Type-2 BMS were described as containing significantly less Os. Critically, Burton et al. (1999) reported that the sulfide-contaminated pyroxenes had significantly higher Os concentrations than the uncontaminated equivalent, with both orthopyroxene and clinopyroxene containing 1.9 ng·g−1 Os compared to ca. 130 pg·g−1 in the clean mineral separates. Moreover, the pyroxenes that were contaminated by sulfide inclusions yielded 187Os/188Os that were remarkably close to that of the enclosed BMS (included sulfide and contaminated orthopyroxene 187Os/188Os = 0.11719 ± 11; contaminated clinopyroxene 187 Os/188Os = 0.11790 ± 28), which were all significantly less radiogenic than the bulk-rock (187Os/188Os = 0.12076 ± 62 to 0.12126 ± 64), the interstitial BMS (187Os/188Os = 0.12086 ± 20), and the uncontaminated silicates (187Os/188Os = 0.12066 ± 11 to 0.12131 ± 22). For this particular Kilbourne Hole lherzolite, this not only confirmed that the Os elemental (and therefore 187Os/188Os) budget is controlled by BMS, but also suggested that the bulk-rock budget was likely dominated by Type-2 BMS in equilibrium with the silicates (Fig. 12) that

orthopyroxene

187Os/188Os

(a)

clinopyroxene spinel olivine

T=0

enclosed sulfide 187

188

Os/

Osi at melt depletion

orthopyroxene

(b)

187Os/188Os

T=1

clinopyroxene

T=1

spinel olivine

T=2 interstitial sulfide enclosed sulfide

187

188

Os/

Os of metasomatic agent

T=0

187

Re/188Os

Figure 12. Evolution of Re–Os isotope systematics amongst peridotitic phases as a result of melting and metasomatism. (a) At the time of melt depletion and the enclosure of an immiscible sulfide, in isotopic equilibrium with co-existing silicate phases, all phases will share a common Os isotope ratio but have varying 187 Re/188Os at T=0. Over time each phase will evolve according to its 187Re/188Os and the Re decay constant (λ = 1.666 × 10−11 yr−1 (T = 1). The enclosed sulfide, with very low 187Re/188Os retains the 187Os/188Osi. (b) At T=2 the mineral assemblage interacts with a metasomatic melt with a higher 187Os/188Os than the original residual solid. This melt may also be sulfur saturated and precipitate metasomatic, interstitial BMS with a high 187 Re/188Os. Interaction of the residual silicates with the metasomatic melt may re-set the Os isotope ratio to that of the metasomatic melt (and interstitial BMS) but the enclosed BMS, buffered against re-equilibration by a high Os content and armored within its host silicate grain, is unaffected, thus preserving Os isotope disequilibrium between the enclosed BMS and the other phases. Modified after Burton et al. (1999).

Harvey, Warren & Shirey

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were slightly more radiogenic than the Type-1 BMS. However, the bulk-rock 187Os/188Os was still significantly less radiogenic than PUM (Meisel et al. 2001) and mean chondrite (Shirey and Walker 1998; Day et al. 2016, this volume). More recently, Harvey et al. (2010, 2011) revisited the issue of Os mass balance in SCLM lherzolites (Fig. 13), examining Os elemental and 187Os/188Os systematics in four lherzolites from Mont Briançon, French Massif Central (Harvey et al. 2010) and two harzburgites and a further two lherzolites from Kilbourne Hole, NM USA (Harvey et al. 2011). The harzburgites tended to have silicates that were more radiogenic than the bulk rock (187Os/188Os = 0.11600 to 0.12131) with low Os concentrations (≤ 142 pg·g−1). Enclosed BMS grains were unradiogenic (187Os/188Os = 0.1185 to 0.1235) with high Os concentrations (6–25 µg·g−1), while interstitial BMS grains were radiogenic (187Os/188Os = 0.1304 to 0.2163) with generally low Os concentrations (0.001–6 µg·g−1, but with one exceptional grain containing 21 µg·g−1). This suggests that enclosed BMS dominates the Os mass balance of these harzburgites, as confirmed by a quantitative Os elemental and isotopic mass balance (Fig. 13). In total, less than 5% of bulk-rock Os was hosted by the silicates, while the contribution from interstitial BMS was no more than 17% (and as low as 3%) and the remainder (up to 95%) of Os was hosted by Type-1 BMS (Harvey et al. 2011). One of the two Kilbourne Hole lherzolites in the Harvey et al. (2011) study was similar to the harzburgites, with the same Os elemental and isotope distributions, but requiring an even smaller contribution of Os from interstitial BMS in order to balance the bulk-rock budget (Fig. 13). In the other Kilbourne Hole lherzolite, both the bulk-rock and silicate phases were supra-chondritic, as were most of the interstitial BMS (187Os/188Os = 0.1291 to 0.1694), even though some subchondritic BMS grains were recovered. In this example, the Os mass balance was controlled more by the interstitial BMS, although Harvey et al. (2011) were unable to calculate an unequivocal Os mass balance in this sample. The lherzolites from Mont Briançon (Harvey et al. 2010) shared many of the Os mass balance characteristics of the Kilbourne Hole peridotites dominated by enclosed BMS (Fig. 13) but, as discussed below, the 187Re/188Os and Re/Os signatures of these lherzolites are more difficult to interpret than a simple enclosed versus interstitial BMS division. 100%

60%

Type-1 sulfide

Olivine

Type-2 sulfide

Orthopyroxene

Silicate / oxide

Clinopyroxene Spinel

5

40%

4 3

20% ?

?

?

2 1

0%

MBr3

MBr8

silicate / oxide Os contribution (%)

80%

MBr20 KH03-15 KH03-16 KH03-21 KH03-24

Figure 13. Osmium mass balance calculations for peridotite xenoliths from Mont Briançon, French Massif Central, and Kilbourne Hole New Mexico, USA. Silicates and spinel account for < 5% of bulk-rock Os budgets, the remainder being accounted for by a balance between the two populations of BMS grains. The difference in Os concentration between enclosed and interstitial BMS is sufficiently large in MBr 3 that interstitial BMS contributes no more than 1% to the whole rock budget whilst maintaining the measured whole rock isotopic ratio. Similarly, in MBr 8 interstitial BMS can contribute no more than 11%. The contribution from interstitial BMS, in the case of KH03-24, is insignificant, and in KH03-15 and KH03-21 account for 3.5–17.5% respectively of the Os present. The relative contributions of the two BMS populations in KH03-16 is less clear but is probably dominated by the enclosed population. The contribution of the interstitial component in KH03-16 is calculated to be less than 17.5% of the total. Modified after Harvey et al. (2010, 2011).

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Overall, the studies of Hart and Ravizza (1996), Burton et al. (1999) and Harvey et al. (2010, 2011) demonstrate that: (i) it is possible to quantify the contributions to bulk-rock peridotite Os mass balance from its constituent phases; (ii) in peridotites where enclosed BMS dominates the Os mass balance, melt depletion is sometimes ancient; and (iii) where interstitial BMS dominates the Os mass balance, bulk-rock Os isotope ratios are not likely to reveal any significant information pertaining to melt depletion. However, this contrasts with the information recorded in interstitial PGM, thought to have formed from extreme melt depletion in spinel lherzolites from Lherz (Luguet et al. 2007). But, as was demonstrated by the radiogenic Kilbourne Hole lherzolite of Harvey et al. (2011), analysis of enclosed BMS from otherwise metasomatized peridotites has the potential to “see through” the late events that precipitated radiogenic interstitial BMS and / or PGM and alloys.

Rhenium mass balance Rhenium has high Kd,sulfide/silicate, but estimated values are significantly lower than the equivalent partition coefficient for Os. For example, the results of experiments designed to measure partitioning of Re and Os between sulfide and silicate melt (Brenan 2008; Brenan et al. 2016 this volume) produced values of Kd Ossulfide/silicate/ Kd Resulfide/silicate >150, which is the minimum required to produce the Re/Os fractionation observed in mantle derived magmas. However, the absolute value for Kd Resulfide/silicate varies over a wide range, from > 20,000 to ~ 20, depending on fO2–fS2 conditions imposed during an experiment (Fonseca et al. 2007; Brenan et al. 2016 this volume). This suggests that Re mass balance in peridotites should also be strongly controlled by BMS and, given the strong partitioning of Re into chalcopyrite and hence interstitial BMS, this phase should dominate the Re bulk-rock mass balance. However, a quantitative Re elemental mass balance that is consistent with observed 187 Re/188Os in peridotites is still lacking. Achievement of this mass balance is difficult due to the low abundance of Re reported in silicates (typically < 100 pg·g−1) combined with difficulties in obtaining low Re analytical blanks. In addition, the partial lithophile character of Re compared to the other HSE (Handler and Bennett 1999; Burton et al. 1999; Harvey et al. 2010, 2011) and associated implications for Re–Os fractionation during mantle melting (e.g., Righter and Hauri 1998; Mallmann and O’Neill 2007) make Re behavior with melting more difficult to predict, especially in the presence of residual sulfide. Rhenium is easily removed from a peridotite source during partial melting as it has a low bulk Kd peridotite/basaltic melt (Brenan et al. 2003). The bulk-rock incompatibility of Re in the mantle leads to its high concentration in melt (Alard et al. 2002; Harvey et al. 2010, 2011; Warren and Shirey 2012), especially as BMS dissolve into undersaturated silicate melt during partial melting. This first order interpretation of Re mass balance is not sufficient to interpret the Re–Os isotope systematics of peridotites with complicated histories such as those from Mont Briançon (Harvey et al. 2010), as the two distinct populations of BMS are difficult to distinguish based on 187Re/188Os, Re and Os concentrations. Similar to the Griffin et al. (2002) study of Siberian xenoliths, Mont Briançon has different populations of BMS that were interpreted as having interacted with each other (see the section on peridotite-hosted BMS above). In both studies, unradiogenic Os isotope ratios have often been preserved in sulfides that at least originated as enclosed BMS, but 187Re/188Os ratios were elevated and high Re concentrations often accompanied high Os concentrations. The original low Re/Os elemental ratio of enclosed BMS may have become elevated through melting and mixing with interstitial BMS melt, but the immediate effect on 187Os/188Os was negligible because of the low Os concentration in interstitial BMS. It is only with the addition of time-integrated 187Os from 187Re decay that the original 187Os/188Os of the hybrid BMS becomes progressively obscured.

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Lead mass balance Although the Pb isotope composition of the mantle has been of interest to geochemists for over half a century (e.g., Holmes 1946; Patterson 1956; Allègre 1969), the role that BMS play in the Pb mass balance of the mantle has been less well explored. Recent interest in the Pb elemental and isotopic composition of mantle sulfides stems from a potential solution to the first Pb paradox (i.e., the position of oceanic Pb isotopic compositions in the “future” Pb position, to the right of the geochron; Allègre 1969), as discussed in the final section of this chapter. A quantitative Pb mass balance (elemental and isotopic) based upon individual peridotites has yet to be accomplished, despite studies that have suggested that mantle Pb resides predominantly in BMS (Meijer et al. 1990; Hart and Gaetani 2006). Hart and Gaetani (2006) predicted that if BMS are the main mantle Pb reservoir they should represent c. 0.05% of the mantle to be consistent with PUM sulfur concentration estimates (Lorand 1990; O’Neill 1991; Palme and O’Neill 2003) and contain ca. 75 µg·g−1 Pb. More recently, Pb concentrations in individual BMS grains extracted from abyssal peridotites from the MAR, Gakkel and SWIR were measured by Burton et al. (2012) and Warren and Shirey (2012). Warren and Shirey (2012) found concentrations that ranged from 0.12–12 µg·g−1 (mean 4 µg·g−1) whereas those in Burton et al. (2012) ranged from 1.5–52 µg·g−1 (mean 19 µg·g−1). These values possibly over-estimate the average Pb concentration as data becomes increasingly difficult to obtain at abundances below about 1 µg·g−1, due to the minute quantities involved in an individual BMS grain analysis. Using these estimates for BMS Pb concentrations yields a significant deficit of Pb in many abyssal peridotites (Warren and Shirey 2012). For example, using the average Pb concentration of 4 µg·g−1 Pb found in Gakkel and SWIR BMS (Warren and Shirey 2012) and the average abyssal peridotite BMS modal abundance of 0.02% (Luguet et al. 2001 2003) gives a bulk peridotite Pb concentration of 0.8 ng·g−1, significantly lower than the estimate of 150 ng·g−1 for PUM (McDonough and Sun 1995) and also lower than the estimate of 21 ng·g−1 for depleted MORB mantle (DMM; Salters and Stracke 2004; Workman and Hart 2005). Using the maximum modal estimate of 0.06% for mantle BMS in abyssal peridotite (Luguet et al. 2003) and the maximum BMS Pb concentration observed by Warren and Shirey (2012) of 12 µg·g−1, only about 40% of Pb in DMM can be accounted for. Using the mean value of Pb in BMS analyzed by laser ablation (Burton et al. 2012) from the 15° 20’ Fracture Zone, MAR (19 µg·g−1) can account for approximately 60% of DMM Pb. Using the maximum value of 52 µg·g−1 and the maximum modal abundance of 0.06 % can account for the mean DMM Pb abundance of 21 µg·g−1, but would still only represent ca. 25% of the estimated PUM Pb abundance. However, these maximum estimates of BMS Pb concentration and abundance are not representative of the oceanic mantle. Observations of xenoliths suggest that Pb is present in the silicate phases at concentrations < 1 µg·g−1 (e.g., Meijer et al. 1990; Hauri et al. 1994; Carignan et al. 1996; Eggins et al. 1998; Norman 1998; Ionov et al. 2002), but given their modal abundance this can account for a significant proportion of the bulk Pb. For example, Carignan et al. (1996) measured the Pb concentrations of the main silicate phases (olivine 11 ng·g−1; orthopyroxene 20 ng·g−1; clinopyroxene 470 ng·g−1) in two peridotite xenoliths and calculated their bulk-rock Pb abundances to be 21 and 29 ng·g−1, without the inclusion of peridotite-hosted BMS. These values were remarkably close to the measured bulk-rock Pb concentrations of 16 and 40 ng·g−1, respectively. In this instance the three main silicate phases in peridotites can account for 73– 100% of mantle Pb concentrations. Overall, it would appear that mantle BMS contain much less Pb than that predicted by Hart and Gaetani (2006). The major difficulty with a Pb elemental and isotopic mass balance for the oceanic mantle is a lack of data on Pb in the silicate phases and bulk rock. Warren and Shirey (2012) provided an elemental mass balance for a single SWIR sample based on Pb concentration data for pyroxenes

Mantle Sulfides’ Role in Re–Os and Pb Isotope Geochronology

615

and bulk peridotite. They found reasonable agreement between the measured bulk-rock concentration of 31 µg·g−1 and the reconstructed concentration of 22 µg·g−1 (excluding sulfides), as olivine would only need to contain 13 µg·g−1 to fully match the measured value. However, the measured bulk-rock concentration of 31 µg·g−1 is higher than the current estimate of 21 µg·g−1 for Pb in DMM, indicating the need for more systematic analysis of Pb concentrations in abyssal peridotite, particularly as studies of peridotite xenoliths have shown that bulk Pb concentrations can be highly variable (e.g., Carignan et al. 1996). This will also require improved constraints on the effect of alteration on peridotite Pb concentrations. Clearly several fundamental unknowns remain to be investigated with respect to the Pb isotopic and elemental mass balance in the mantle, most importantly, the contribution that mantle BMS make to solving the first Pb paradox.

GEOCHRONOLOGICAL METHODS, MODEL AGES, AND POTENTIAL PITFALLS. Sulfide Re–Os isochrons, TMA, TRD, and γOs. An isochron is the “gold standard” in geochronology, notwithstanding the possiblility that an isochron may in fact represent a mixing line. For sulfides, Re–Os isochrons (Fig. 14) can be generated when a selection of co-genetic BMS in chemical equilibrium at the time of system closure, but with a range of Re–Os, evolve over time, under closed-system conditions. This, in theory at least, allows their isochronous relationship to be used as a precise measure of the time elapsed since they were isolated from open-system conditions. Isochrons also have the advantage of providing both age and initial isotopic composition. By projecting the slope of the isochron back to the intercept on the y-axis, i.e., at zero 187Re/188Os, this allows the 187Os/188Os at the time the system closed to be calculated, and hence inferences about the source of the sulfides and petrogenetic history can be made. Figure 14 shows examples of true isochrons among diamond-hosted BMS populations from Aulbach et al. (2004a) and Westerlund et al. (2006). Harvey et al. (2006) came close to a true isochron with a population of enclosed serpentinite-hosted BMS from the 15° 20’ N Fracture Zone on the MAR (mean

(a)

3.52 +/- 0.17 Ga (MSWD = 0.46)

(b)

0.120

0.120 0.115

0.116

0.110

187Os/188Os initial = 0.10725 +/- 0.00014

0.105 0.100 0.0

0.124

187Os/188Os

187Os/188Os

0.125

3.27 +/- 0.34 Ga (MSWD = 0.75)

Aulbach et al. (2004) 0.1

0.2 0.3 187Re/188Os

0.4

187Os/188Os initial = 0.1093 +/- 0.0001

0.112

0.108 0.0

Westerlund et al. (2006) 0.1

0.2 187Re/188Os

0.3

Figure 14. Rare examples of true Re–Os isochrons (MSWD < 2.4) derived from diamond-hosted BMS grains recovered from the Slave Craton, Canada. (a) Re–Os isochron age of 3.27 ± 0.34 Ga (MSWD = 0.75) for BMS grains recovered from xenocrystic olivine and pyroxenes from kimberlites in the Lac de Gras area, central Slave Craton (modified after Aulbach et al. 2004a). (b) Re–Os isochron age of 3.52 ± 0.17 (MSWD = 0.46) derived from diamonds in harzburgitic host rocks from the 53 Ma Panda kimberlite pipe, Ekati Mine, North West Territories, Canada (modified after Westerlund et al 2006).

Harvey, Warren & Shirey

616

square weighted deviation or MSWD of 3.3), as shown in Figure 15. For this same location, Burton et al. (2012), in a plot of 207Pb/204Pb versus 206Pb/204Pb (Fig. 16), determined an isochron age of 1.83 ± 0.23 Ga (MSWD = 0.18) for enclosed serpentinite-hosted BMS. This value agrees, within analytical uncertainty, with the Re–Os age of 2.06 ± 0.26 Ga from Harvey et al. (2006); i.e., compare Figs. 15 and 16. Unfortunately, many arrays of data derived for non-diamond related sulfides have MSWD values that are > 2.4, indicating poor statistics for the data regression (Faure 2001), and cannot meet the conditions for a precise Re–Os geochronometer. Instead, isotopic data for these BMS can be compared to reference isochrons as an alternative to actual isochron regressions. In many instances, this is the closest to an isochron age that many studies are able to achieve. In BMS from diamonds, such data arrays can give a useful indication of the general age for the diamonds in question. That such BMS grains might not fall on a strict isochron is understandable given that diamonds are xenocrysts whose spatial relationship to each other in the lithosphere is not known and thus may be too far apart to have achieved strict chemical equilibrium.

0.119

(a)

187Os/188Os

bulk rock 0.117

100

1000

10000

peridotite / seawater interaction

see inset below

0.115 enclosed sulfides interstitial sulfides

0.113 0.001

0.01

0.1

1

10

187Os/188Os

0.119

(b) 2.06 +/- 0.26 Gyr (MSWD = 3.3) 187 Os/188Osinitial = 0.11396 +/- 0.00030

0.117

0.115

0.113

100

Hole 1274a ODP Leg 209 15o 20’ N Fracture Zone Mid-Atlantic Ridge Harvey et al. (2006)

0

0.02

0.04

0.06

0.08

0.1

187Re/188Os

Figure 15. (a) 187Re–187Os isotope evolution diagram for serpentinite-hosted BMS grains from sample 2R1 31–37, Hole 1274A, 15º 20’ N Fracture zone, Mid-Atlantic Ridge. Most of the BMS grains yield an indistinguishable 187Os/188Os isotope composition from that of the whole-rock, but possess a wide range of 187 Re/188Os. The solid curve illustrates the calculated trajectory for mixing of the low 187Re/188Os sulfides with seawater, and shows that even for sulfide/seawater ratios of 104 there is no measurable shift in 187Os/188Os ratio. (b) Regression of high-Os sulfides, with a rounded morphology, yields a best-fit line corresponding to an age of 2.06 ± 0.26 Gyr (MSWD = 3.3) using Model 3 of Ludwig (1997). Modified after Harvey et al. (2006).

Mantle Sulfides’ Role in Re–Os and Pb Isotope Geochronology

617

15.8 enclosed MAR sulfides interstitial MAR sulfides SWIR sulfides Gakkel sulfides

15.4 on chr

2G

on

chr

iso

geo

15.2

ce

ren

fe a re

15.0 15

enclosed MAR sulfides 1.83 +/- 0.23 Ga (MSWD = 0.18) Burton et al. (2012)

4.57

Ga

207

Pb/204Pb

15.6

16

17 18 206Pb/204Pb

19

20

Figure 16. 207Pb-206Pb/204Pb isotope diagram for BMS grains from abyssal peridotites from a single MAR abyssal peridotite (Burton et al. 2012) and from multiple Gakkel and SWIR peridotites (Warren and Shirey 2012). For the MAR sample, interstitial BMS (white circles) possess radiogenic Pb isotope compositions consistent with recent seawater interaction. In contrast, included MAR BMS grains (black circles) are unradiogenic, yielding a 207Pb–206Pb age of 1.83 ± 0.23Gyr (MSWD = 0.18, model 2; Ludwig 1997). Base metal sulfide grains from the Gakkel (5 samples) and SWIR (7 BMS grains from 5 samples) show no clear distinction in Pb composition between included and interstitial BMS. Two standard error bars are smaller than the symbol sizes in all cases. Modified after Burton et al. (2012) and Warren and Shirey (2012).

The complexity of melt depletion and subsequent metasomatism, tectonism and metamorphism often obscures any meaningful information that could have been extracted from a population of BMS grains prior to their disturbance. Diamond-hosted BMS are less susceptible to these complexities because they are encapsulated and protected from secondary melts and fluids, but they do have xenocrystic-related uncertainties. As a consequence of these difficulties in producing Re–Os isochron ages, model ages are often employed instead. A model age represents a time when the isotopic growth curve for a mineral coincides with the isotopic evolution of a reference reservoir such as the mantle. A model age is an attempt to derive meaningful geochronological information with imperfect knowledge of the sequence of events that have affected the samples, or when there is simply an insufficient number of samples with which to construct an isochron. For example, they allow an “age” to be derived from a single data point. In the case of the Re–Os isotope system, two types of model age calculations are in common use (Fig. 17). Both calculations compare 187Os/188Os in an unknown to the 187 Os/188Os evolution curve for mantle with chondritic Re/Os. The evolution of 187Os/188Os from the initial osmium isotope ratio of IIIA iron meterorites (the most primitive Solar system initial value: 187Os/188Osi = 0.09531; Smoliar et al. 1996) is used to estimate the average composition of chondritic meteorites for the present day (Fig. 17a). A 187Re decay constant (λ) of 1.666 × 10−11 yr−1 is used for this and all other age-dependent calculations: 187

9

Os/188Ost = 187Os/188Osi + 187 Re/ 188Oschon (eλ (4.558×10 ) − eλt ).

(1)

The measured composition for chondrites includes early measurements that gave 0.1262 ± 0.0005 (Allègre and Luck 1980; Luck et al. 1980; Walker and Morgan 1989). Meisel et al (1996) reported a narrow range of present-day 187Os/188Os ratios for 11 ordinary chondrites of 0.1289 ± 0.0022 (2σ) and a similar present-day 187Os/188Os ratio of 0.1283 ± 0.0008 (2σ) for 8 enstatite chondrites. The latter was later refined by Walker et al. (2002a) and termed the enstatite chondritic reservoir (ECR), with present-day 187Os/188Os = 0.1281 and 187Re/188Os = 0.421. In contrast, a suite of 4 carbonaceous chondrites defined a 1–2% lower 187Os/188Os ratio of 0.1263 ± 0.0008. A detailed discussion of the variability of chondrite 187Os/188Os can be found in this volume in the chapter on HSE in planetary bodies (Day et al. 2016, this volume).

0.15

(a)

187Re/188Os

187

Os/

188

Os

0.14

s=

0.13

an

Os

s=

ch

on

dri

te

ch

γO

s=

0.11 γOs

(b)

0.04 0

dit

ic

-10

ev

olu

tio

n IIIA

iro

=-

nm

ete

20

ori

tes

(i) sulfide in peridotite xenolith brought to surface in modern basalt (iii) TRD assumes no Re in sulfide: calculated from intersection of 187Os/188Os in sulfide with chondrite evolution curve

Os/

0.13

187

(ii) No correction for eruption age needed

0.11(ii)

(iv) In growth of 187Os from ingrowth of measured 187Re corrected for in calculation of TMA. TRD is therefore a minimum and will underestimate TRD when Re abundance is undisturbed (iii)

(i)

(iv)

0.10 0.15

0.13

(c)

(i) sulfide recovered from xenolith with eruption age of 0.3 Ga. Re added at time of eruption (ii) 187Os corrected for ingrowth since eruption

iv

Os/ 187

(i)

0.12 0.11 0.10

TMA

TRD

0.14

Os

0

0.14

0.12

188

0.40

+2

+1

on

0.10

188

γO

me

0.12

0.15

200

γO

(iii)

ii

(iii) TRD calculated from intersection with chondrite evolution curve (iv) Re addition leads to meaningless “future” TMA age

TMA 0

TRD 1

2

Time (Ga)

3

4

Figure 17. Graphical representation of (a) the chondrite evolution curve over the course of Earth history and materials with higher and lower γOs (from −20 to +20). Starting composition for chondrite evolution curve is 187Os/188Os = 0.09531 (Smoliar et al. 1996). Present day mean chondrite 187Os/188Os = 0.1270 (Shirey and Walker 1998). (b) Graphical representation of the difference between TMA and TRD derived from a BMS grain recently erupted in a peridotite xenolith (and no perturbation of original Re). This contrasts with (c) where the sulfide experienced Re-enrichment from the host lava en-route to the surface. Recent Re addition does not materially affect the 187Os/188Os of the sulfide but gives meaningless “future” ages. Modified after Shirey and Walker (1998).

Harvey, Warren & Shirey

618

Mantle Sulfides’ Role in Re–Os and Pb Isotope Geochronology

619

The two commonly employed model age calculations (Walker et al. 1989; Shirey and Walker 1998) are TMA (Eqn. 2) and TRD (Eqn. 3), where both model ages use Equation (1) to calculate a mantle evolution curve. For TMA, both the 187Re/188Os and 187Os/188Os ratios in the unknown are used to calculate the time of separation from asthenospheric mantle: TMA =

 (187 Os /188Os −187 Os/188Os  1 chon sample ) ln  187 + 1. 188 187 188 λ   Re / Oschon − Re / Ossample

(

(2)

)

This model age can be used both for BMS with low Re/Os (e.g., peridotite-hosted Type-1 BMS and diamond-hosted peridotitic BMS) and for BMS with much higher Re/Os (e.g., pyroxenite-hosted BMS, peridotite-hosted Type-2 BMS, and diamond-hosted eclogitic BMS). It assumes that the Re/Os measured today in a sample is representative of its long-term history in the mantle. If this assumption is correct, this implies that the geochronometer started at the time that the sulfide cystallized and no subsequent gain or loss of Re occurred. As such, TMA ages can provide robust geochronological information regarding the age of an individual BMS grain (Fig. 17b). It is not always safe to assume that undisturbed Re/Os systematics have been preserved. For example, where an elevated Re/Os elemental ratio accompanies an unradiogenic 187Os/188Os, or the converse, where an unsupported radiogenic 187Os/188Os (i.e., high 187Os/188Os with low Re/Os) is observed, a TMA calculation will yield an erroneous number that likely bears no resemblance to any actual event in the history of the sulfide (Fig. 17c). Such a situation can occur when Re from the host lava (such as a kimberlite) is added to a xenocryst or xenolith during transport to the surface. Under these circumstances a TMA age might yield a nonsense or “future” age and in this case a TRD, or rhenium depletion age, is likely to be the most appropriate model age calculation. TRD is calculated as an intersection time with chondritic mantle growth, assuming a melt-depletion event previously removed all Re from the sample. It is calculated using the following equation: TRD =

1 ln λ

  

(

187

Os/188Oschon −187 Os/188Ossample (EA) 187

188

Re/ Oschon

) + 1,

(3)

 

where 187Os/188Ossample (EA) is the ratio at the time of separation from the mantle: 187

(

)

Os/188Ossample(EA) = 187Os/188Ossample −187 Re/188Ossample e λt – 1 .

(4)

For a xenolith-hosted BMS (Fig. 17c), 187Os/188Ossample (EA) represents the osmium isotope ratio at the time the host xenolith was entrained and brought to the surface in a lava. However, the correction in the sample’s isotopic composition defined by 187Os/188Ossample(EA) is only necessary if the eruption of the host lava is far enough in the past or enough Re has been added by the host lava to dramatically change the 187Re/188Os such that sufficient ingrowth of 187Os could be anticipated. With TRD, the assumption is made that Re was completely removed from the sulfide during the event that isolated it from the convecting mantle. Therefore, in the absence of any ingrowth of 187 Os from 187Re, the measured 187Os/188Os of the sulfide has remained unchanged since the sulfide crystallized. This is clearly an oversimplification, as even undisturbed diamond-hosted peridotitic BMS and Type-1 peridotite-hosted BMS contain some Re and some 187Os ingrowth is inevitable. Hence the measured 187Os/188Os used in a TRD age calculation will only yield a minimum age (Fig. 17) but one that could nonetheless be useful if the Re concentration is low. A quick examination of equation (3) reveals that relying entirely upon a measured 187Os/188Os ratio for a model age may not yield the most accurate geochronological information and that, in examples where the measured 187Os/188Os exceeds that of the chondritic reference value, it is not possible to obtain any useful geochronological information. Meaningless TRD ages in the future

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and TMA ages older than 4.5 Ga reflect derivation of Os from a source more radiogenic than PUM and / or the local disturbance of the Re–Os system. Fortunately, most host lavas have very low Os content and thus their ability to modify the Os isotopic composition of a mantle peridotite or a BMS grain is very limited (e.g., Walker et al. 1989; Shirey and Walker 1998). An additional method has been developed to handle datasets produced by LA MC ICPMS as a consequence of the relative ease with which large numbers of 187Re/188Os and 187Os/188Os ratios can be generated (e.g., Griffin et al. 2002, 2004; Marchesi et al. 2010; González-Jiménez et al. 2014). Calculating model ages for these sulfides can give a high density of “ages” for a given region, allowing model ages to be plotted on histograms. This in turn allows peaks in the frequency of ages to be determined. More recently, large datasets of model ages have been presented as relative probability plots. For example, the model ages generated from BMS inclusions in olivines from Udachnya, Siberia (Griffin et al. 2002, using the models of Ludwig 2000) were incorporated into a relative density plot and used to suggest that most of the lithospheric mantle beneath the Daldyn kimberlite field formed during the period 3–3.5 Ga. Bulk-rock data and BMS data generated by TIMS could also be used for this type of analysis, but the time taken to acquire sufficient data using TIMS is considerably longer. A case in point are the relative probability plots for the Kaapvaal and Slave cratons presented in Pearson and Wittig (2008), which are each comprised of TIMS data compiled from measurements acquired over several years (Kaapvaal: Pearson et al. 1995a; Carlson et al. 1999; Menzies et al. 1999; Irvine et al. 2001; Carlson and Moore 2004; Simon et al. 2008; Slave craton: Irvine et al. 2003; Aulbach et al. 2004a; Westerlund et al. 2006). In addition to a large peak in BMS Re–Os model ages at 2.5–3.0 Ga for the Kaapvaal craton (Fig. 18), the relative probability plots of Pearson and Wittig (2008) also demonstrated a main peak in BMS model ages around 2.8–2.9 Ga (see also Aulbach et al. 2004a) and subsidiary peaks at 1.3 Ga, 2.2 Ga, and 3.9 Ga for the Slave Craton (Fig. 18). These examples demonstrate the power of large numbers of Re–Os model ages to constrain the timing of regional scale events and the link, or lack thereof, between ages preserved in the SCLM and crust forming events (i.e., largescale partial melting). Aulbach et al. (2016, this volume) discuss the significance of relative probability plots in more detail, including the use of HSE in dating cratons. Finally, often quoted alongside Re–Os model ages, although not a model age itself, is the concept of γOs (gamma osmium), where the Os isotope ratio of an unknown is compared to that of a chondritic reference, with the difference at a specific time between the unknown and the chondritic reference (γOs) reported as either a positive or negative percentage (Fig. 17a). The underlying concept is exactly the same as the delta notation (δ) commonly used for many stable isotope systems (per mil deviations from a known standard) and the epsilon notation (ε) for Nd and Hf isotope ratios (parts per 10,000 deviation from a known standard):  γOs(t ) =  

( (

187

) – 1  ) 

Os/188Ossample(t )

187

Os/188Oschon(t )

× 100.

(5)

Samples with positive γOs often are described as enriched, radiogenic or suprachondritic and imply long-term elevated 187Re/188Os. Samples with negative γOs often are described as depleted or unradiogenic and imply long-term lowered 187Re/188Os.

Potential pitfalls with sulfide geochronology. The fractionation of Re and Os during mantle melting preserves distinct Re/Os and 187Os/188Os systematics in primary BMS grains enclosed within either a silicate or diamond host (i.e., Type-1). These contrast with Type-2 BMS that are associated with metasomatism, which makes the Re–Os isotope system an ideal tool for dating melt depletion in the former and metasomatism in the latter. However, as with all isotope systems, there are limits to the utility of the Re–Os isotope system.

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Relative probability

Kaapvaal craton

bulk rock TRD (n = 228)

sulfide TMA (n = 71)

bulk rock TMA

(n = 228)

Relative probability

Slave craton

bulk rock TRD (n = 36)

sulfide TMA (n = 48)

0

0.5 1.0 1.5 2.0 2.5

3.0 3.5 4.0 4.5

Model age (Ga)

Figure 18. Probability density plots of Re–Os model ages for bulk-rock peridotites and BMS grains for the Kaapvaal and Slave cratons. Kaapvaal: Pearson et al. (1995a), Carlson et al. (1999), Menzies et al. (1999), Irvine et al. (2001), Carlson and Moore (2004) and Simon et al. (2008); Slave: Irvine et al. (2003), Aulbach et al. (2004a) and Westerlund et al. (2006); TRD ages are corrected to the eruption age of the host kimberlite. Model age equations for TRD and TMA of Walker et al. (1989) and Pearson and Shirey (1999) are calculated using for 187Os/188Os= 0.1283 and 187Re/188Os = 0.422 chondritic mantle. Modified after Pearson and Wittig (2008).

Ages calculated from sulfide isotopic compositions are subject to a number of potential pitfalls, including uncertainties in model age calculations, modifications induced during sample preparation or analysis, and changes to original Re–Os isotope systematics that derive from mantle processes. All of these modifications may obscure or even render useless model age information preserved in mantle BMS. Uncertainties in model age calculations. Although the sulfides which are the subject of this chapter are mantle-derived, to obtain geologically meaningful TRD and TMA ages a choice of the reference reservoir for 187Os/188Os at time in the past must be made. The range and heterogeneity of published mantle values and, in particular, the wide disparity between mean 187Os/188Os for abyssal peridotite versus that of PUM (ca. 4% difference in 187Os/188Os between these two reservoirs, discussed in detail in Rudnick and Walker 2009; cf. Gannoun et al. 2016, this volume) means that a wide range of model ages for a single BMS grain could be derived depending upon

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which mantle value was chosen for normalization. Using abyssal peridotite or PUM as a basis for calculating TMA and TRD ages yields younger and older model ages, respectively, than using mean chondrite. The main problem with PUM estimates is that PUM is an artificial construct designed to predict the composition of a notional pristine mantle by projecting back from samples that have been modified over the course of Earth’s history. The disparity in the evolution models for the mantle impacts model ages by ca. 300 Ma (e.g., Pearson and Wittig 2008; Rudnick and Walker 2009). These differences are relatively unimportant for ancient systems but can be significant when attempting to date Phanerozoic events using the Re–Os geochronometer. Therefore model age calculations relative to PUM are probably not appropriate (Rudnick and Walker 2009). The carbonaceous chondrite average 187Os/188Os of 0.1262 ± 0.0005 (see Day et al. 2016, this volume) is probably still the best, and is the most commonly used reference for calculating model ages. In general, however, model ages rarely can carry a precision better than several hundred million years and this imprecision will offset some of the need for an accurate knowledge of the compositional evolution of the reference reservoir (e.g., the mantle). Sample preparation issues. Some diamond-hosted BMS break up during the process of removing them which can result in skewed Re/Os ratios (Richardson et al. 2009). This is because Re is preferentially partitioned into chalcopyrite relative to pyrrhotite and/or pentlandite whereas the reverse is true for Os (e.g., Craig and Kullerud 1969; Li et al. 1996; Piña et al. 2012). The separation of chalcopyrite from pyrrhotite or former MSS will typically occur at a much later time (e.g kimberlite eruption) relative to the time averaged Re/Os that was established for the whole sulphide grain. This can introduce a different measured 187 Re/188Os if chalcopyrite or pyrrhotite makes up a disproportionate part of the analysis. For example, Richardson et al. (2001) reported up to 30% difference in 187Re/188Os amongst fragments of chalcopyrite, pentlandite and pyrrhotite recovered from eclogitic diamonds. An artificial lowering of 187Re/188Os will skew the calculated TMA of an affected sulfide to an artificially young model age (see Eqn. (2); Fig. 19). Consequently, Richardson et al. (2001) attributed scatter on Re–Os isochrons to the possible incomplete recovery of sulfide inclusions, in particular chalcopyrite-rich rims. The absence of well correlated linear arrays for BMS Re–Os isotope data was also attributed by Beyer et al. (2004) to the potential loss of Re-rich chalcopyrite rims during the preparation of polished sections for in situ analysis by LA MC ICPMS. With careful sample preparation, these artificially induced Re/Os fractionations and potential corruptions of model age calculations may be avoided, or at least minimized. Mantle-derived changes to Re–Os systematics. Negative model ages, often meaninglessly referred to as “future ages”, along with model age calculations that return “ages” that are greater than the age of the Earth are the result of a departure away from the idealized Re–Os isotope systematics of Type-1 peridotite-hosted (or diamond-hosted peridotitic) BMS grains. Examination of Equations (2) and (3) reveals that if melts with Re or Os interact with a population of BMS grains, the original geochronological information that was contained within that population may be obscured. Hence, BMS model ages should not be used indiscriminately, as their compositions may reflect cumulative effects of several events. Of particularly importance are the addition of Re (both recent and ancient) and radiogenic Os, and the implications of these on TMA and TRD are summarized in Figure 19. Osmium enrichment. Unsupported radiogenic Os (i.e., insufficient 187Re to account for the 187Os present in a geologically reasonable time, also referred to as suprachondritic; Fig. 19) can be attributed to subduction-related processes. Subduction can involve sulfides equilibrated with the slab (the crustal portion of which is basaltic with a high Re/Os) and fluids in the wedge which can move Os. Highly oxidising subduction-related fluid liberates radiogenic Os, but not Re, from the downgoing slab (Brandon et al. 1996), resulting in a higher 187Os/188Os that suggests derivation from a time-integrated high-Re/Os reservoir. This elevated Os isotopic composition appears to be a general feature of arc peridotite whole rocks (e.g., Widom et al

Mantle Sulfides’ Role in Re–Os and Pb Isotope Geochronology Os/188Os > chondrite Re/188Os < chondrite

187

187

187

Os/188Os > chondrite Re/188Os > chondrite

Addition of radiogenic Os Re addition in antiquity TRD underestimated TMA exaggerated or meaningless TRD meaningless “future” ages TMA more reliable chondrite 187

187

187

Os/188Os < chondrite Re/188Os < chondrite

187

TRD and TMA broadly equivalent

0

Both model ages could give plausible ages

0

Os/188Os < chondrite Re/188Os > chondrite TRD < TMA

chondrite

0.1270

187Os/188Os

187

623

Recent Re addition, insufficient time for ingrowth of 187Os TRD give plausible ages

0.40186 187Re/188Os

Figure 19. The effects of recent and ancient Re addition on sulfide TMA and TRD compared to seemingly undisturbed 187Os/188Os and 187Re/188Os (i.e., both < mean chondrite). Values which plot in the lower left quadrant do not necessarily preclude any metasomatism since early melt depletion (cf. Wang et al. 2009, where interaction with modest amounts of silicate melts may still yield plausible TMA and TRD), but may be considered the least likely to be seriously compromised.

2003). For this to be observed in an individual BMS grain requires a multistage history for that grain, which may have crystallized as an unradiogenic MSS with a low Re/Os ratio, followed by interaction with a melt or fluid that contains radiogenic Os, but relatively little Re (e.g., Wang et al. 2009). This process has the effect of raising the 187Os/188Os, but suppressing still further the originally low Re/Os ratio of the sulfide. These features were evident in the peridotitic BMS inclusion suite studied from the Ekati Mine, in the Panda kimberlite pipe, Slave craton, where low Re/Os BMS grains were enriched in radiogenic Os associated with specific zones of diamond growth (Westerlund et al. 2006). Complicated, multistage scenarios are not as common in sulfide inclusions in diamond as they are in sulfides from mantle peridotites (see below), because the encapsulation in diamond gives the sulfide an impervious host. While a detailed appraisal of subduction-related HSE behavior is beyond the scope of this chapter, some selected examples that are specific to BMS Re–Os isotope systematics are highlighted here. Gonzáles-Jiménez et al. (2014) reported suprachondritic 187Os/188Os in a population of BMS grains analyzed in situ in a suite of peridotite xenoliths from Calatrava, Spain (Fig. 20). As a consequence of elevated 187Os/188Os, none of the disturbed BMS Re–Os isotopic ratios yielded geologically reasonable TMA or TRD model ages. They attributed these perturbed Re–Os isotope systematics to interaction with fluids originating in a supra-subduction zone environment where dehydration produces a fluid enriched in radiogenic Os (e.g., Brandon et al. 1996; Reisberg et al. 2004; Walker et al. 2002b). Beyer et al. (2004) also reported radiogenic initial 187Os/188Os (0.1282 ± 0.0061) in BMS grains from the Western Gneiss Region of Norway and implicated the introduction of Os derived from sources with higher long-term Re/Os than the peridotites, likely local garnet pyroxenite and eclogite dykes. Similarly, in a study of diamond-hosted BMS

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P = Pangea G = Gondwana R = Rodinia C = Columbia 0

0.5

P

G

7

2.0

1.5

R

C

2.5

supercontinents

0.55 Sulfides with 187Re/188Os 2 Ga, whereas inclusions in the outer portions of the diamond yielded much younger ages that approached the age of the kimberlite eruption that brought the diamonds to the surface. A consequence of the uncertainties associated with the Pb isotope evolution of the mantle was that early Pb–Pb model ages (Rudnick et al. 1993b) were associated with large uncertainties. In Siberian diamonds, small BMS grains are frequently found in the innermost zones of diamonds and it has been suggested that these inclusions acted as seeds for the nucleation of the diamonds (Bulanova et al. 1996), although how long these BMS grains exist before the actual growth of the diamond is uncertain. Pearson et al. (1998) use this line of evidence to argue for syngenetic growth of diamond and BMS, which implies that BMS growth was effectively isochronous with the host diamond. Shirey and Richardson (2011) suggested that although there is evidence for silicate inclusions often being 500 to 1500 million years older than the diamond crystallization age, BMS inclusions tend to preserve ages that are only 10–100 million years older than the diamond that hosts them, making BMS inclusions a much more reliable measure of the age of diamond formation. To improve upon Sm–Nd isochron ages derived from composites of silicate inclusions (e.g., Richardson and Harris 1997), a means of accurately and precisely dating individual BMS inclusions was sought in order to better elucidate the timing of diamond growth. Bulanova et al. (1996), using a proton probe, determined that many diamond-hosted BMS contained substantial Os that, in retrospect, would have been high enough to date with the Re–Os isotope system,

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either to produce model ages on individual grains or to generate isochron style Re–Os plots with multiple cogenetic BMS grains. A modified version of the Os micro-distillation method of Roy-Barman and Allègre (1995) was used to obtain total procedural blanks approaching 10−15 g (Pearson et al. 1998; Pearson and Shirey 1999). This level of blank, which had previously been precluded by the conventional Carius tube digestion method (Shirey and Walker 1995), permitted the analysis of sulfides containing sub-µg·g−1 to µg·g−1 concentrations of Os (corresponding to an absolute amount of Os ranging from 0.1 to 10 pg). This technique produced the first successful date on a suite of diamonds from Koffiefontein, South Africa (Pearson et al. 1998). Since then this has been the preferred method for the dating of diamond formation (e.g., Pearson et al. 1999a,b; Pearson and Shirey 1999; Richardson et al. 2001, 2004, 2009; Shirey et al. 2001, 2002; Aulbach et al. 2004a,b, 2009a,b,c, 2011; Westerlund et al. 2004, 2006; Smit et al. 2010).

Diamond formation through time Peridotitic diamond protoliths mostly date from the Paleoarchean and silicate inclusions in these diamonds have the depleted major element composition expected for refractory continental mantle (Richardson et al. 1984; Westerlund et al. 2006). These paleoarchean BMS grains have low Re/Os because they formed in equilibrium with mantle peridotite that was residual after melt extraction, consistent with their observed low Re concentrations (< 10 pg·g−1 Re). Eclogitic BMS, in general, forms from protoliths that have a systematically higher initial Os isotopic composition than peridotitic BMS, implying that some degree of prior Re/Os enrichment must have occurred, i.e., they are derived from metamorphosed basaltic or komatiitic lavas. Isochron and mantle model ages for peridotitic BMS grains extend to older ages (> 3.2 Ga; Gurney et al. 2010; Shirey and Richardson 2011) than those of eclogitic BMS. To date, no Paleoarchean BMS inclusions with high, eclogitic Re/Os ratios have been found (Shirey and Richardson 2011). Consequently, eclogitic diamonds are only commonplace after 3.0 Ga, which Shirey and Richardson (2011) used to suggest represents the age of the earliest slab subduction and thus the establishment of modern plate tectonics. In their model, eclogitic diamonds usually occur where lithospheric mantle has been subject to continental collision. For example, the collision of two continental blocks of the Kaapvaal craton, the Witwatersrand and Kimberley blocks at 3.1–2.9 Ga (Moser et al. 2001; Schmitz et al. 2004), is suggested to have resulted in underthrusting of oceanic lithosphere and eclogite capture (Shirey and Richardson 2011). The absence of Paleoarchean eclogite xenoliths and eclogitic BMS in diamonds from cratonic lithospheric mantle suggests that the earliest continental nuclei may have formed by processes that were different to what today are recognized as the Wilson cycle. The absence of a true Wilson cycle in the Palaeoarchean does not rule out the presence of some form of recycling or even shallow plate subduction, but the style of subduction observable today had not been initiated at this point. The earliest detected geochemical signatures consistent with some form of recyling of hydrated oceanic lithosphere date back as far as 3.9 Ga and perhaps even 4.2 Ga, based upon the changing Nb/Th (Jochum et al. 1991) and Th/U ratios in basaltic and ultramafic rocks (Collerson and Kamber 1999; Zartman and Richardson 2005), and their trace and major element systematics (O’Neil et al. 2011). Crustal rocks with a signature consistent with derivation from depleted mantle (εNd : εHf = 1:2) became apparent after 3.6 Ga (Shirey et al. 2008). The appearance of subduction-related imprints on the stable keels of cratons supports the hypothesis that the generation of continental crust was progressively depleting the mantle from 3.6 Ga onwards. By the Neoarchean (after ca. 2.8 Ga), recycling processes more strongly resembled the style of subduction consistent with the Wilson cycle (Shirey et al. 2008).

The age of the continental lithospheric mantle and the assembly of its domains The timing of diamond formation and the conditions of their formation are invaluable in unravelling the series of events that together resulted in the assembly of the cratons. Moreover, the geochronological information preserved in BMS grains that become isolated within the

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SCLM as a result of melt depletion can help to constrain the timings of melt depletion and the inherent link between the formation of continental crust and the stabilization of the SCLM. With specific regard to mantle BMS geochronology, three cratons in particular have been studied in detail: the Kaapvaal craton, South Africa, the Siberian craton, Russia, and the Slave Craton, Canada. A detailed discussion of the HSE systematics of cratonic mantle is covered in depth in Aulbach et al. (2016, this volume) and only the role of BMS in the determination of cratonic structure and history will be briefly discussed here. The history of the Kaapvaal craton, from the perspective of BMS inclusions in diamonds, was summarized by Shirey et al. (2004) and more recently by Aulbach et al. (2016, this volume). Briefly, the southern African cratonic keel, comprising the Kaapvaal–Zimbabwe cratons, separated by the Limpopo mobile belt, is underlain by Archean mantle peridotite and subordinate quantities of eclogite, which host multiple generations of diamonds of Archean to Proterozoic age (Kramers 1979; Richardson et al. 1984, 1993, 2001; Navon 1999; Pearson and Shirey 1999; Shirey et al. 2002). Some parts of the craton are underlain by portions of the mantle keel that probably relate to the generation of the oldest crust observed in the Kaapvaal. This hypothesis in entirely consistent with the numerous bulk-rock Re–Os model ages derived from xenolith studies of the region (e.g., Walker et al. 1989; Carlson et al. 1999, 2000; Janney at al. 2010) and more recently in situ Re–Os model ages on silicate-hosted BMS grains derived by LA MC ICPMS (e.g., Griffin et al. 2003a,b). Of the suites of eclogitic diamonds associated with the Kaapvaal craton, four (De Beers Pool, Jwaneng, Koffiefontein, and Orapa) preserve Archean ages that overlap the average Neoarchean Re–Os mantle model ages of cratonic peridotite (Carlson et al. 1999; Irvine et al. 2001; Shirey et al. 2001). The eclogitic BMS grains form an array in 187Re/188Os–187Os/188Os space that highlights two general groups of ages. A large number of inclusions from the De Beers Pool, Jwaneng and Orapa plot on a 2.9 Ga isochron (Richardson et al. 2001), with the remainder displaying more scatter (Shirey et al. 2001). At Koffiefontein, only two inclusions correspond to a ca. 2.9 Ga age, while the majority yield much younger ages of 1 Ga (Pearson et al. 1998). The younger Proterozoic diamond ages vary from locality to locality within the Kaapvaal (Pearson et al. 1998; Richardson et al. 2001; Shirey et al. 2004). The sequence of events proposed for the formation of the Kaapvaal craton begins with formation of the nucleus of the craton at ca. 3.7–3.3 Ga (Aulbach et al. 2009a). This would have been accompanied by hot, large melting intervals, such as those inferred to have been responsible for the 3.5 Ga Barberton komatiites. This process would have left a highly depleted residue with refractory olivine and garnet compositions (Stachel et al. 2003; Stachel and Harris 2008) and a sulfide-free residuum (e.g., Keays 1995). In this hot, volatile-poor peridotitic environment, diamond formation would be difficult. Percolation of S-undersaturated melts during secondary processes at this time may also have worked to strip the SCLM of its sulfide content (Reisberg et al. 2005). Those diamonds that could form may have been able to trap refractory silicate inclusions, but the sulfur-undersaturated nature of the protolith may have impeded the encapsulation of large numbers of peridotitic-BMS grains despite the abundance of peridotitic silicate inclusions (Aulbach et al. 2009a). Continental collision of the eastern and western blocks of the Kaapvaal craton at 2.9 Ga allowed for incorporation of volatile bearing and sulfur rich oceanic lithosphere into the mantle keel peridotite. BMS added at this time were eclogitic and preferentially added in the west, explaining both the timing and distribution of 2.9 Ga eclogitic diamond growth (Shirey et al. 2004a, 2013; Aulbach et al. 2009a). Subsequent repeated episodes of sulfide re-introduction into the lithosphere accompanying metasomatism and marginal subduction are documented by peridotitic BMS grains in mantle xenoliths and xenocrysts (Griffin et al. 2004) and multiple episodes of Proterozoic diamonds with eclogitic sulphides (e.g., Pearson et al. 1998; Richardson et al. 2004; Aulbach et al. 2009b). The few examples of peridotitic BMS-containing diamonds in the Kaapvaal craton have proved to be young. (Pearson et al. 1998; Aulbach et al. 2009b).

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Archean peridotitic BMS inclusion-bearing diamonds are more abundant in the central Slave craton, where diamond formation appears to have been coeval with lithosphere formation (Aulbach et al. 2009a, 2011). The Slave craton comprises a juvenile eastern and an ancient central to western domain that may have become amalgamated as a result of east-oriented subduction at ca. 2.7 Ga ago (Kusky 1989; Bleeker et al. 1999a,b). The formation of shallow lithosphere beneath the Slave craton is thought to have commenced through accretionary processes as early as ca. 3.5 Ga (Aulbach et al. 2011). Several peridotitic diamond-hosted BMS grains derived from Diavik plot along a 3.3 Ga isochron, whereas 11 similar inclusions yield an isochron age of 3.52 Ga i.e., from the Palaeoarchean (Westerlund et al. 2006). This early generation of peridotitic BMS that is predominant in the Slave craton is completely missing in the Kaapvaal craton. In contrast to the Kaapvaal craton, the less depleted deep lithospheric mantle beneath the central Slave craton is proposed to have formed as a result of plume-related underplating, but with a lower overall degree of melt depletion because of the presence of a pre-existing lithospheric mantle lid. This smaller degree of melt depletion did not result in a complete exhaustion of the SCLM sulfur supply and hence permitted the formation of peridotitic BMS inclusions coeval with diamond formation (Aulbach et al. 2009a). However, like the Kaapvaal craton, the subsequent evolution of the Slave craton has been dominated by a series of collision and subduction-related events recorded in the multiple generations of diamond-hosted eclogite- and pyroxenite-hosted BMS (Aulbach et al. 2009a,c).

The relationship between the age of the SCLM and the overlying crust In regions where the overlying crustal record is preserved, ages derived for crust generation from the crust itself can be compared to peaks in model ages preserved in SCLM-derived BMS grains derived from peridotite xenoliths (e.g., Pearson et al. 2007; González-Jiménez et al. 2013, 2014). Where the crustal record is imperfect, for example when ancient crust has been eroded or has become obscured through burial or tectonism, the geochronological evidence preserved in relative probability plots of BMS model ages still provides valuable information regarding the history of crust formation and SCLM stabilization. Peaks in model ages on relative probability diagrams of Re–Os TRD or TMA (e.g., Figs. 18, 20) are attributed significance as corresponding to the stabilization of SCLM, which may have been coupled to major crust formation events. Both bulk-rock and BMS Re–Os ages frequently show a correspondence between the crust and the underlying mantle, sampled as xenoliths, from which it is inferred to be derived (Pearson et al. 1995a,b 2002; Handler et al. 1997; Carlson et al. 1999; Chesley et al. 1999; Menzies et al. 1999; Hanghøj et al. 2001; Irvine et al. 2001 2003; Lee et al. 2001; Griffin et al. 2002; Schmidt and Snow 2002; Carlson and Moore 2004; Marchesi et al. 2010; González-Jiménez et al. 2013, 2014). However, there are some examples where underlying mantle is apparently older than the overlying crust (Parkinson and Pearce 1998; Peslier et al. 2000; Handler et al. 2003; Smit et al. 2010). In the Udachnaya kimberlite, Siberia, Griffin et al. (2002) used sulfide model ages to construct a relative probability plot to date the formation of the Siberian craton, though the BMS enclosed in olivine macrocrysts are complex in nature. Of the 53 BMS grains that had 187 Re/188Os of < 0.07 (the cut-off chosen by Griffin et al. 2002, as ensuring a model age that was not affected by Re-addition or mixing with Re-rich secondary sulfides) 45 of the BMS grains produced TMA ages between 2.5 and 3.6 Ga, and 35 of the grains gave TMA of > 2.8 Ga. This lead Griffin et al. (2002) to conclude that the Siberian SCLM formed during the period 3–3.5 Ga, with a peak in lithosphere stabilization at around 2.9 Ga, and little evidence for significant additions to the lithosphere afterwards. However, the recent Re–Os isotope study of Yakutian diamond-hosted BMS by Wiggers de Vries et al. (2013) demonstrated that the assemblage of the Siberian craton is more protracted than the initial 2.9–3.5 Ga growth phase suggested by Griffin et al. (2002). The ages obtained by Wiggers de Vries et al. (2013) for the different diamond populations demonstrate two major periods of eclogitic and lherzolitic Yakutian

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diamond formation at ca. 2.1–1.8 Ga (Mir, 23rd Party Congress and Udachnaya kimberlites) and at ca.1.0–0.9 Ga (Mir and 23rd Party Congress kimberlites). These correspond to the collision between different terranes of the Siberian Craton during the formation of a Palaeoproterozoic supercontinent at ca. 2.0–1.8 Ga and accretion leading to formation of the supercontinent Rodinia ca. 1.1 Ga. The very radiogenic initial Os isotope ratios of the eclogitic and lherzolitic BMS grains (187Os/188Os = 0.14–2.22) imply the incorporation of radiogenic Os from subducted oceanic lithosphere, a conclusion that is also relevant to the Kaapvaal and Slave cratons. However, more recently, the age of the Siberian craton as a whole has been called into question, as the vast majority of Re–Os dates for the region have been derived from a single kimberlite province, i.e., Udachnaya. Doucet et al. (2015) argue that although the Siberian SCLM contains components of material that retain an Archean signature, a wide range of Re-depletion ages have been derived (3.4–1 Ga and commonly ≤ 2 Ga for peridotite xenoliths), consistent with craton-forming activity not limited to the Archean. Using data derived from a second Siberian kimberlite, Obnazhennaya, Ionov et al. (2015) have recently proposed that the lithospheric mantle beneath the Siberian craton was formed in at least two events, one in the late Archean and the other in the Paleoproterozoic. An example of the utility of relative probability plots in coupling melt depletion and the crustal growth history of non-cratonic SCLM was recently published by González-Jiménez et al. (2013). Here, using a similar approach to that employed by Griffin et al. (2002) for olivinehosted BMS, they determined TRD of BMS grains from peridotite xenoliths recovered from the Calatrava volcanic field, Spain. Their study revealed that episodes of mantle magmatism and/or metasomatism in the Iberia microplate were linked to supercontinent assembly and/or breakup at ca. 1.8, 1.1, 0.9, 0.6, and 0.3 Ga (Fig. 20). In addition, they found that the mantle and crust constituting the Iberian microplate have coexisted since at least Paleozoic–Proterozoic time. Rhenium-depletion ages of 68 carefully screened BMS grains, either possessing 187 Re/188Os < 0.07 (n = 12) or 187Re/188Os > 0.07 but with uncorrelated Re/Os, and hence no systematic ingrowth of 187Os (n = 56), produced peaks on a relative probability plot coinciding with major tectonic events recorded in the crust (Fig. 20). The formation of the Iberian lithosphere at ca. 1.6–2.0 Ga coincides with crust-mantle differentiation that occurred during the assemblage of the Columbia supercontinent (ca. 1.8–1.5 Ga). Peaks in TRD at 1.3–0.9 Ga coincided with the amalgamation of Rodinia at ca. 1.25 Ga and its breakup at ca. 0.75 Ga. Peaks in TRD at 0.6 Ga and 0.3 Ga were interpreted by González-Jiménez et al. (2013) as representing the beginning of continental collisions to form Gondwana and the initiation of the breakup of Pangea respectively (Orejana et al. 2009). These interpretations should be treated with some caution given the caveats associated with interpreting Phanerozoic TRD in peridotites. The interpretation of the older sulfide TRD are broadly consistent with the earlier study of Marchesi et al. (2010), which examined individual BMS grains recovered from the Ronda orogenic massif, Spain. This earlier study also considered the possibility that residual BMS grains with Os model ages of ca. 1.2–1.4 Ga may represent the reworking of older SCLM in the Mesoproterozoic. Critically, if the massif became part of the lithosphere at about 1.2–1.4 Ga, the preferred interpretation of whole-rock Os isotopic data (Reisberg et al. 1991; Reisberg and Lorand 1995), then the residual BMS grains with Os model ages of ~ 1.6–1.8 Ga from Marchesi et al. (2010) could be interpreted as inherited grains that survived within the convecting mantle for hundreds of millions of years. The studies of Marchesi et al. (2010) and González-Jiménez et al. (2013) agree that the Proterozoic Os model ages exhibited by BMS grains in the Iberian SCLM tend to correspond to different stages of generation of crust, which was later recycled in the Gondwana supercontinent. Evidence for ancient lithospheric mantle pre-dating the oldest exposed cratonic crust has also been found in the Kimberley craton, north-western Australia in BMS-bearing diamonds recovered from the ca. 20 Ma Ellendale lamproite pipes (Smit et al. 2010). Lherzolitic

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diamond–hosted BMS grains yield an age of 1426 ± 130 Ma, with an initial 187Os/188Os ratio of 0.1042 ± 0.0034. The upper limit on the 187Os/188Os initial ratio suggests a Re depletion age of 2.96 Ga, indicating the presence of SCLM beneath Ellendale since at least the Mesoarchaean. This is supported by independent evidence for the presence of deep SCLM below the King Leopold Orogen and the Kimberley craton by seismic tomography (van der Hilst et al. 1998; Kennett 2003; Fishwick et al. 2005) and cool cratonic geotherms for the West Kimberley province (Griffin and Ryan 1995). Archean SCLM also extends to the south-east of the Kimberley craton, with mantle xenolith TRD ages as old as the Neoarchean (2.2–2.9 Ga; Luguet et al. 2009) or possibly even older, as suggested by an imprecise Re–Os isochron age of 3.4 Ga (Graham et al. 1999). However, the age of the mantle below the Kimberley craton is significantly older than the oldest exposed crust within the North Australian Craton (ca. 2 Ga; Page et al. 1995; Tyler et al. 1999; Worden et al. 2008). Therefore, despite early Proterozoic convergent margins and new crust formation (Tyler et al. 1999; Griffin et al. 2000), remnants of the underlying continental mantle were preserved (Smit et al. 2010). This indicates that crust and mantle in the Kimberley region were decoupled (Luguet et al. 2009), with preservation of the pre-existing lithosphere during accretion (Smit et al. 2010). The generation of BMS grains during ancient melt-depletion events that no longer represent any known event recorded in the overlying crust is the most likely explanation for some populations of BMS. These could either be related to domains where ancient crust is not exposed, either because it has been removed or because newer mantle, containing an old heterogeneous component has become part of the lithosphere. This is particularly pertinent to the subject of osmium and lead isotope heterogeneity in the oceanic mantle.

The inherent heterogeneity within the oceanic mantle Isotopic studies of oceanic samples have long demonstrated the occurrence of isotopic variations in the mantle that must have been created by geologic processes much older than the 180 My age of the oldest ocean floor (e.g., Gast et al. 1964; Sun and Hanson 1975; O’Nions et al. 1977). In addition, compositional heterogeneities in the oceanic mantle at length scales of thousands of kilometres (e.g., Hart 1984; Meyzen et al. 2007) to sub-km scales have often been reported (e.g., Shirey et al. 1987; Dosso et al. 1999; Meibom et al. 2002; Standish et al. 2002; Warren et al. 2009). These observations contrast with studies of the SCLM, where Re–Os ages for the mantle often show a correspondence to the age of the overlying crust (e.g., Griffin et al. 2002; Pearson et al. 2002), though examples exist of the SCLM having ages older than the overlying crust (e.g., Peslier et al. 2000; and discussion above). Although the Pb isotope composition of the oceanic mantle has been of interest to geochemists for nearly half a century (e.g., Allègre 1969), the role that BMS plays as a reservoir for mantle Pb has been less well explored, despite the early realization that Pb isotope heterogeneity may be present at a small scale (e.g., Saal et al. 1998, and more recently Maclennan 2008). Recent interest in the Pb elemental and isotopic composition of mantle BMS stems from a potential solution to the first Pb paradox (Allègre 1969), which is the observation that global basalts have ubiquitously radiogenic Pb isotope compositions (Hofmann 1997) that fall at much higher 206Pb/204Pb than the geochron. If estimates of the bulk silicate Earth Pb isotope composition are correct (i.e., broadly chondritic for U, Th, and Pb; e.g., Palme and Jones 2003) and that the U/Pb ratio was set at ~ 4.50 Ga when core formation ended (Kleine and Rudge 2011), then there must be a reservoir of unradiogenic Pb. Observations of Pb isotopes in both bulk-rock peridotite and individual BMS grains suggest that the mantle may represent one of the reservoirs for unradiogenic Pb based upon isotopic measurements of peridotites and their minerals (e.g., Malaviarachchi et al. 2008; Burton et al. 2012; Warren and Shirey 2012), and predictions based upon trace element depletion (Godard et al. 2005; Kelemen et al. 2007; Hanghøj et al. 2010).

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The first hint that the unradiogenic Pb reservoir was at least partially hosted in the mantle came from a Pb isotope study of the Horoman peridotite (Malaviarachchi et al. 2008). Here, bulk-rock plagioclase lherzolites were analyzed for their Pb isotope ratios and some plotted to the unradiogenic side of the 4.5 Ga geochron. Among abyssal peridotites, one clinopyroxene analysis out of three samples analyzed has yielded an unradiogenic composition (Warren et al. 2009). To date, further Pb isotope data on pyroxene mineral separates from abyssal peridotites are not available due to the analytical difficulties of working at Pb concentrations of ~ 10 ng·g−1. More recently, Burton et al. (2012) and Warren and Shirey (2012) measured Pb isotope compositions in individual BMS grains extracted from abyssal peridotites recovered from the MAR, Gakkel Ridge and SWIR (Fig. 16). For the MAR, Burton et al (2012) found that 206 Pb/204Pb ratios ranged from 18.4063 to 18.4672 for interstitial BMS, i.e., all radiogenic compared to the geochron. In contrast, enclosed BMS ranged from 16.2276 to 17.7839, i.e., extending to values considerably less radiogenic than the geochron. Similarly, Warren and Shirey (2012) reported BMS 206Pb/204Pb ratios ranging from 17.034 (less radiogenic than the geochron) to 19.640 (considerably more radiogenic than the BMS grains of Burton et al. 2012). As seen in Os isotope ratios of abyssal peridotite BMS grains (e.g., Alard et al. 2005; Harvey et al. 2006), Pb isotope ratios of abyssal peridotite BMS show little systematic covariation with basalt Pb isotopic compositions from the area of the ridge from which they were recovered. For example, BMS grains from Gakkel abyssal peridotites span the full range of Gakkel Ridge basalt compositions (Warren and Shirey 2012; Blusztajn et al. 2014), but do not show the break in Pb isotope ratios at 15° E observed among basalts (Goldstein et al. 2008). The recent discoveries of ancient Pb isotope signatures preserved in BMS grains recovered from abyssal peridotites (Fig. 16) suggests that at least some of the missing unradiogenic Pb resides in the upper mantle. This conclusion is supported by the observation that the Pb isotope composition of bulk-rock peridotites from a variety of continental settings also spans both sides of the geochron (Warren and Shirey 2012). In addition, Os and Pb isotopic compositions are correlated (Fig. 21), with both isotope systems yielding similar model ages. For the combined dataset from the MAR, Gakkel and SWIR, the Pb isotopic composition of mantle BMS grains aligns along the ~2 Ga data array for oceanic basalts and peridotite mineral separates, while the Re–Os datasets plot along the 2 Ga Re–Os reference isochron (Harvey et al. 2006; Burton et al. 2012; Warren and Shirey 2012). This “age”, combined with the observed Pb–Os isotopic correlation (Fig. 21) and a very small (≪ 1 km) length scale of peridotite compositional variability can be explained either by crustal extraction at 2 Ga or an average mixing age for the mantle. A single depletion event that can explain the isotope datasets from three widely separated ridges (Gakkel, SWIR and MAR) seems implausible and is inconsistent with episodes of peak crustal formation at 2.7, 1.0, 0.6, and 0.3 Ga (Condie and Aster 2010). In contrast, the mixing of ancient components back into the mantle by subduction has long been suggested as an explanation for the heterogeneity of the oceanic basalt array (e.g., Chase 1981; Hofmann and White 1982; Zindler and Hart 1986; Kellogg et al. 2007). Analysis of the isotopic composition of individual BMS grains from abyssal peridotites allows mantle heterogeneity to be probed at very small length scales, which is particularly powerful when Pb and Os isotopes can be analyzed in the same grain (Fig. 21; Warren and Shirey 2012). Base metal sulfide grains recovered from the same drill core (Harvey et al. 2006; Burton et al. 2012) or dredge haul (Alard et al. 2005; Warren and Shirey 2012; Blusztajn et al. 2014) have large magnitude Os and Pb isotopic variations, demonstrating that these can be preserved in the oceanic mantle at length scales of significantly less than 1 km. For example, the Os and Pb isotope heterogeneity in BMS from serpentinized peridotite, reported by Harvey et al. (2006) and Burton et al. (2012) came from a single 6 cm section of quarter-core from Hole 1274a on the 15° 20’ N Fracture Zone, MAR. Blusztajn et al. (2014) found considerable Pb isotope variation among BMS grains in some individual samples from the Gakkel Ridge and SWIR,

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19

8

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Figure 21. Co-variation of Pb isotopes with 187Os/188Os. The solid lines are regressions through the data (correlation coefficient given by r). PS86-6-38 is excluded from the regressions due to radiogenic Os, which reflects recent interaction with melts from Bouvet hotspot. After Warren and Shirey (2012).

including one sample with BMS grains that cover 25% of the Pb isotope range of oceanic basalts. Pyroxenes from abyssal peridotites have also been used to show considerable Nd (e.g., Cipriani et al. 2004), Sr (e.g., Warren et al. 2009) and Hf (e.g., Stracke et al. 2011) isotope variability, but these isotopes so far have only been measured in mineral separates and not individual grains. The scale and magnitude of oceanic Pb and Os isotope heterogeneity will remain poorly constrained, especially compared to current knowledge of the SCLM, until a larger dataset for abyssal peridotite BMS becomes available. However, what is clear is that the composition of basalts alone cannot be relied upon to constrain either the extent of mantle depletion or the degree of small scale variations in the mantle from which they were derived. Sulfides indicate the occurrence of Os and Pb heterogeneities in the upper mantle with extreme compositions that are not resolvable in pooled basaltic melts. Either these anomalies reside in BMS within domains that are too refractory to contribute to the melting process or their size is such that more voluminous melts smooth out the most extreme isotopic signatures of the source (e.g., Warren and Shirey 2012).

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CONCLUDING REMARKS AND FUTURE DIRECTIONS Base metal sulfides in mantle peridotite and diamonds preserve evidence for a variety of processes, including melt depletion, re-fertilization, metasomatism, and alteration. The ability of sulfides in mantle rocks to preserve different processes has led to an unraveling of these different generations of processes in bulk mantle samples. Combined with new advances in analytical sensitivity and spatial resolution, an improved understanding of the geological history of mantle samples on all scales is being achieved. The emergence of improved amplifier resistors for TIMS and MC ICP-MS instruments has recently seen a leap in possible analytical resolution for silicates (e.g., Koornneef et al. 2015; Sarkar et al. 2015) and meteoritic material (Peters et al. 2015). It seems likely that these methods will be applied to problems where high precision analyses of micro-metric BMS grain and / or alloys are required in the near future. Base metal sulfides hosted in peridotites and pyroxenites of the oceanic lithosphere carry a memory of melt depletion and melt infiltration that is an evident effect of the plate tectonic cycle of ridge melting and slab subduction. For the first time, isotopic ages can be anchored to their mineral hosts. By looking at primary BMS grains, we can peer through the prevalent alteration that masks much of the work done on the bulk-rock scale. Future work will increase the now sparse sites on the ocean floor that have been studied and will allow us to approach a global understanding of specific episodes of depletion and melting, with the ultimate goal to relate these episodes to specific plate geometries and geodynamic cycles. Geochronology on BMS grains in diamonds has achieved the long-held goal of producing ages for every diamond locality that carries amenable BMS inclusions. With the caveat that all diamonds in kimberlite are xenocrysts, age arrays that have geological significance emerge from the diamond suites in many well-mined kimberlites. These arrays are being used to identify when portions of the deep continental lithosphere have been added and reworked, giving a new and deep dimension—one not evident from the surface—to the evolution of the continents. Future work will more solidly anchor these deep mantle lithosphere ages to crustal tectonothermal histories, so that we can achieve a thorough picture of crustal differentiation through time and the nature of the earliest stable crust forming processes.

ACKNOWLEDGMENTS The authors would like to thank Ambre Luguet, Claudio Marchesi, and Akira Ishikawa for thought-provoking reviews of this chapter and James Day for editorial handling. JH was supported by a NERC Advanced Research Fellowship (NE/J017981/1), a Blaustein Visiting Professorship at Stanford University, and a visiting investigator appointment at the Carnegie Institution, Washington. JMW received support from the US National Science Foundation through grant OCE-1434199; SBS received support from NSF grant EAR-1049992.

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Geology 3:297–302 Szabó CS, Bodnar RJ (1995) Chemistry and origin of mantle sulfides in spinel peridotite xenoliths from alkaline basaltic lavas, Nògràd-Gömör volcanic field, northern Hungary and southern Slovakia. Geochim Cosmochim Acta 59:3917–3927 Thomassot E, Cartigny P, Harris RW, Lorand J-P, Rollion-Bard C, Chaussidon M (2009) Metasomatic diamond growth: A multi-isotope study (13C, 15N, 33S, 34S) of sulphide inclusions and their host diamonds from Jwaneng (Botswana). Earth Planet Sci Lett 282:79–90 Tyler I, Page R, Griffin T (1999) Depositional age and provenance of the Marboo Formation from SHRIMP U–Pb zircon geochronology: implications for the early Palaeoproterozoic tectonic evolution of the Kimberley region, Western Australia. Precamb Res 95:225–243 Usselman TM (1975) Experimental approach to the state of the core: Part I. The liquidus relations of the Fe-rich portion of the Fe–Ni–S system from 30 to 100 kb. Am J Sci 275:278–290 Vakhrushev VA, Prokoptsev NG (1972) Primary magmatic formations in basalts from oceanic crust and inclusions of ultrabasic rocks in basalts. Int Geol Rev 14: 125–135 Vakhrushev VA, Sobolev NV (1973) Sulfidic formations in deep xenoliths from kimberlite pipes in Yakutia. Int Geol Rev 15:103–110 van Acken D, Becker H, Walker RJ (2008) Refertilization of Jurassic oceanic peridotites from the Tethys Ocean— implications for the Re–Os systematics of the upper mantle. Earth Planet Sci Lett 268:171–181 van Acken D, Becker H, Walker RJ, McDonough WF, Wombacher F, Ash RD, Piccoli PM (2010a) Formation of pyroxenite layers in the Totalp ultramafic massif (Swiss Alps)—Insights from highly siderophile elements and Os isotopes. Geochim Cosmochim Acta 74:661–683 van Acken D, Becker H, Hammerschmidt K, Walker RJ, Wombacher F (2010b) Highly siderophile elements and Sr–Nd isotopes in refertilized mantle peridotites - A case study from the Totalp ultramafic body, Swiss Alps. Chem Geol 276:257–268 van Acken D, Humayun M, Brandon AD, Peslier AH (2012) Siderophile trace elements in metals and sulfides in enstatite achondrites record planetary differentiation in an enstatite chondritic parent body. Geochim Cosmochim Acta 83:272–291 van der Hilst R, Kennett B, Shibatani T (1998) Upper mantle structure beneath Australia from portable array deployment. In: Structure and Evolution of the Australian Continent. Doolet J, Goleby J, van der Hilst B, Klootwijk R (eds). Am Geophys Union, Washington, DC, p 39–57 Walker RJ (2009) Highly siderophile elements in the Earth, Moon and Mars: update and implications for planetary accretion and differentiation. Chemie der Erde 69:101–126 Walker RJ, Morgan JW (1989) Rhenium osmium isotope systematics of carbonaceous chondrites. Science 243:519–22 Walker RJ, Carlson RW, Shirey SB, Boyd FR (1989) Os, Sr, Nd, and Pb isotope systematics of Southern African peridotite xenoliths; implications for the chemical evolution of subcontinental mantle. Geochim. Cosmochim. Acta 53:1583–95 Walker RJ, Horan MF, Morgan JW, Becker H, Grossman JN, Rubin AE (2002a) Comparative 187Re–187Os systematics of chondrites: implications regarding early solar system processes. Geochim Cosmochim Acta 66:4187–4201 Walker RJ, Prichard HM, Ishiwatari A, Pimentel M (2002b) The osmium isotopic composition of convecting upper mantle deduced from ophiolite chromites. Geochim Cosmochim Acta 66:329–345 Wang KL, O’Reilly SY, Griffin WL, Pearson NJ, Zhang M (2009) Sulfides in mantle peridotites from Penghu Island, Taiwan: melt percolation, PGE fractionation, and the lithospheric evolution of the South China block. Geochim Cosmochim Acta 73:4531–4557 Warren JM (2015), Global Variations in Abyssal Peridotite Compositions, Lithos, submitted Warren JM, Shirey SB (2012) Lead and osmium isotopic constraints on the oceanic mantle from single abyssal peridotite sulfides. Earth Planet Sci Lett 359–360:279–293 Warren JM, Shimizu N, Sakaguchi C, Dick HJB, Nakamura E (2009) An assessment of upper mantle heterogeneity based on abyssal peridotite isotopic compositions. J Geophys Res 114(B12203) Westerlund KJ, Gurney JJ, Carlson RW, Shirey SB, Hauri EH, Richardson SH (2004) A metasomatic origin for late Archean eclogitic diamonds: implications from internal morphology of Klipspringer diamonds and Re –Os and S isotope characteristics of their sulfide inclusions. S Afr J Geol 107:119–130 Westerlund KJ, Shirey SB, Richardson SH, Carlson RW, Gurney JJ, Harris JW (2006) A subduction wedge origin for Paleoarchean peridotitic diamonds and harzburgites from the Panda kimberlite, Slave craton: evidence from Re–Os isotope systematics. Contrib Mineral Petrol 152:275–294. doi:10.1007/s00410-006-0101-8 White RW (1966) Ultramafic inclusions in basaltic rocks from Hawaii. Contrib Mineral Petrol 12: 245–314

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Wicks FJ, Whittaker EJW (1977) Serpentinite textures and serpentinization. Can Mineral 15:459–488 Widom E, Kepezhinskas P, Defant M (2003) The nature of metasomatism in the sub-arc mantle wedge: evidence from Re–Os isotopes in Kamchatka peridotite xenoliths. Chem Geol 196:283–306 Wiggers de Vries DF, Pearson DG, Bulanova GP, Smelov AP, Pavlushin AD, Davies GR (2013) Re–Os dating of sulphide inclusions zonally distributed in single Yakutian diamonds: Evidence for multiple episodes of Proterozoic formation and protracted timescales of diamond growth. Geochim Cosmochim Acta 120:363–394 Wilshire HG, Meyer CE, Nakata JK, Calk LC, Shervais LW, Nielson JW, Schwarzman EG (1988) Mafic and Ultramafic Xenoliths from Volcanic Rocks of the Western United States. US Geological Survey Professional Paper 1443. p 1–179 Wood SA (1987) Thermodynamic calculations of the volatility of the platinum group elements (PGE): the PGE content of fluids at magmatic temperatures. Geochim Cosmochim Acta 51:3041–3050 Worden K, Carson C, Scrimgeour I, Lally J, Doyle N (2008) A revised Palaeoproterozoic chronostratigraphy for the Pine Creek Orogen, Northern Australia: evidence from SHRIMP U–Pb zircon geochronology. Precamb Res 166:122–144 Workman RK, Hart SR (2005) Major and trace element composition of the depleted MORB mantle (DMM). Earth Planet Sci Lett 231:53–72 Yaxley GM, Kamenetsky V (1999) In situ origin for glass in mantle xenoliths from south-eastern Australia: insights from the trace element compositions of glasses and metasomatic phases. Earth Planet Sci Lett 172:97–109 Yefimova ES, Sobolev NV, Pospelova LN (1983) Sulfide inclusions in diamonds and specific features of their paragenesis. Zap Vses Mineral Ova 112:300–310 Zangana NA, Downes H, Thirwall MF, Hegner E (1997) Relationship between deformation, equilibration temperatures, REE and radiogenic isotopes in mantle xenoliths (Ray Pic, Massif Central, France): an example of plume-lithosphere interaction? Contrib Mineral Petrol 127:187–203 Zartman RE, Richardson SH (2005) Evidence from kimberlitic zircon for a decreasing mantle Th/U since the Archean. Chem Geol 220:263–283 Zindler A, Hart S (1986) Chemical Geodynamics. Annu Rev Earth Planet Sci 14:493–571

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Reviews in Mineralogy & Geochemistry Vol. 81 pp. 651-724, 2016 Copyright © Mineralogical Society of America

Highly Siderophile Element and Os Isotope Systematics of Volcanic Rocks at Divergent and Convergent Plate Boundaries and in Intraplate Settings Abdelmouhcine Gannoun Laboratoire Magmas et Volcans Université Blaise Pascal, CNRS-IRD, BP 10448 63000 Clermont Ferrand, France [email protected]

Kevin W. Burton Department of Earth Sciences Durham University, Science Labs Durham DH1 3LE, United Kingdom [email protected]

James M.D. Day Geosciences Research Division Scripps Institution of Oceanography La Jolla, CA 92093-0244, USA [email protected]

Jason Harvey Institute of Geophysics and Tectonics School of Earth and Environment University of Leeds Leeds, LS2 9JT, United Kingdom [email protected]

Pierre Schiano Laboratoire Magmas et Volcans Université Blaise Pascal, CNRS-IRD, BP 10448 63000 Clermont Ferrand, France [email protected]

Ian Parkinson School of Earth Sciences University of Bristol Bristol BS8 1RJ, United Kingdom [email protected]

1529-6466/16/0081-0011$10.00

http://dx.doi.org/10.2138/rmg.2016.81.11

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Gannoun, Burton, Day, Harvey, Schiano & Parkinson INTRODUCTION

Terrestrial magmatism is dominated by basaltic compositions. This definition encompasses mid-ocean ridge basalts (MORB), which account for more than eighty percent of Earth’s volcanic products and which are formed at divergent oceanic plate margins, as well as intraplate volcanic rocks such as ocean island basalts (OIB), continental flood basalts (CFB) and continental rift-related basalts, and highly magnesian ultramafic volcanic rocks that dominantly occur in Archean terranes, termed komatiites. All of these broadly basaltic rocks are considered to form by partial melting of the upper mantle, followed by extraction from their source regions and emplacement at the Earth’s surface. For these reasons, basalts can be used to examine the nature and extent of partial melting in the mantle, the compositions of mantle sources, and the interactions between the crust and mantle. Because much of Earth’s mantle is inaccessible, basalts offer some of the best ‘proxies’ for examining mantle composition, mantle convection and crust–mantle interactions. By contrast, at arcs, volcanism is dominated by andesitic rock compositions. While some arcs do have basaltic and picritic magmatism, these magma types are rare in convergent plate margin settings and reflect the complex fractional crystallization and often associated concomitant assimilation processes occurring in arcs. Despite the limited occurrence of high MgO magmas in arc volcanic rocks, magmas from this tectonic setting are also important for elucidating the behavior of the HSE from creation of basaltic compositions at mid-ocean ridges to the subduction of this crust beneath arcs at convergent plate margins. The highly siderophile elements (HSE; comprising Re and Au, along with the six platinumgroup elements [PGE] Os, Ir, Ru, Rh, Pt, and Pd) combined with the 187Re–188Os and 190Pt–186Os systems that are embedded within these elements, have found significant utility in the study of basaltic rocks (e.g., Shirey and Walker 1998; Carlson 2005; Day 2013). The greatest strengths of the HSE lies in the fact that they strongly partition into metal or sulfide phases, and so record evidence for processes that are not revealed from other isotope systems commonly used in hightemperature geochemical studies (e.g., He–O-Sr–Nd–Hf–Pb). Partial melting over much of Earth’s geological history has resulted in significant fractionation of the HSE between the mantle and the crust (oceanic and continental). The HSE show contrasting behavior during melting, with the platinum-PGE (PPGE; Pt, Pd), Re, and Au usually behaving as moderately compatible to moderately incompatible elements during melting and crystallization, and the iridium-PGE (IPGE; Os, Ir, and Ru) acting as highly compatible elements (Barnes et al. 1985). The differential response of the HSE to partial melting is demonstrated by differences in both the absolute and relative abundances of the HSE in mantle-derived melts and in residual mantle rocks themselves. High degree melts, such as komatiites (e.g., Puchtel et al. 2009) show a smaller enrichment of PPGE over IPGE than relatively lower degree melts, such as MORB (e.g., Rehkämper et al. 1999; Bezos et al. 2005) (Fig. 1a). Mantle peridotites often show a complementary depletion of PPGE relative to the IPGE that reflects the degree of melt depletion (Fig. 1b), consistent with preferential removal of Re > Au > Pd > Pt > Rh > Ir ≥ Ru ≥ Os (Pearson et al. 2004; Becker et al. 2006; FischerGödde et al. 2011). In the broadest sense, these observations suggest that the HSE in mantle and mantle-derived melts are controlled by both: (i) the degree of melting and; (ii) the mineralogy of mantle rocks. The IPGE are preferentially retained in mantle rocks at low degrees of melting, consequently, low-degree melts such as MORB have relatively low IPGE abundances. Furthermore, because Pt is moderately compatible, Re is moderately incompatible and Os is highly compatible during melt generation, the Re–Os and Pt–Os isotope systems differ significantly from other geologically useful long-lived radiometric systems (e.g., Rb–Sr, Sm–Nd, Lu–Hf, U–Th–Pb), where both the parent and the daughter elements are preferentially concentrated into the melt. In this chapter, we review the distribution of the HSE amongst mantle minerals and their behavior during melting, the HSE abundances and Os isotope compositions preserved at mid-oceanic ridge settings (divergent plate boundaries), intraplate settings, and of magmas formed at arcs (convergent plate boundaries), to examine the behavior of these elements in a range of tectonic settings.

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Figure 1. CI-chondrite normalized HSE abundances in (a) mantle derived melts and (b) primitive upper mantle and residual mantle rocks (oceanic abyssal peridotites). Due to extraction of low melting temperature Cu–Ni sulfide melt, which concentrates Pt and Pd, the HSE patterns of residual mantle rocks are depleted in Re, Pd and Pt. The depletion factor increases with the degree of melting (10 to 40%), and therefore with the amount of magma extracted from the mantle column, due to the concentration of the HSE in monosulfide solid solution (mss) and also to the fact that an increase in the degree of melting decreases the amount of mss remaining in the residual mantle. Mantle derived rocks show the opposite behavior. MORB are IPGEdepleted (Ru, Ir, Os) relative to the mantle composition because base-metal sulfides are not exhausted. In contrast the very high degree of partial melting (>35%) needed to generate the Archean komatiite melts consumed all the base-metal sulfides in the mantle, generating HSE pattern close to the primitive mantle estimate. Data sources: MORB (Rehkämper et al. 1999; Bézos et al. 2005; Gannoun et al. 2007; Jenner et al. 2012; Yang et al. 2013, 2014; Burton et al. 2015); komatiites (Puchtel et al. 2004, 2005, 2009; Connolly et al. 2011); abyssal peridotites (Reisberg and Lorand 1995; Pearson et al. 2004; Harvey et al. 2006; Luguet et al. 2007); primitive mantle (Becker et al. 2006).

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HIGHLY SIDEROPHILE ELEMENT DISTRIBUTION AND BEHAVIOR IN THE UPPER MANTLE Core formation and the late accretion of impactor material The HSE have high affinity for both Fe-metal and sulfide over coexisting silicate minerals or silicate melt. Low-pressure metal–silicate partition coefficients determined experimentally are extremely high (between 104 and 1015) (Kimura et al. 1974; Jones and Drake 1986; Peach et al. 1990, 1994; Fleet et al. 1991, 1996; Borisov et al. 1994; O’Neill et al. 1995; Holzheid et al. 2000; Ertel et al. 2001; Fortenfant et al. 2003; Yokoyama et al. 2009; Mann et al. 2012; Brenan et al. 2016, this volume). Consequently, these elements should have been substantially partitioned into Earth’s metallic core, leaving the silicate mantle effectively stripped of the HSE. Yet, HSE concentrations in Earth’s upper mantle are much greater than predicted from low-pressure experimental data (see Day et al. 2016, this volume). Moreover, their relative abundances display a broadly chondritic pattern, rather than reflecting differences in their respective metal-silicate partition coefficients (Fig. 2). However, the siderophile behavior of some of the HSE may be greatly reduced at high P–T conditions, and on this basis it has been suggested that high-pressure equilibration at the base of a deep molten silicate layer or ‘magma ocean’ on the early Earth, may account for their abundances in the upper mantle (Murthy 1991). High-pressure experiments that simulate the conditions of core formation do indeed indicate that the HSE are less siderophile under these conditions (e.g., Mann et al. 2012). However, the range of HSE partition coefficients, even at elevated P–T conditions, cannot account for either the absolute or relative abundances in the terrestrial mantle, suggesting that high-pressure equilibration was not the dominant process controlling their present distribution. Therefore, mantle HSE abundances have long been taken to suggest that between 0.5% and 0.8% by mass of ‘late accreted’ broadly chondritic material was added to Earth after core formation was complete (e.g., Kimura et al. 1974; Chou 1978). Differing absolute abundances, but similar chondrite-relative HSE abundances have also been inferred for the Moon, Mars and other Earth's mantle (Sun and McDonough, 1995) Earth's mantle (Becker et al., 2006) Predicted composition at 1 atm Predicted composition at 20 GPa Late veneer addition 0.55%

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Figure 2. Highly siderophile elements concentrations, normalized to CI-chondrite (Lodders et al. 2009). Primitive mantle compositions are from Becker et al. (2006) and from McDonough and Sun (1995). Predicted composition of Earth’s mantle as a result of metal–silicate partitioning at low pressure (1 atm) are from Borisov et al. (1994, 1995); Borisov and Plame (1997); Ertel et al. (1999); Ertel et al. (2001); Fortenfant et al. (2003, 2006) and at high pressure (20 GPa) are from Ohtani and Yurimoto (1996); Holzheid et al. (2000); Cottrell and Walker (2006); Ertel et al. (2006); Righter et al. (2008); Brenan and McDonough (2009). The late accretion ‘veneer’ addition calculation uses the average composition of all chondrite groups (Walker 2009; Day et al. 2016, this volume). Note the overlap of primitive mantle estimates with the late accretion addition calculation.

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meteorite parent-bodies (Day et al. 2007, 2010a, 2012, 2016 this volume; Brandon et al. 2012; Dale et al. 2012a; Riches et al. 2012; Day and Walker 2015), suggesting that late accretion was a common phenomenon to terrestrial planets, setting the HSE abundances in planetary mantles. In this way, core formation and late addition of meteorite material are thought to have established the HSE abundance in Earth’s silicate mantle, providing a framework for understanding the long-term effects of mantle melting.

Highly siderophile elements in mantle minerals The behavior of the HSE during partial melting of the mantle is controlled by their distribution amongst sulfides, platinum group metal alloys (PGM) and coexisting silicates and oxides in mantle rocks (see also Harvey et al. 2016, this volume; Lorand and Luguet 2016, this volume; O’Driscoll and González-Jiménez 2016, this volume). Sulfide. In addition to their strongly siderophile (iron-loving) behavior, the HSE are also known to be highly chacophile (sulfur-loving), with sulfide in mantle rocks exerting a dominant control over the behavior of the HSE (e.g., Mitchell and Keys 1981), despite the extremely low abundance of these minerals (the proportion of sulfide in mantle rocks is thought to be in the range 0.0014–0.008%, Luguet et al. 2003). The exact magnitude of partitioning of the HSE between sulfide and silicate, however, remains poorly constrained, with values ranging from 1000 to > 106 (Fig. 3) (Peach et al. 1990, 1994; Fleet et al. 1996; Crocket et al. 1997; Andrews and Brenan 2002a; Gannoun et al. 2004, 2007; Fonseca et al. 2009; Mungall and Brenan 2014). At least some of this variation is likely to relate to compositional variations of sulfide and silicate, or the conditions under which equilibration occurred. Values at the low end of the range are usually found in natural occurrences of glass and sulfide (e.g., Gannoun et al. 2004, 2007), while the highest values are indirect estimates based on alloy–sulfide and alloy–silicate partitioning (e.g., Fonseca et al. 2009). A particular problem with the “indirect” estimates of alloy–silicate partitioning (Fonseca et al. 2009) is that they were determined for Fe and S-free compositions, precluding the possible formation of metal-sulfide complexes (e.g., Gaetani and Grove 1997). Moreover, the solubility of at least some of the HSE is enhanced in sulfur-bearing experiments, relative to sulfur-free experiments (Laurenz et al. 2013), bringing partition coefficients into the range of other experimental estimates (Andrews and Brenan 2002a; Mungall and Brenan 2014). While the differences in partition coefficients that remain still span up to three orders of magnitude, 109

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estimates based on individual experiments or natural coexisting sulfide–silicate show significantly less variation. These data indicate that the PGE (Os, Ir, Ru, Pt, and Pd) partition similarly into sulfide, with only Re showing a significant difference to the other HSE. During mantle melting, sulfide will be removed in the silicate melt, as a function of temperature, fO2, pressure, and the iron content of the melt (Wallace and Carmichael 1992; Mavrogenes and O’Neill 1999; O’Neill and Mavrogenes 2002). Given the estimated sulfur content of both the primitive mantle (~ 250 µg g−1 S; Lorand 1990; O’Neill 1991; Palme and O’Neill 2003) and the depleted mantle (~ 120–150 µg g−1 S; Salters and Stracke 2004), and the relatively low degrees of partial melting required to produce most basalts, it is likely that they leave their source sulfide saturated (that is, sulfide remains as a stable mantle mineral). For example, the low HSE content of some low-degree alkali basalt partial melts can be explained by the presence of residual sulfide in the mantle source, while the high-HSE content of highdegree mantle melts, such as komatiites, can be explained by exhaustion of sulfide in the source. However, sulfide behavior alone cannot account for the systematic depletion of HSE seen in mantle rocks, or the variable HSE content and high Re abundances seen in MORB. Silicate and oxides. Rhenium not only partitions into sulfide, but also into other mantle phases including clinopyroxene, orthopyroxene, garnet, and spinel (Hart and Ravizza 1996; Righter and Hauri 1998; Burton et al. 1999, 2000, 2002; Mallman and O’Neill 2007), particularly under reducing conditions (Mallman and O’Neill 2007). The relatively low partition coefficients for Re between silicate phases and melt, and the much lower coefficient for its partitioning between sulfide and silicate melt compared to other HSE, makes this element moderately incompatible during terrestrial partial melting (Fig. 3). The partitioning of Re into silicate then raises the question of to what degree the HSE may also be incorporated into silicates or oxides in mantle rocks. Overall, natural and experimental data suggest that silicate or oxide phases in the mantle do not exert a strong control on the behavior of HSE during partial melting. Taking estimates of the proportion of silicate phases present in the upper mantle (e.g., Workman and Hart 2005), partial melting of a sulfide-free mantle would yield melts that are slightly depleted in Os, Ir, and Ru, relative to their source. Such a pattern is consistent with that seen for high-degree melts, such as komatiites. Nevertheless, silicate and oxide behavior cannot account for the fractionation of the HSE, particularly the low Os, Ir, and Ru contents, observed in basaltic rocks. Spinel. Empirical estimates of partitioning derived from mineral separates suggest that Os, Ru, and Ir are highly compatible in Cr-bearing spinel, with partition coefficients of up to 150, while Pt and Pd are moderately compatible (Hart and Ravizza 1996; Puchtel and Humayun 2001). Experimental work on spinel–silicate melt partitioning at moderate to high fO2 suggests that for Fe-bearing spinels Ru, Rh, and Ir are all highly compatible with partition coefficients of 20 to > 1000, whereas Pd is barely compatible (Capobianco and Drake 1990; Capobianco et al. 1994, Righter et al. 2004). More recently it has been shown that the partition coefficients for Ir, Rh, and Ru are strongly controlled by the ferric-iron content of the spinels. For Cr-bearing spinels, in which Fe3+ is replaced by Cr3+, partition coefficients for Ir and Rh are much lower, and Pt and Pd are highly incompatible (Brenan et al. 2012). Olivine. Some of the first empirical data for olivine mineral separates were taken to indicate that Os may be compatible in olivine with an inferred olivine–silicate partition coefficient of ~ 20 (Hart and Ravizza 1996). However, other work on separated olivine suggested that Os is highly incompatible in olivine (Burton et al. 1999, 2000, 2002; Walker et al. 1999; Harvey et al. 2010, 2011). At this stage it is not clear whether these variations reflect compositional differences between samples, or simply the presence of micro-nuggets of sulfide or PGM in the separated silicate phase. Experimental work, however, suggests that many of the HSE are weakly compatible or only slightly incompatible, particularly under reducing conditions (Brenan et al. 2003, 2005).

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Figure 4. Mineral–melt partition coefficients of HSE determined by experiment (Brenan et al. 2003, 2005, 2012; Chazey and Neal 2005; Mallman and O’Neill, 2007; Righter et al. 2004) and from natural samples (Hart and Ravizza 1996; Burton et al. 1999, 2000, 2002; Puchtel and Humayun, 2001; Gannoun et al. 2004; Gao et al. 2008; Debaille et al. 2009; Puchtel et al. 2009; Harvey et al. 2010, 2011; Connolly et al. 2011; Jackson and Shirey 2011).

Orthopyroxene and clinopyroxene. Empirical constraints from Hart and Ravizza (1996) suggest that Os may be compatible in orthopyroxene and clinopyroxene (Fig. 4), but other studies yield much lower Os concentrations for these phases (relative to coexisting sulfide or olivine) (e.g., Burton et al. 1999, 2000; Harvey et al. 2010, 2011). Experimental work indicates that Re may be mildly compatible in orthopyroxene and clinopyroxene under reducing conditions (e.g., Mallman and O’Neill 2007), but is incompatible under more oxidizing conditions (e.g., Watson et al. 1987; Righter and Hauri 1998; Righter et al. 2004; Mallman and O’Neill 2007). Platinum and Pd appear to be mildly compatible in clinopyroxene (Hill et al. 2000; Righter et al. 2004). Overall, natural samples and experimental data suggest that silicate or oxide phases in the mantle do not exert a strong control on the behavior of the HSE during partial melting. Taking estimates of the proportion of silicate phases present in the upper mantle (e.g.,4Workman and Figure Hart, 2005) partial melting of a sulfide-free mantle would yield melts that are slightly depleted in Os, Ir, and Ru, relative to their source. Such a pattern is consistent with that seen for highdegree melts, such as komatiites (Fig. 1a). Nevertheless, silicate and oxide behavior cannot account for the observed fractionations of the HSE, and in particular the low Os, Ir, and Ru contents, in basaltic rocks (Fig. 1a). Refractory mantle sulfide. For natural magmatic and experimentally produced sulfide the data suggests that while the HSE are strongly partitioned into this phase there is little fractionation between the elements (with the exception of Re). Mantle sulfides, however, dominantly comprise refractory monosulfide solid solution (MSS) and Cu-rich sulfides, which together control much of the HSE budget of the upper mantle (e.g., Alard et al. 2000). Petrographic observations suggest that MSS often occurs as inclusions trapped in silicate phases, and is characterized by high Os, Ir, and Ru abundances, whereas the interstitial Cu-rich sulfides possess lower Os, Ir,

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Figure 5. CI-chondrite-normalized HSE patterns for refractory mantle sulfides and intergranular Cu-rich sulfides. Reported patterns are a combination of different peridotites (Alard et al. 2000, 2005; Lorand et al. 2001; Harvey et al. 2006). Calculated mixture of residual included sulfide and an appropriate amount of intergranular sulfides produces a primitive mantle-like HSE pattern.

and Ru contents (Fig. 5). The silicate hosted MSS sulfides were interpreted to be the refractory residues of partial melting, and the interstitial sulfides as having crystallized from a sulfidebearing melt. On the basis of these observations it has been argued that the fractionation of the HSE during mantle melting might be accomplished by partitioning between refractory “solid” MSS and liquid sulfide (Bockrath et al. 2004a). However, at mantle temperatures of 1300– 1400 °C and pressures of 5–16 kbar—that is, those appropriate for the generation of MORB (e.g., Klein and Langmuir 1987)—any refractory sulfide is likely to be completely molten well before the peridotitic silicate and oxide phases start to melt (Rhyzenko and Kennedy 1973; Hart and Gaetani 2006). Consequently, two phases of sulfide are unlikely to be stable during the melting that produces MORB, consistent with modeled depletion of mantle peridotites where MSS–sulfide melt partitioning cannot explain the observed variations in Pd, Pt, and Au (FisherGödde et al. 2011). However, under conditions of melting at lower temperatures, for example, due to the presence of volatiles such as H2O and at fO2 lower than that at which sulfide is oxidized to sulfate, MSS fractionation may play a role in generating melts with low Os, Ir, and Ru contents (Mungall 2002; Mungall et al. 2006; Dale et al. 2012b; Botcharnikov et al. 2013). Os–Ir–Ru metallic alloys. Osmium, Ir, and Ru (the IPGE) are not only strongly concentrated in refractory MSS, but also in platinum-group minerals (PGM), which encompass alloys and sulfides where Ru, Os, and/or Ir are the major metallic elements. It is clear from the distribution and absolute concentration of the IPGE in PGM (Fig. 6) that precipitation and accumulation of such phases will have a profound effect on IPGE/PPGE fractionation (Brenan and Andrews 2001). In addition, Pt-rich PGM, such as Pt–Ir alloys, have also been found in upper Figure mantle lithologies (Luguet et al. 2007; Lorand et al. 2010). Palladium-rich PGM also exist, which may contain Pt, but typically Pd combines with bismuth and/or tellurium to form bismuthotellurides which are thought to be indicators of refertilization rather than being residual to melting (e.g., Lorand et al., 2010). Thus, Os–Ir–Ru and, to a lesser extent, Pt can all be retained by PGM during melting, while Pd is not. Some have argued that these alloys may represent material that was once part of the core, either as a result of incomplete segregation of metal to the core, or

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due to the entrainment of outer core material into the mantle at the core mantle boundary (Bird and Weathers 1975; Bird and Bassett 1980; Bird et al. 1999). However, recent experimental data suggests that metal originating in the outer core would possess similar concentrations of Os, Pt, and Re, rather than show an enrichment in Ru, Os, and Ir (Van Orman et al. 2008; Hayashi et al. 2009). The solubility of Os, Ir, and Ru is extremely low in silicate melts (e.g., Borisov and Palme 2000; Brenan et al. 2005). Therefore, it has been argued that Os–Ir–Ru-rich PGM may precipitate directly from a silicate melt, through nucleation on nanoclusters of HSE molecules (Tredoux et al. 1995). Furthermore, on the basis of the high solubility of Ir and Ru in sulfide melts it has been proposed that crystallization of Ru–Ir–Os alloys in the presence of a sulfide liquid is unlikely (Brenan and Andrews 2001). Rather it has been argued that such alloys can only precipitate from a melt that is sulfide-undersaturated (Brenan and Andrews 2001; Andrews and Brenan 20002b; Bockrath et al. 2004b; Barnes and Fiorentini 2008). Together, these observations have been taken to suggest that the relationship between Os–Ir–Ru alloys and refractory sulfides in the mantle is key to understandingFigure the behavior 6 of the HSE during higher degrees of partial melting (e.g., Fonseca et al. 2012), where the removal of sulfur in silicate melts leads to a decrease in the proportion of sulfide in the source. All the while that sulfide remains present the HSE are quantitatively retained, and can reach wt% levels in sulfide. However, as soon as sulfide has been completely dissolved, Os–Ir–Ru–Pt alloys form in response to lowering of fS2 and diminished metal-sulfide complexation in the silicate melt (Fonseca et al. 2012). Effectively, much of the HSE budget of the mantle, with the exception of Re, remains in the mantle until sulfide has been completely removed, after which time Os–Ir–Ru and Pt are hosted by PGM phases rather than being liberated in a silicate melt. This model is consistent with an increasing number of petrographic observations indicating the presence of alloy phases in melt-depleted mantle peridotite (Luguet et al. 2003, 2007; Pearson et al. 2004; Brandon et al. 2006; Kogiso et al. 2008, Lorand et al. 2010, 2013; Fisher-Gödde et al. 2012) The degree of partial melting needed to trigger PGM formation will depend on how much sulfur there is in the mantle source at the onset of melting, and is also a result of the solubility of S being inversely proportional to pressure (Mavrogenes and O’Neill 1999). The mantle that melts to produce MORB is already significantly depleted (e.g., Hofmann 1997), and the melting occurs at relatively shallow levels (e.g., Klein and Langmuir 1987). However, there is considerable uncertainty as to the amount of sulfur in the depleted mantle, with estimates ranging down to ~120 µg g−1 (Salters and Stracke 2004) compared to the concentration in

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primitive “fertile” (unmelted) mantle at 250 µg g−1 (Lorand 1990; O’Neill 1991; Palme and O’Neill 2003). Taking the S content of the MORB source mantle to be 120 µg g−1, then 15% melt extraction is needed to exhaust sulfide from the source, and thereby allow the generation of alloys in the mantle residue (Fonseca et al. 2011, 2012; Mungall and Brenan 2014). While these calculations indicate that even the depleted mantle requires significant degrees of melting to remove sulfide, such melt proportions are well within the range of estimates for the generation of MORB (e.g., Klein and Langmuir 1987). In this case PGM formation in the upper mantle may be a potential cause for the characteristic depletion of Os, Ir, Ru, and Rh, relative to Pt and Pd observed in MORB. The absence of significant fractionation of the HSE in komatiites, considered to represent higher degrees of melting than MORB, suggests that alloys are not stable at the higher pressure and temperature conditions required for the generation of such melts (cf. Mungall and Brenan 2014). Overall, the natural and experimental data for mantle minerals indicates that all the while sulfide is present in the mantle, the HSE are largely retained during partial melting, the exception being Re that is not as strongly incorporated into sulfide, and is relatively soluble in silicate melts. However, if sulfide is removed from the system during high degrees of melting, at the pressure–temperature conditions appropriate for MORB melting, then this will result in the formation of Os–Ir–Ru alloys and/or sulfides.

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The major and trace element variations preserved in MORB indicates that their composition has been extensively modified by fractional crystallization, prior to eruption on the ocean floor (e.g., Klein and Langmuir 1987). The principal silicate phases involved in the fractional crystallization that generates MORB are olivine, plagioclase, and clinopyroxene (e.g., Klein and Langmuir 1987; Grove et al. 1992). In general, the more evolved MORB (that is, those with lower MgO and Ni contents, due to the crystallization and removal of olivine) possess lower HSE contents (Fig. 7). On the basis of early empirical estimates for the partitioning of Os into olivine, this relationship has led some to suggest that the HSE

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are compatible in this phase and removed from the silicate melt. However, as discussed previously, with the exception of Re, there is little evidence to suggest that the HSE are strongly partitioned into olivine, plagioclase, or clinopyroxene (Fig. 4). Most MORB are considered to be sulfur saturated (Wallace and Carmichael 1992) and sulfide is a ubiquitous phase. Nevertheless, even if MORB melts are sulfur saturated at their source, they are likely to arrive at the surface undersaturated, because the sulfur content at sulfide saturation increases dramatically at lower pressures (e.g., Mavrogenes and O’Neill 1999). In this case the only viable mechanism by which MORB melts can become sulfur saturated is through extensive fractional crystallization, driving the residual melt to higher S contents. Therefore, it seems most likely that it is the fractional crystallization of olivine, plagioclase and clinopyroxene that drives the melt to sulfur saturation, resulting in the precipitation of sulfide. Hence, the relationship between Ni (concentrated in olivine) and the HSE (concentrated in sulfide) can be attributed to the coupled crystallization of silicates and sulfide. Sulfide may be present in relatively high proportions in MORB (up to ~ 0.23% by mode, K. Kiseeva, 2015), which strongly incorporates most of the HSE, with sulfide/silicate melt partition coefficients of between 104 and 106. In contrast, Re, while still being compatible in sulfide, has a sulfide-silicate melt partition coefficient at least two orders of magnitude lower than that of the other HSE (DRe ~ 10–103). MORB sulfides have high Os (and other HSE) contents, and low Re/Os relative to their parental melt. Consequently, the effect of sulfur saturation and sulfide crystallization will be to decrease absolute HSE abundances, and to raise Re/Os in the residual melt.

THE 187Re–187Os ISOTOPE SYSTEM AND THE FORMATION OF MID-OCEAN RIDGE BASALT (MORB) Introduction Mid-ocean ridge basalts form by partial melting of the Earth’s upper mantle, and variations in their radiogenic isotope compositions or concentration ratios of incompatible elements are considered to reflect compositional heterogeneity in the mantle source (Tatsumoto 1966; O’Nions et al. 1977; Kay 1985; Hofmann 1997). These compositional variations occur on a variety of scales and tectonic settings, ranging from the global-scale of the so-called DUPAL anomaly (centered on the Indian ocean) (Dupré and Allègre 1983; Hart 1984; Hamelin and Allègre 1985; Hamelin et al. 1986; Michard et al. 1986; Price et al. 1986; Dosso et al. 1988; Mahoney et al. 1989, 1992; Rehkamper and Hofmann 1997; Escrig et al. 2004); to those associated with oceanisland volcanics or near-ridge seamounts (White and Schilling 1978; Zindler et al. 1984; Brandl et al. 2012); to minor pervasive variations within ridge segments of normal MORB (e.g., Hofmann 1997; Agranier et al. 2005). A number of processes have been put forward to account for these compositional variations including variable degrees of mantle depletion by prior partial melting (e.g., DePaolo and Wasserburg 1976; Zindler et al. 1984), the infiltration of silicate melts or fluids (e.g., Green 1971), or recycling of lithospheric material into the mantle (e.g., Hofmann 1997). The 187Re–187Os isotope system, based on the long-lived b− decay of 187Re to 187Os, potentially provides an exceptional tracer of recycled lithosphere in Earth’s mantle. This is because both oceanic and continental crust possess exceptionally high Re/Os (parent/daughter ratios), and develop radiogenic Os isotope compositions over time (e.g., Pegram and Allègre 1992; Shirey and Walker 1998; Hauri 2002). In contrast, portions of the lithosphere have low Re/Os, and evolve to unradiogenic Os isotope compositions relative to that of the primitive upper mantle (PUM) (Walker et al. 1989; Pearson et al. 1995). These distinctive isotope signatures can be readily traced as recycled material if mixed back into the convective mantle. For example, the 187Os/188Os variations seen in HIMU ( = high µ = elevated 238U/206Pb) ocean

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island basalts indicate the presence of material that has evolved over a long-time period with a high Re/Os, consistent with models indicating recycled oceanic lithosphere in the source of these volcanic rocks (Day et al. 2010b; Day 2013). Some of the earliest measurements of 187Os/188Os in MORB also yielded isotope compositions more radiogenic than estimates for the primitive upper mantle (e.g., Martin 1991; Roy-Barman and Allègre 1994) and these were attributed either to contamination by seawater derived Os, or melting of a heterogeneous mantle (e.g., Martin 1991; Roy-Barman and Allègre 1994). The work of Schiano et al (1997) on normal MORB, however, not only indicated relatively radiogenic Os isotope compositions but also that these compositions appeared to co-vary with the Sr–Nd and Pb isotopes of the same samples. For the DUPAL anomaly, radiogenic Os isotope compositions were taken to indicate the presence of mafic continental crust in the mantle source (Escrig et al. 2004). On the other hand, radiogenic 187Os/188Os compositions for MORB from the south Atlantic were attributed to metasomatism of the asthenospheric mantle, and local effects from plume–ridge interaction (Escrig et al. 2005a). At first sight, the data from these studies might be taken to suggest that the Os isotope variations reflect those of the MORB mantle source, rather than a secondary process, and that Os isotopes do indeed act as a sensitive tracer of different recycled or enriched material in the mantle. However, these data also indicate a covariation between the Os isotope composition and the Os elemental abundance in these samples (Schiano et al. 1997; Escrig et al. 2005a). Covariations between Os, Ni, and Mg contents in MORB are most readily explained by fractional crystallization (e.g., Burton et al. 2002), but in this case it is then difficult to attribute the Os isotope variations to a mantle source, leading some to propose that the radiogenic Os isotope ratios reported by these studies must result from seawater derived contamination (e.g., Shirey and Walker 1998; Hart et al. 1999; Standish et al. 2002; PeuckerEhrenbrink et al. 2003). Subsequent work demonstrated that many of the MORB previously analyzed (Schiano et al. 1997; Escrig et al. 2004, 2005a) had been affected by an analytical artefact (Gannoun et al. 2007), nevertheless a number of MORB samples still possessed relatively radiogenic isotope compositions (Gannoun et al. 2004, 2007; Yang et al. 2013). Despite the potential utility of the Re–Os isotope system, in particular for tracing the presence of recycled material in MORB, these studies highlight the particular difficulties of both the measurement and the interpretation of 187Re–187Os isotope data in MORB. Midocean ridge basalts possess extremely low Os concentrations, usually less than 10 parts per trillion (pg g−1) which, not only makes their accurate measurement challenging, but also renders MORB highly susceptible to effects that are rarely seen in lithophile elements isotope systems (such as Rb-Sr or Sm-Nd). Such effects include; (i) radiogenic ingrowth of 187Os, produced from the decay of 187Re over very short periods of time (< 10 kyr), (ii) seawater contamination, both directly, on the sea floor, or indirectly in the magmatic plumbing system, and (iii) sample heterogeneity, due to variable contamination in glass or amongst coexisting magmatic phases or through sulfide nugget effects.

Analytical techniques Osmium has seven naturally occurring isotopes, two of which 187Os and 186Os are the decay products of long lived radioactive isotopes, 187Re and 190Pt. Of these two decay schemes, the Re–Os method has been used as dating tool and geochemical tracer for over four decades (Shirey and Walker 1998). Despite its great potential as a geochemical tool, analytical difficulties initially limited the application of the osmium isotope method, mainly because of the high ionization potential of Os (ca. 9eV). The discovery that a solid Os sample could yield negative molecular ions by conventional thermal ionization (Creaser et al. 1991; Volkening et al. 1991) rendered largely obsolete all the excitation methods for atomic osmium used before (Hirt et al. 1963; Luck and Allègre 1982; Walker and Fasset 1986). In the negative thermal ionization mass spectrometry (N-TIMS) method Os is measured as osmium trioxide (OsO3−) via heating on platinum filaments with an electron donor. A Ba–Na emitter solution is employed to lower

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the work function of the filament, which enhances the emission of negative ions. The formation of the Os oxide species is also advantaged by bleeding oxygen into the source (Walczyk et al. 1991). The ionization efficiency increases significantly with decreasing Os loads and can reach above 30% at the pg Os level (Birck 2001; Gannoun and Burton 2014). Another major problem with Re–Os isotopic analysis has been the chemical behavior of Os in solution because of the numerous oxidation states including the volatile tetraoxide species (OsO4). At present, no single technique is equally applicable to all matrices particularly when organic matter and/or refractory mineral phases are present because the variable oxidation states may inhibit the complete homogenization of Os between sample and spike. High temperature (~ 250 °C) oxidizing digestions using either Carius tubes (Shirey and Walker 1995) or high-pressure asher (HPA) digestion vessels (Meisel et al. 2003) have the merit of dissolving acid-resistant phases such as chromite and noble metal alloys. These methods have been supplemented by employing HF digestion after Carius tube/HPA digestion (e.g., Ishikawa et al. 2014), but with mixed results (Day et al. 2015). However, such techniques can potentially yield high total analytical blanks that can contaminate low-HSE abundance samples, such as MORB. Mid-ocean ridge basalt glass possesses low Os abundances, with some samples in the range of 0.2 and 3 ppt, in which refractory minerals are usually absent. For these reasons low-temperature digestion techniques have been used in preference to other approaches when analyzing Os in MORB. These use HF and HBr in sealed Teflon vessels at temperatures of ≤ 140 °C, followed by extraction of Os in liquid bromine (Birck et al. 1997). Extremely low blanks of < 50 fg of Os have been achieved with this method (Gannoun et al. 2004, 2007). Furthermore, MORB glasses are likely to be completely dissolved in HF–HBr acids mixtures even at room temperature. Mid-ocean ridge basalt sulfide grains can be extracted directly using a magnet and handpicked under a binocular microscope (Gannoun et al. 2004, 2007; Harvey et al. 2006) or removed from hand-polished slabs using a diamond scribe to etch around and under the grains (Warren and Shirey 2012). The grains are weighed, spiked with 185Re–190Os and dissolved with high purity HBr. The Os fraction is then purified using microdistillation (Birck et al. 1997; Harvey et al. 2006; Gannoun et al. 2007). It is also possible to undertake dissolution simultaneously with microdistillation (Warren and Shirey 2012). The purified Re and Os are analyzed by N-TIMS following the method described by Pearson et al. (1998). Osmium analysis in sulfides can also be achieved using in situ laser ablation techniques. The strength of this technique lies in the ability to relate Os isotope information from individual sulfide to their precise spatial and textural setting in the rock (Pearson et al. 2002). However, single-sulfide Os data analyzed by the N-TIMS technique are typically of a much higher precision than in situ analysis (cf. Pearson et al. 1998; Harvey et al. 2006; Gannoun et al. 2007) even for sulfide with low Os contents (i.e., less than 10 µg g−1). Moreover, for in situ analysis, because of the isobaric interference of 187Re on 187Os accurate measurement of 187 Os/188Os is only possible for sulfides with low 187Re/188Os (Pearson et al. 2002). Such conditions are typically only met in the case of mantle sulfides.

Rhenium–Osmium elemental variations in MORB glass The fractionation of Re and Os accompanying the generation of MORB is one of the key processes controlling the distribution of these elements between Earth’s mantle and crust. Osmium behaves as a highly compatible element during partial melting, and is preferentially retained in the residual mantle. Consequently, MORB have much lower concentrations, ranging from 0.18 to 170 pg g−1 (with a mean of 10 pg g−1) than mantle peridotite, ranging from 800 to 13000 pg g−1 (with a mean of 3900 ng g−1). In contrast, Re is moderately incompatible during partial melting and preferentially enters the melt. Accordingly, MORB have high Re concentrations, ranging from 480 to 3000 pg g−1 (with a mean of 1023 pg g−1) compared to 10–450 pg g−1 in mantle peridotite (with a mean of 200 pg g−1) (Fig. 8).

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By comparison, komatiiites have generally much higher Os concentrations, up to 10,000 pg g−1, with a similar range of Re concentrations as MORB. These high Re and Os concentrations are generally attributed to higher degrees of melting. Ocean island basalts (OIB) have Os concentrations that range from 1 to 500 pg g−1, and arc lavas from 0.1 to > 10 pg g−1. The low Os concentration of many arc lavas is likely due to extensive removal during fractional crystallization and, indeed, in cases where basaltic compositions have been sampled, Os concentrations can be greater than 50 pg g−1 (e.g., Woodland et al. 2002; Dale et al. 2012b). The relatively low Re concentration of many arc lavas and OIB was originally thought to reflect differences in the mineralogy of the mantle source or the extent of melting, but it is likely that for many of these samples the low Re concentrations result from volatile behavior during sub-aerial eruption (e.g., Lassiter 2003; Day et al. 2010b; Gannoun et al. 2015b). As outlined previously, the low Os concentration of MORB is likely to result, in part, from preferential partitioning into residual Figuremantle 8 sulfide and/or PGM phases and, in part, to the low solubility of Os in silicate melts. In addition, the Os composition of primitive MORB melts will be further reduced by sulfide segregation during fractional crystallization. In contrast, the relatively high Re concentrations result, in part from Re being much less strongly incorporated in mantle sulfide and PGM phases and, in part, from much of the Re budget being controlled by silicate phases, and having a much higher solubility in silicate melts. Rhenium is removed into both silicates and sulfide during fractional crystallization. A remarkable feature of MORB, and indeed all other terrestrial basalts, is the relatively constant increasing fractionation of Re/Os with decreasing Os content. The values range from mantle Re/Os values of around 0.01 for Os concentrations of 2–7 ng g−1, to Re/Os values of ~1000 for lavas with concentrations of 0.1 pg g−1 (Fig. 9). The systematic nature of this fractionation, suggests either that it is dominantly controlled by a single process, such as mantle melting or fractional crystallization, or else that several process act to have the same effect, for example, fractionation by refractory mantle sulfide and also by sulfide segregation during fractional crystallization.

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Rhenium shows a broad positive covariation with Al2O3 and sulfur consistent with the incompatibility of all these elements during mantle melting (Fig. 10). The positive Re–S covariation might be explained by the fact that both elements will be removed into sulfide during fractional crystallization, resulting in a decreasing S and MgO content during the differentiation of S-saturated MORB (Mathez 1976; Bezos et al. 2005; Ballhaus et al. 2006). Despite significant scatter, Os broadly covaries with Ni in MORB (Fig. 11), consistent with a role for olivine crystallization in Os partitioning. Although previous studies have attributed the Os–Ni covariation directly to the compatibility of Os in olivine (Brügmann et al. 1987; Hart and Ravizza 1995), natural samples and experiments indicate that Os is much less compatible. Burton et al. (2002) have shown that Os is in fact extremely incompatible in olivine. Rather it is the crystallization of olivine that drives the melt to sulfur saturation, which in turn results in sulfide precipitation (in which Os is highly compatible) that is trapped within the olivine as ‘melt inclusions’ (Walker et al. 1999; Burton et al. 2002; Brenan et al. 2003, 2005). In summary, Re and Os display similar overall behavior in MORB from the three major ocean basins. Osmium is highly compatible during melting Figure and fractional 9 crystallization, whereas Re is moderately incompatible

The 187Os/188Os isotope variations in MORB glass The 187Os/188Os isotope compositions for MORB from the Pacific, Atlantic and Indian oceans are shown against the reciprocal of the concentration in Figure 12. Mid-ocean ridge basalts from the three major oceans show a similar range of 187Os/188Os isotope compositions, ranging from 0.126 to 0.148 with a mean value of 0.133 ± 0.009 (2σ st. dev.) (Gannoun et al. 2004, 2007; Yang et al. 2013). There is no overall correlation with Os concentration (cf. Schiano et al. 1997; Escrig et al. 2004), however, in general MORB glasses have Os concentrations in the following order: Indian > Atlantic > Pacific, and those samples with a higher Os concentration have a tendency to possess more radiogenic 187Os/188Os compositions. Comparison of 187Os/188Os with 187Re/188Os on a conventional isotope evolution diagram (Fig. 13) indicates that there is no systematic covariation. The data do, nevertheless, indicate that MORB glasses with lower 187Re/188Os are generally found in the Indian > Atlantic > Pacific. In addition, those samples with the lowest 187Re/188Os tend to possess the most radiogenic isotope compositions.

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With regard to the long-lived radiogenic isotopes of Sr, Nd, and Pb, while the crosslinked data are limited, there are no systematic variations between 187Os/188Os and 87Sr/86Sr, 143 Nd/144Nd, and 206Pb/204Pb (Fig. 14). Similarly, there is no correlation between 187Os/188Os composition and ridge bathymetry or spreading rate (Fig. 15) (using data compilation of DeMets et al. 2010 and Argus et al. 2011).

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Figure 14. 187Os/188Os isotope composition of MORB glasses (Burton et al. 2015; Gannoun et al. 2004, 2007; Yang et al. 2013) shown against (a) 87Sr/86Sr (b) 143Nd/144Nd and (c) 206Pb/204Pb (see text for discussion) (Sr, Nd, and Pb data from Vidal and Clauer 1981; Hamelin et al. 1984, 1986; Hamelin and Allègre Figure 14 1985; Prinzhofer et al. 1989; Dosso et al. 1993; Schiano et al. 1997; Escrig et al. 2004).

Analytical issues associated with MORB Several studies have reported 187Os/188Os data for MORB glass (Schiano et al. 1997; Escrig et al. 2004) that could not be reproduced elsewhere, using lower blank techniques (Gannoun et al. 2007). Comparison of these data shows that for many of the relatively unradiogenic samples there is reasonably good agreement between studies (Figs. 16 and 17) but notably none of the very radiogenic values previously reported were reproduced for the same samples. Such a difference might be attributed either to the nature of the samples or

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Spreading rate (cm/yr) Figure 15. 187Os/188Os isotope composition of MORB glass shown against (a) ridge depth (metres below sea level) and (b) spreading rate (cm/year) (calculated using Argus et al. 2011 and De Mets et al. 2010).

(recent work)

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Figure 16. Comparison of 187Os/188Os for MORB glass investigated previously (Schiano et al. 1997; Escrig et al. 2004) and re-analyzed in Gannoun et al. (2007), unpublished data.

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Os (pg g-1) Figure 17. Osmium abundance shown against the deviation of the measured 187Os/188Os (in %) between recent studies (Gannoun et al. 2007; unpublished data) and earlier work (Schiano et al. 1997; Escrig et al. 2004). The highest deviation in the reported 187Os/188Os is observed for the glass samples with the lowest Os contents.

the methods involved in their preparation for chemistry. The earlier studies used leaching techniques to remove any Fe–Mn oxyhydroxides that may have accumulated on the glass while on the sea floor. Iron-manganese precipitates, if present, are likely to possess a radiogenic Os isotope composition acquired from seawater (187Os/188Os = ~1), therefore they might shift the measured 187Os/188Os to more radiogenic values. However, experiments on some of the same glasses indicate that extensive leaching, with oxalic acid and HBr, yields Figure 17 indistinguishable results to those for the same glass samples simply rinsed in dilute HCl, ethanol and water. Another possibility is that because of the large samples sizes used in the earlier studies, between 1 and 5 g (Schiano et al. 1997; Escrig et al. 2004) compared to 300–500 mg (e.g., Gannoun et al. 2007), phenocrysts possessing radiogenic isotope compositions may have been inadvertently included in the material measured. Likewise, entrainment of included sulfides possessing radiogenic compositions may have the same affect. If the radiogenic 187Os/188Os were due to the presence of entrained silicates or sulfides, then some variation in the parent/daughter ratio might be expected (cf. Fig. 9 of Day 2013). Such heterogeneity is spectacularly displayed in two samples from the same locality in the Indian Ocean, where significant variations in the isotope and elemental composition of MORB glass can be attributed to the variable the presence of sulfide inclusions. However, duplicate and triplicate measurement of eleven of the samples showed no resolvable variation, and there is no evidence for isotope and elemental heterogeneity in any of these glass samples. Therefore, it seems more likely that the difference in measured 187Os/188Os composition is an analytical artefact. One possibility is that this is due to interference from 187 ReO3− on the measured 187OsO3−, although this can be carefully monitored during N-TIMS analysis through the direct measurement of 185ReO3−. More likely is that the earlier data were under-corrected for the total procedural blank during chemical purification. The blanks of the original studies possessed a radiogenic 187Os/188Os composition, and the difference between the earlier data (Schiano et al. 1997; Escrig et al. 2004) and those samples that were re-analyzed increases with decreasing Os concentration in the sample, consistent with increasing contribution from the blank (Fig. 17). Overall, these studies highlight the analytical difficulties of obtaining accurate 187Os/188Os data for MORB glass many of which possess low Os concentrations (i.e., between 0.2 and 5 pg g−1). The origin of Os isotope variations in MORB glass. Notwithstanding any shifts that arise from analytical problems, the data obtained thus far, for all the major oceans, indicates a resolvable variation in the 187Os/188Os isotope composition of MORB, ranging from

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values similar to those expected for the primitive upper mantle (e.g., Meisel et al. 1996) to radiogenic compositions akin to those found in ocean island basalts (e.g., Day 2013). It is unlikely that these data have also been compromised by analytical problems; first, because there is no covariation between the corrected 187Os/188Os and the Os concentration, as might be expected if the blank concentration was not correctly determined. Second, replicates with differing sample weights and subject to different dissolution technique yield indistinguishable 187 Os/188Os values (Gannoun et al. 2007; Yang et al. 2013, Burton et al. 2015). Moreover, those samples with radiogenic 187Os/188Os compositions are actually those with the highest Os concentrations, and therefore would be less susceptible to any blank effect. Finally there is no significant covariation between Os and Sr, Nd or Pb isotopes, as might be expected if the variations were due to compositional heterogeneity in the mantle source. Radiogenic growth of 187Os since MORB eruption. For lithophile elements, such as Sr or Nd, parent/daughter ratios in MORB glass and coexisting silicates are relatively low, consequently shifts in their radiogenic isotope composition are unlikely to have a measurable effect for timescales less than 103 million years (e.g., Hofmann 1997). Therefore variations in Sr or Nd isotope composition preserved in MORB can be attributed to compositional heterogeneity in the upper mantle source (e.g., Hofmann 1997). For the 187Re–187Os system however, silicate phases and glass possess exceptionally high 187Re/188Os (parent/daughter). This then raises the possibility that radiogenic 187Os could be produced in situ from the decay of 187Re over relatively short periods of time (that is a few hundred thousand years or less; e.g., Hauri et al. 2002, Gannoun et al. 2004, 2007). For example, MORB glass possesses 187Re/188Os with values ranging from 30 to 8000 (Gannoun et al. 2007; Yang et al. 2013), and a glass with 187 Re/188Os = 4000 would produce a shift in 187Os/188Os from mantle values of 0.1296 to a value of 0.14 in less than 250 thousand years (Gannoun et al. 2007). This effect is illustrated in Figure 13, where timescales of between 50 ka and > 1 Ma could produce the range of 187Os/188Os preserved in the MORB glasses if they were simply due to the decay of Re. One approach to determining the age of crystallization of the MORB glasses is the measurement of short-lived isotopes of Th–U and Ra in the same samples. Such Th–U–Ra data was obtained for a few MORB glasses spanning much of the observed range of 187Os/188Os compositions for the datasets in Gannoun et al. (2004, 2007). Of those samples measured, if it is assumed that they initially possessed a PUM-like composition at the time of crystallization, then between 700 kyr and 1.25 Myr would be required to generate their given 187Os/188Os isotope compositions. However, the same samples possess 230Th/232Th activity ratios greater than 1, suggesting that they must be ≤ 350 kyr old (that is, the maximum time available before all 230Th has decayed). Moreover, all but one sample has a 226Ra/230Th activity ratio that is also greater than 1, suggesting those samples must be ≤ 8 kyr old. Therefore, for these samples, at least, the radiogenic 187Os/188Os compositions cannot be explained solely as a result of in situ decay of 187Re subsequent to igneous crystallization (Gannoun et al. 2004, 2007). An alternative approach that can be used with phenocryst-bearing MORB samples is to obtain Re–Os isotope data for the constituent phases in MORB, including sulfide, glass, spinel, olivine, clinopyroxene and spinel (Gannoun et al. 2004). If these coexisting phases are in Os isotope equilibrium, then they may yield an isochron that will give the age of crystallization, and the initial Os isotope composition defined by the best-fit line will correspond to that of the mantle source. However, if some of the phases were assimilated from previously crystallized basalts, gabbro (from deeper in the oceanic crust), or contaminated by seawater, then they may possess different isotope information to that of the host glass or other minerals (Gannoun et al. 2004). 187Re–187Os data were obtained for coexisting phases from two MORB samples from the FAMOUS region on the mid-Atlantic ridge (Figs. 18 and 19). These results illustrate the age information that can be obtained from MORB glass and coexisting phases, some of the processes involved in MORB genesis, and

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Figure 18. 187Re–188Os isotope evolution diagram for coexisting phases from the olivine–basalt ARP1974-011-018 (Gannoun et al. 2004). Olivine, plagioclase, glass, and matrix yield a best-fit line corresponding to an age of 565 ± 336 ky (2σ). Clinopyroxene (not shown) does not lie on this best-fit line, suggesting either an older age or a different and more radiogenic source for this phase.

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Figure 18

Figure 19. 187Re–188Os isotope evolution diagram for coexisting phases from the picritic basalt ARP1973-010-003 (Gannoun et al. 2004). Olivine, plagioclase, glass, and sulfide lie on a best-fit line corresponding to an age of 2.53 ± 0.15 My(2σ). Spinel possesses a distinct isotope composition from this best-fit line and is probably the phase responsible for the displacement of the matrix from the same line.

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the 187Os/188Os composition of the MORB source. Sample ARP1974-011-018 (36.85°N; 33.25°W) is an olivine basalt containing olivine (Fo90–Fo80), plagioclase (An91–An95), and clinopyroxene (Wo44En15Fs5–Wo40En15Fs9) phenocrysts (up to 1–2 mm in diameter) and microphenocrysts in a hyalocrystalline matrix, and, in places, a glassy pillow rim (e.g., Le Roex et al. 1981). The 187Re–187Os isotope data for matrix, glass, plagioclase, and olivine yield a best-fit line corresponding to an age of 565 ± 336 ky and an initial 187Os/188Os of 0.1265 ± 0.0046 (Fig. 18). The data for clinopyroxene are distinct from this best-fit line, suggesting either an older age or a different and more radiogenic source for this phase. Sample ARP1973-010-003 (36.8372°N; 33.2482°W; 2760-m water depth) is a porphyritic, picritic basalt with abundant olivine phenocrysts (Fo91–Fo89; up to 5 mm in diameter) set in a glassy to hyalocrystalline matrix. Cr-spinel [Cr/(Cr + Al) = 48.01] phenocrysts and sulfide [~14 weight percent (wt %) Ni] blebs (up to 1 mm in diameter) occur as inclusions in olivine or discrete crystals in the groundmass. Plagioclase (An80) microlites are also common (Le Roex et al. 1981, Su and Langmuir 2003). The 187Re–187Os data for olivine, plagioclase, glass, and sulfide yield a best-fit line corresponding to an age of 2.53 ± 0.15 My and an initial 187 Os/188Os ratio of 0.129 ± 0.002 (Fig. 19). Spinel, which is relatively Os-rich (Table 1 of Gannoun et al. 2004), possesses a distinct isotope composition from this best-fit line and is probably the phase responsible for the displacement of the matrix from the same line. The simplest interpretation of these data is that the ages represent the time of igneous crystallization and the initial Os isotope composition represents that of the mantle source. The crystallization ages are, however, much older than might be expected from age-distance relations with the ridge axis that suggest ages of 5–10 kyr (Selo and Storzer 1979). They are also different to the ages inferred from the Th–U–Ra isotope composition of the glass. Glass from sample ARP1974-011-018 glass gives a 226Ra/230Th activity ratio close to 1, suggesting that the sample is ≤ 8 ky old, whereas the 230Th/232Th activity ratio is 1.273, suggesting that the sample is ≤ 350 ky old, consistent with previous 230Th data for the same sample (Condomines et al. 1981). Arguably the 187Re–187Os age of 565 ± 336 kyr is indistinguishable from the 230Th age constraints. Glass from sample ARP1973-010-003 gives 226Ra/230Th ratio of 1.3, which might at first be taken to indicate that the sample is less than 8 ky old. However, the same sample has a 234U/238U ratio of 1.043, and such elevated values are often taken to indicate seawater contamination, consistent with previously published data for this sample (Condomines et al. 1981), which raises the possibility that Ra has also been affected by the same seawater contamination. It might be argued that the best-fit lines are due to contamination by radiogenic Os from seawater, rather than having some age significance. This would require that the contamination occurred during mineral crystallization and has affected phases such as olivine and plagioclase in a systematic manner; otherwise, it is difficult to imagine how different phases would align to yield the correlations observed. Alternatively, the data may indicate that few if any of the constituent phases crystallized in their present basalt host (i.e., they are xenocrysts not phenocrysts). There is evidence for assimilation of xenocrystic phases in samples from the FAMOUS region (e.g., Clocchiati 1977; le Roex et al. 1981; Shimizu 1998). For example, in this sample high-Al spinel is considered to be a relict from high-pressure crystallization (Sigurdsson and Schilling 1976), which suggests that spinel is not in Os isotopic equilibrium with the other phases. However, if most of the phases lie on the same best-fit line, then this interpretation demands that all such minerals are xenocrysts. For the picritic basalt, if eruption occurred about 5–10 kyr ago, then the Re–Os isotopic data indicate that original crystallization of the minerals occurred about 2.5 Myr prior to this event. In this case, the xenocrysts were assimilated from previously solidified “olivine–plagioclase” basalts, or cumulates through which the present host basalts have ascended. Taken together, these results demonstrate that the radiogenic 187Os/188Os composition of MORB glass can be readily generated from the decay of 187Re over very short timescales (that is, a few hundred thousand years or less). Nevertheless, the ages obtained for the samples

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from the FAMOUS region on the mid-Atlantic ridge are much older than might be expected on the basis of their distance from the ridge axis, and this can only be explained either by seawater contamination (that occurred during the crystallization of magmatic minerals) or by the entrainment of crystals (i.e., xenocrysts) from older oceanic crust. Extreme 187Os/188Os heterogeneity in MORB glass. Occasionally MORB itself shows significant Os isotope and elemental heterogeneity. For example, replicate measurements of the MORB sample EN026 10D-3 show significant heterogeneity, with 187Os/188Os isotope compositions that range from 0.128 to > 0.15 (Day et al. 2010b). For MORB glass this is exemplified by two samples from the same locality on the central Indian ridge, MD57 D9-1 and D9-6 (8.01°S; 68.07°E) which show 187Os/188Os compositions ranging from 0.126 to 0.254, with covariations in Os concentration (Fig. 20). Those samples with the least radiogenic 187Os/188Os composition possess unusually high Os concentrations (up to 220 pg g−1). Sulfides from the same samples possess 187Os/188Os between 0.126 to 0.132, and concentrations between 136 and 246 ng g−1. Given the presence of Os-rich sulfide in these samples, it seems most likely that this heterogeneity is due to the entrainment of this phase. If the radiogenic 187Os/188Os isotope composition of the glass is simply due to the radiogenic growth of 187Os from the decay of 187Re since the time of igneous crystallization, then the initial ratio determined from elemental or parent/daughter ratios may reflect the composition of the source (cf. Day 2013). Alternatively, if the radiogenic composition of the glass is due to seawater contamination or altered oceanic crust then the initial 187Os/188Os isotope composition determined from such covariations may have little relationship with that of the mantle source.

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1/[Os] (ng g-1) Figure 20. 187Os/188Os versus 1/[Os] for heterogeneous Indian MORB. Two samples from the central Indian ridge, MD57 D9-1 and D9-6 show high range of 187Os/188Os from 0.126 to 0.254 which covaries with Os concentrations (unpublished data).

Seawater contamination or assimilation of altered oceanic crust. The age constraints from spreading rates, Th–U–Ra disequilibria and 187Re–187Os isotope data for MORB glass and coexisting minerals suggest that the radiogenic 187Os/188Os compositions of MORB glass cannot be solely explained by an age effect following igneous crystallization. An alternative

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unpublished work.

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possibility is that these radiogenic compositions could be due to seawater contamination, either occurring directly during quenching of the glass on the ocean floor or through the assimilation of hydrothermally altered oceanic crust in the magmatic plumbing system. Seawater possesses a radiogenic 187Os/188Os composition (~1.026–1.046) (e.g., Sharma et al. 2012, Gannoun and Burton 2014) and a 187Re/188Os ratio of ~3400, (calculated using the Os concentrations from Sharma et al. 2012, Gannoun and Burton 2014 and Re from Anbar et al. 1992, Colodner et al. 1993). In this case, seawater contamination could account for both the radiogenic Os isotope composition and the tendency of such samples to possess relatively low 187Re/188Os. Trace elements that are enriched in seawater, such as Cl or B could potentially be used as indicators of seawater contamination. At first sight, however, there is no apparent covariation of either B or Cl with 187Os/188Os in the MORB glasses. Rather the variations that do exist indicate Figure that many of the samples with radiogenic Os compositions possess low Cl and B concentrations, inconsistent with seawater contamination (Fig. 21). The difficulty in interpreting Cl and B is that both are highly incompatible elements, and therefore they are strongly affected by partial melting Figure and fractional crystallization (Michael and Schilling 1989; Chaussidon and Jambon 1994; Jambon et al. 1995; Michael and Cornell 1998). Indeed, Cl and B for the same MORB glasses show a negative covariation with MgO suggesting that fractional crystallization has strongly influenced their abundances, thereby masking any subtle effects from seawater contamination. Like Cl and B, K also behaves as a highly incompatible element during melting and crystallization, in this

21 21

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case an alternative approach is to use incompatible element ratios such as B/K or Cl/K that are not significantly fractionated during crystallization to place some constraints on potential contamination by seawater. For example, mantle Cl/K ratios are low (< 0.08), whereas altered oceanic crust has a Cl/K ratio ~ 0.1, and seawater ~ 50 (Michael and Schilling 1989; Jambon et al. 1995; Michael and Cornell 1998). However, again there is no clear co-variation of Cl/K with 187 Os/188Os, rather the radiogenic Os values appear to possess low Cl/K (Gannoun et al. 2007). A more robust tracer of seawater interaction is provided by 11B/10B of the MORB glasses. The upper mantle is thought to possess a δ11B value (Chaussidon and Marty 1995) of −10‰, (where δ11B = 1000 × [(11B/10Bsample /11B/10Bstandard) − 1] relative to the borate standard NBS 951 with an 11B/10B ratio of 4.04558). In contrast, for altered oceanic crust δ11B ranges from +2 to +9‰, seawater has a δ11B = +39.5‰ (e.g., Spivak and Edmond 1987; Smith et al. 1995) and serpentinized oceanic mantle samples can range from +9‰ to +39‰ (Boschi et al. 2008; Vils et al. 2009; Harvey et al. 2014a). While melting and crystallization processes are unable to significantly fractionate boron isotopes, mixing with altered oceanic crust and mantle can account for the δ11B range of −7 to −1‰ observed in MORB (Chaussidon and Jambon 1994). The δ11B values of MORB glasses for which 187Os/188Os data are available range from −9 to +2‰, and those samples with high δ11B values also possess radiogenic 187Os/188Os compositions (Fig. 22). The B concentration of seawater is ∼ 4.6 µg g−1 which is some 5–10 times higher than that of unaltered MORB ( −5‰), and the radiogenic Os/188Os could be due to the presence of recycled oceanic crust (present as pyroxenite) in the MORB mantle source. Recycled oceanic crust can lose substantial amounts of Re during subduction (∼ 50% or more, Becker 2000; Dale et al. 2007) but Re/Os ratios are still sufficiently elevated to produce radiogenic 187Os/188Os values with time. However, recent studies suggest that during dehydration of the subducting slab, B is preferentially partitioned into the released fluids, leaving a depleted residue (Moran et al. 1992; Bebout et al. 1993; Peacock and Hervig 1999; Nakano and Nakamura 2001; Harvey et al. 2014b). Furthermore, boron-isotope fractionation occurs during such dehydration and the residue becomes increasingly enriched in the light B isotope (10B) generating light δ11B values (You et al. 1996; Ishikawa et al. 2001; Leeman et al. 2004; Dale et al. 2007), rather than the heavy values required to generate the ranges observed in MORB. 187

Notwithstanding analytical difficulties, the Os isotope and elemental variations in MORB glass, the mismatch in age constraints and measured 187Os/188Os compositions, and the covariations with B isotopes suggest that assimilation of seawater-altered oceanic crust is likely to be the dominant process responsible for the radiogenic Os-isotope signal seen in many of the MORB glasses studied thus far.

SULFIDES IN MID-OCEAN RIDGE BASALTS Petrology and chemistry Sulfide is a ubiquitous phase in MORB glass, indicating that these melts were sulfur saturated (Wallace and Carmichael 1992). Because decompression will drive the melt away from sulfide saturation (e.g., Mavrogenes and O’Neill 1999) it might be expected that most MORB would be undersaturated when transported to lower pressures during eruption. The presence of sulfide globules in early crystallizing phases, however, clearly indicates that MORB are sulfur saturated during the initial stages of magmatic evolution (Mathez and Yeats 1976; Patten et al. 2012; Yang et al. 2014) and, as previously suggested for MORB, this sulfur saturation is most likely to result from fractional crystallization itself. In addition, MORB contain more sulfur than subaerially erupted basalt, because degassing is impaired by the overlying pressure of seawater. Sulfides occur as spherules embedded in the walls of large vesicles (Moore and Calk 1971; Moore and Schilling 1973), as small irregular grains in microcrystalline aggregates of plagioclase and olivine (Mathez and Yeats 1976) and as well-developed spherical globules, in glass or in phenocrysts (Mathez and Yeats 1976; Czamanske and Moore 1977; RoyBarman et al. 1998; Patten et al. 2012) (Figs. 23a,b). The globules, which range from 5 to 600 μm in diameter, have different textures that can be divided in three groups (Moore and Calk 1971; Mathez 1976; Mathez and Yeats 1976; Czamanske and Moore 1977; Peach et al. 1990; Roy-Barman et al. 1998; Patten et al. 2012, 2013). The first, comprise a fine grained micrometric intergrowth of Fe–Ni-rich and Cu–Fe-rich sulfide phases that represent quenched monosulfide solid solution (MSS) and intermediate solid solution (ISS). The second, comprise globules of coarser grained intergrowth of MSS and ISS with pentlandite and oxide (Mathez 1976; Czamanske and Moore 1977; Patten et al. 2012) and the third group comprise zoned globules that consist of two massive and distinct grains of MSS and ISS, first identified recently by Patten et al. (2012).

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ISS

Pn MSS

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Figure 23. Backscattered-electron (BSE) images and chemical maps of typical MORB sulfides from the picritic basalt ARP1973-010-003 (Famous area, Mid-Atlantic ridge). Chemical maps were produced using a wavelength dispersive spectrometry (WDS) coupled to a CAMECA SX-100 microprobe at Blaise Pascal University (Clermont-Ferrand, France). Shading indicates the relative abundance of a given element. MSS: monosulfide solid solution; ISS: intermediate solid solution; Pn: pentlandite. a. spherical sulfide globule inclusion in olivine. (continued on next page).

Pentlandite and oxide occur to a lesser extent in all types of textures. Sulfide droplets with different sizes and textures may coexist in the same MORB sample. Patten et al. (2012) have shown that sulfide droplets exhibiting all three textures may be present in the same sample separated by only few mm, (cf. Czamanske and Moore 1977). Patten et al. (2012) also observed a relationship between the size of the droplets and their textures. Below 30 μm, over

Figure 23a

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MSS Pn ISS

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Figure 23 (cont’d). b. sulfide globule inclusion in basalt matrix. Both grains have coarse grained texture.

90% of the droplets have a fine grained texture and between 30 and 50 μm, 60% of the sulfide droplets are coarse-grained. In contrast, above 50 μm all the droplets are zoned. Sulfide globules usually comprise fine-grained exsolution of Fe-Ni and Cu-rich sulfide phases. When the bulk compositions of sulfide are calculated to 100%, in order to estimate liquidus temperature of the MSS using the Ebel and Naldrett (1997) approach for O-free systems, they showed low variability in S content, moderate variability in Fe contents and high variability in Cu and Ni contents (Patten et al. 2012). Figure 24 shows the bulk composition of sulfide globules in terms of the system Fe–Ni–Cu. The limited field of such bulk compositions confirms the agreement between different studies (Czamanske and Moore 1977; Roy-Barman Si The dashed lines in Figure 24 indicate the sulfide liquid at et al. 1998; Patten et al. 2012). crystallization temperatures of the MSS at 1100, 1050, and 1000 °C from Ebel and Naldrett Figure 23b (1997). The liquidus temperature of the sulfide globules from MORB determined in this way, range from slightly above 1100 °C to 1030 °C where globules are randomly distributed over this temperature interval irrespective of their size or textures (cf. Patten et al. 2012).

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Roy-Barman et al., 1998 0 Yang et al., 2014 Czamenske and Moore, 1977 0.1 Patten et al., 2013 Gannoun, unpublished Gannoun, unpublished 0.2

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Figure 24. Bulk composition of sulfide droplets in the system Fe–Ni–Cu in weight fraction. The grey zone corresponds to the bulk composition of sulfide droplets from Czamanske and Moore (1977). Dashed line represents the composition of sulfide liquid composition at Mss crystallization at 1100, 1050, and 1000 °C from Ebel and Naldrett (1997). Note that texture of sulfide droplets is not dependent on their composition. Droplet liquidus range between more than 1100 to 1050 °C. Modified from Czamanske and Moore (1977) and Patten et al. (2012).

Pentlandite occurs to a lesser extent than MSS and ISS in all textures of sulfides. Oxide also occurs either inside MSS, inside ISS or at their interface, comprising up to 7% of some sulfide globules. Oxides are best developed in zoned droplets and electron probe analyses reveal that they are Ti-free magnetite (Patten et al. 2012) in agreement with Czamanske and Moore (1977), who suggested that a few percent of magnetite is common in sulfide globules in MORB. 187

Re–187Os behavior in MORB sulfide

Figure 24

If present, sulfide dominates the Os budget in MORB, where sulfide-silicate partition coefficients for Os in basaltic system are in the range ∼104–106 (Roy-Barman et al. 1998; Gannoun et al. 2004, 2007). In contrast, Re while still being highly compatible in sulfide, has a partition coefficient at least two orders of magnitude lower than that of Os (∼ 101–103; RoyBarman et al. 1998; Gannoun et al. 2004, 2007), and similar to that of Cu (Peach et al. 1990; Gaetani and Grove 1997). As a result of the difference in partitioning of Re and Os, MORB sulfides have high Os concentrations (tens to a few hundreds of ng g−1) and low Re/Os relative to their coexisting glass (some three orders of magnitude lower). Consequently, sulfide is much less susceptible to the effects of seawater assimilation, or radiogenic in-growth, than coexisting silicate minerals or glass (Roy-Barman et al. 1998; Gannoun et al. 2004, 2007). For those sulfides for which Os isotope and elemental abundances have been measured thus far, there is a clear covariation between 187Os/188Os and the Os concentration (Fig. 25). Where those sulfides with low Os concentrations (i.e., ≤ 10 ng g−1) possess 187Os/188Os compositions > 0.15, and those with high Os concentrations (i.e., ≥ 100 ng g−1) possess 187Os/188Os compositions around ~ 0.13 or less. This relationship might be taken to indicate that the sulfide globules, like their host glass have been systematically affected by contamination with material derived from

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Figure 25. 187Os/188Os isotope composition shown against Os concentration (ng g−1) for individual sulfides from MORB. This indicates a negative covariation between 187Os/188Os and Os concentration in the sulfides, where low Os sulfides possess more radiogenic Os isotope compositions. These radiogenic values may indicate that such sulfides are more susceptible to seawater derived contamination. Data taken from Gannoun et al. (2004, 2007), Roy Barman et al. (1998), and unpublished data (see text for discussion).

altered oceanic crust. There is no clear relationship between the Os concentration of the sulfide and that of the host glass. However, with one exception, sulfides possess 187Os/188Os values that are less radiogenic than their glass host, where in general, the more radiogenic the host glass the greater the difference in 187Os/188Os with coexisting sulfide (Fig. 26). It is difficult to explain such a difference between sulfide and glass simply by radiogenic decay of 187Re. Rather it suggests that the 187Os/188Os composition of the glass has been more significantly affected by the assimilation of older oceanic crustal material than the coexisting sulfide. If MORB sulfides preserve 187Os/188Os compositions that are systematically less radiogenic than their host silicate glass then this has some important implications for the timing of contamination relative to crystallization. If contamination of the silicate melt occurred before sulfide precipitation then the sulfide should possess an Os isotope composition that is indistinguishable from that of the melt. Therefore, the contrasting Os

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Figure 26. 187Os/188Os isotope composition of individual sulfides shown against the 187Os/188Os value of the glass host. In all cases, sulfide grains possess 187Os/188Os values that are less radiogenic than their glass host. Sulfides also show a much reduced range of Os isotope compositions compared to the corresponding host glass. Data from Gannoun et al. (2007) and unpublished work.

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isotope composition of the glass and sulfide suggests that the silicate melt experienced contamination after the segregation of sulfide in the melt. At the high temperatures of MORB eruption (~ 1200 °C) most sulfides will be present as liquid globules rather than as a solid phase, and diffusional equilibration between silicate and sulfide liquids is likely to be rapid. The time in which a sulfide globule will equilibrate its Os-isotope composition with a melt can be assessed using simple diffusion calculations. Using an implicit finite difference model (Crank 1975) and assuming a sulfide globule radius of 250 μm and a silicate-sulfide melt diffusion coefficient of 10−8 cm2 s−1 the sulfide will equilibrate with the melt in ∼12 h (Gannoun et al. 2007). This is a relatively conservative estimate because cation diffusion in most basaltic melts is 10−5 to 10−6 cm2 s−1 (Watson and Baker 1975), whereas diffusion rates in pyrrhotite are likely to be faster than 10−9 cm2 s−1 at magmatic temperatures (Brenan et al. 2000). Therefore, under normal circumstances, complete equilibration between sulfide and glass would be expected, with both possessing an indistinguishable 187Os/188Os composition. However, because of the large concentration difference between the sulfide and the silicate liquid, a large amount of melt has to exchange with a small sulfide bleb before the sulfide reaches Os isotope equilibrium with the glass. It is possible to calculate the volume (and mass) of melt that is needed to equilibrate the sulfide using simple mass balance equations and the concentration and isotopic data for the glass and sulfides obtained here. Assuming initial 187 Os/188Os for the sulfides of 0.125 and a sulfide globule radius of 250 μm, then sulfides will have only equilibrated with < 0.5 cm3 of melt (or less if the sulfide blebs were smaller). This suggests that the sulfides have only exchanged with the immediate melt surrounding the sulfide. Furthermore, a sulfide that contains > 200 ng g−1 Os would have to exchange with < 50 cm3 of melt in order to completely equilibrate with that melt. Thus, the absence of any Os isotope or elemental covariation between the sulfides and their host glass suggests that Os isotope exchange is likely to have been limited. These observations are consistent with Pd elemental data for MORB from the south west Indian ridge taken to suggest that segregated sulfides were poorly equilibrated with their host silicate magmas (Yang et al. 2013). Nevertheless, many, if not all of the sulfides analyzed thus far are likely to have been modified by contamination, depending on their Os concentration. The sulfides with 187Os/188Os compositions > 0.13 have most likely been significantly modified through partial exchange with the contaminated silicate melt. Although those sulfides with a high Os concentration (> 20 ng g−1) may have also been affected by such exchange they do, however, yield the least radiogenic compositions yet observed in normal MORB samples.

The 187Os/188Os composition of the MORB mantle source The MORB glass measured thus far preserves variations in 187Os/188Os extending from unradiogenic values as low as 0.125, comparable to estimates for the primitive upper mantle, to radiogenic values up to 0.25. There are no clear covariations with lithophile element isotopes, such as Sr or Nd, as might be expected from Os isotopic heterogeneity inherited for a mantle source. Rather, the radiogenic Os isotope compositions show a relationship with B isotopes that is most simply attributed to seawater-derived contamination that occurs during magma ascent. In this case, to a greater or lesser extent all MORB glass has been affected by seawater contamination. Individual sulfide grains appear to provide a much more robust record of the primary Os isotope signature (Roy-Barman et al. 1998; Gannoun et al. 2004, 2007) although even this phase appears to be susceptible to seawater contamination. In this case it is difficult to assess the extent to which any radiogenic signal, preserved in either glass or sulfide, is due to an age effect caused by 187Re decay following igneous crystallization, or the presence of Re-enriched material, such as recycled oceanic crust in the MORB source. Assuming that the Os isotope information preserved by high-Os sulfide grains has been minimally affected by seawater contamination then they potentially provide some unique

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constraints on the nature of the MORB source. A fundamental assumption underlying the use of radiogenic isotopes, such as Sr, Nd and Os, in mantle derived basalts is that they are in equilibrium with their mantle source (e.g., Hofmann and Hart 1978). Abyssal peridotites are ultramafic rocks thought to represent the residue of the melting responsible for generating MORB (Dick et al. 1984; Johnson and Dick 1992; Brandon et al. 2000). Consequently during melting and basalt genesis the composition of long-lived isotopes of heavy elements in both MORB and residual abyssal peridotites should be the same. The average 187Os/188Os composition of a compilation of abyssal peridotites is 0.127 ± 0.015 (n = 129) (Fig. 27), however, like MORB, abyssal peridotites are also susceptible to seawater alteration during their exhumation on the sea floor, which may shift the composition towards radiogenic values. In this case individual abyssal peridotite sulfides are likely to yield a more reliable indication of their primary Os isotope composition, and these yield an average 187Os/188Os composition of 0.125 ± 0.021 (n = 63). The best estimate for the 187Os/188Os composition of the primitive upper mantle, that is a theoretical mantle composition with high Al2O3 that is considered to have experienced no depletion through melting, is 0.1296 ± 0.0009 (2σ; n = 117) (Meisel et al. 2001). By comparison, the high-Os (> 20 ng g−1) sulfides yield an average composition of 0.129 ± 0.005 (n = 31) with values as low as 0.1236 (Fig. 27). Therefore, these high-Os sulfides show no evidence for significant Re enrichment in the MORB source, as might accompany the presence of recycled oceanic crust. Rather they indicate that the upper mantle source of these samples has experienced a long-term depletion of Re, similar to that observed in abyssal peridotites, and consistent with the incompatible nature of this element during mantle melting.

LOWER OCEANIC CRUST The oceanic crust comprises some 1–1.5 km of basalt and dolerite that is underlain by 4–5 km of gabbro. Therefore, MORB are thought to be evolved lavas formed by fractional crystallization in the lower oceanic crust, that itself comprises plutonic rocks and cumulates from primitive magmas. Given that Re is moderately incompatible while Os is compatible during mantle melting, one might expect that gabbros in the lower crust would have higher Os and lower Re concentrations and accordingly lower Re/Os ratios than evolved MORB, assuming that the phases that control solid/liquid partitioning of Re and Os during crystallization are similar to those involved during partial melting. Gabbroic lower oceanic crust should therefore dominate the HSE budget of the oceanic crust as whole. However, the first reported siderophile element data for gabbros from Ocean Drilling Program (ODP) Site 735 (Blusztajn et al. 2000) yielded rather low HSE concentrations (Fig. 28)—even lower than average MORB (Bézos et al. 2005; Gannoun et al. 2007)— pointing to their evolved compositions. Indeed, Dick et al. (2000) and Hart et al. (1999) noted that the average composition of gabbro from ODP Site 735B is closer to that of average MORB (on the basis of major and trace element systematics). Consequently, the gabbro recovered at this site cannot be considered as the primitive complement to typical evolved MORB. More recently, Peucker-Ehrenbrink et al. (2012) have argued that all prior geochemical work on in situ upper oceanic crust such as DSDP-ODP sites 417, 418, and 504 (Bach et al. 2003; Peucker-Ehrenbrink et al. 2003), and 801 (Reisberg et al. 2008), and evolved gabbros at ODP 735 (Hart et al. 1999; Blusztajn et al. 2000), and site 894 (Lecuyer and Reynard 1996) failed to reproduce the true average for the complementary crustal reservoir to MORB lavas and therefore needs to be complemented with more detailed geochemical and petrologic studies of primitive gabbroic material from the lower crust. In order to more accurately assess the global HSE chemistry of the whole oceanic crust Peucker-Ehrenbrink et al. (2012) obtained data for an oceanic crust section from the Oman ophiolite that includes the crust–mantle transition. The mean weighted

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Figure 27. Histogram showing measured 187Os/188Os for (a) N-MORB glass data (Gannoun et al. 2004, 2007; Yang et al. 2013; Burton et al. 2015) (b) single grain sulfide data for MORB (unpublished data; Roy Barman et al. 1998; Gannoun et al. 2004, 2007; Burton et al. 2015) (c) abyssal peridotite whole-rock data (Martin 1991; Snow and Reisberg 1995;Brandon et al. 2000; Standish et al. 2002; Alard et al. 2005; Harvey et al. 2006) (d) single grain sulfide data for abyssal peridotites (Alard et al. 2005; Harvey et al. 2006; Warren and Shirey 2012). The estimate for the primitive upper mantle (PUM; Meisel et al. 1996) is also shown. The average 187Os/188Os composition of abyssal peridotites is 0.127 ± 0.015 (n = 129) while individual sulfides yield an average 187Os/188Os composition of 0.125 ± 0.021 (n = 63). N-MORB analyzed thus far show no evidence for a subchondritic source which may reflect local melting of abysFigure 27 sal peridotites (Brandon et al. 2000), resistance of depleted peridotites to remelting (Hirth and Kohlstedt 1996; Manga 1996) or that Os from undepleted (fertile) mantle dominates the MORB budget. However, the high-Os (> 20 ppb) sulfides yield an average composition of 0.129 ± 0.005 (n = 31) close to the PUM estimate with values as low as 0.1236. Therefore, these high-Os sulfides show no evidence for significant Re enrichment in the MORB source, as might be expected from the presence of recycled oceanic crust. Rather they indicate that the upper mantle source of these samples has experienced a long-term depletion of Re, similar to that observed in abyssal peridotites, and consistent with the incompatible nature of this element during mantle melting.

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Figure 28. CI-chondrite normalized HSE abundance plot for oceanic crustal rocks. Open squares: average Deep Sea Drilling Project (DSDP) Sites 417/418; diamonds: DSDP Hole 504B (Peucker-Ehrenbrink et al. 2003); solid squares: Ocean Drilling Program Hole 735B (Blusztajn et al. 2000); open circles: Oman crustal section (Peucker-Ehrenbrink et al. 2012); solid circles: composite ocean crust (Peucker-Ehrenbrink et al., 2012). The pattern of average MORB (This chapter) and abyssal peridotites (Reisberg and Lorand 1995; Pearson et al. 2004; Harvey et al. 2006; Luguet et al. 2007) are added for comparison. Normalization data from Lodders et al. (2009).

composition of the 4680 m Oman section yielded Re 427 pg g−1, Os 55 pg g−1, Ir 182 pg g−1, Pd 2846 pg g−1, Pt 4151 pg g−1 and initial 187Os/188Os of 0.142, indicating higher PGE concentrations and lower Re concentrations than all data previously reported on partial sections of ocean crust that lack cumulate lower crust. Assuming that these data are truly representative of the lower oceanic crust, they suggest that these rocks are the main HSE reservoir in the oceanic crust as a whole and that the average Re in these gabbros is much lower than in MORB lavas (Re ~ 1070 pg g−1; Hauri and Hart 1997; Gannoun et al. 2007; Gannoun et al. 2016, this volume). The Oman gabbros are characterised by a distinct subchondritic average Os/Ir ratio of ~ 0.3 which is significantly different from the chondritic ratio or the primitive upper mantle Figure 28 value of ~ 1.1 (Becker et al. 2006; Lodders et al. 2009). This difference is surprising because Ir is generally viewed as a geochemical analogue of Os during magmatic processes (Becker et al. 2006; Puchtel and Humayun 2000). The Os/Ir fractionation observed in the Oman gabbros, while within the range observed in MORB (0.2–1.4, average 0.6), is the opposite of that observed in the upper crustal part from DSDP 504B (average Os/Ir of ~ 2.4; Peucker-Ehrenbrink et al. 2003). However the Os/Ir of abyssal peridotites in general and in the harzburgitic mantle section of Oman in particular, remains chondritic (Hanghøj et al. 2010). If such harzburgites are representative of the mantle source then the subchondritic Os/Ir ratio in Oman gabbros cannot reflect a source signature. Hanghøj et al. (2010) report both superchondritic and subchondritic Os/Ir ratios in Oman dunites (0.5–8.3). As Os and Ir alloys included in chromites have been observed in Oman dunites (Ahmed and Arai 2002; Ahmed et al. 2006), Peucker-Ehrenbrink et al. (2012) suggested that such a phase may be responsible for the fractionation of Os from Ir during melting, melt extraction or crystal fractionation.

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Estimating the HSE inventory of the whole ocean crust remains challenging because of the discontinuous nature of field sampling and the question of how representative the samples that have been analyzed thus far actually are. Peucker-Ehrenbrink et al. (2012) used data from Site 504B for the upper oceanic crust (Peucker-Ehrenbrink et al. 2003) combined that for the Oman ophiolite for the lower oceanic crust (Peucker-Ehrenbrink et al. 2012). The weighted chemical and isotope characteristics of this “composite” oceanic crust (Fig. 28), corrected for Re decay since emplacement, are 736 pg g−1 Re, 45 pg g−1 Os, 133 pg g−1 Ir, 2122 pg g−1 Pd, 2072 pg g−1 Pt, 187Re/188Os: 80 and 187Os/188Os: 0.144. Such crust is more enriched in Re and less depleted in PGE than observed in average gabbros from ODP Hole 735D. Therefore, unless fundamentally altered during subduction, subducted oceanic crust will evolve to form a PGE-depleted, Re-rich mantle component that over time will evolve to radiogenic 187Os/188Os compositions. However, the projected ingrowth of radiogenic 187Os/188Os may be inhibited by the loss of Re from the basaltic upper part of the crust during eclogite-facies metamorphism (Becker et al. 2000; Dale et al. 2007), but this is not true for the gabbroic lower part of the crust (Dale et al. 2007).

Assimilation of gabbroic lower crust Recent work has shown that the crystallization of gabbros, troctolites, and other plutonic rocks of the lower oceanic crust may be protracted, and that these rocks sometimes possess ages that are several million years older than predicted from the magnetic ages of the overlying basaltic crust (e.g., Schwartz et al. 2005; Grimes et al. 2008). This extended timescale for the growth of the lower oceanic crust has been attributed to the crystallization of gabbros in the mantle followed by uplift to lower crustal depths (Schwartz et al. 2005; Grimes et al. 2008). Such uplift may relate to unroofing by low-angle detachment faults, typical of asymmetrical spreading ridge segments (e.g., Lissenburg et al. 2009). Over a timescale of several million years gabbros and troctolites, and their constituent phases, in the lower oceanic crust will rapidly evolve to radiogenic Os isotope compositions. This raises the possibility that younger melts passing through older lower crust may acquire a radiogenic Os isotope composition, either by remelting and assimilation of older material or through the physical entrainment of older crystals. Primitive xenocrysts are commonly found in MORB (e.g., Dungan and Rhodes 1978; Coogan 2014) with evidence for mixing shortly before eruption (e.g., Moore et al. 2014). Indeed, as discussed previously, the old ages of phenocryst phases in basalts that are thought to have been erupted just 5–10 kyr ago (Figs. 18 and 19), may indicate that these are xenocrysts physically entrained from previously solidified “olivine– plagioclase” bearing plutonic rocks through which the present host basalts have ascended. In this case, it is possible that some MORB glass may acquire a radiogenic Os isotope composition without interaction with seawater altered oceanic crust, or the presence of a radiogenic mantle source. For MORB glass such a signature might be distinguished by the absence of any covariation with Cl abundance or B isotopes.

HSE ABUNDANCES AND Re–Os ISOTOPE SYSTEMATICS OF INTRAPLATE VOLCANISM The HSE and Re–Os systematics of intraplate volcanism were reviewed recently by Day (2013). The purpose of this section is to briefly summarize the likely origins of intraplate volcanism, based specifically upon HSE abundance and Re–Os isotope constraints, and to provide an update of developments in the field since 2013. In particular, and mostly as a function of the difficulties associated with producing precise 186Os/188Os data (e.g., Chatterjee and Lassiter 2015), there have been limited advances in the application of the Pt–Os isotope system to intraplate volcanism since Day (2013); the interested reader is referred to this earlier review article for an up-to-date appraisal of Pt–Os isotope systematics.

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The origin of intraplate volcanism has been variously attributed to (i) mantle plumes (Wilson 1963; Morgan 1971), (ii) plumes which are not particularly “hot” (e.g., Falloon et al. 2007; Putirka et al. 2007), (iii) stress-driven processes (Anguita and Hernan 1975) or (iv) chemical heterogeneities preserved in the upper mantle (e.g., Courtillot et al. 2003; Arndt 2012). The occurrence of intraplate volcanism does not appear to be related to proximity to plate boundaries (cf. Hawai’i; Wilson et al. 1963 versus the Canary Islands; Morgan 1971) and does not occur systematically on either the continents or within oceanic basins, even spanning continental–oceanic margins (i.e., the Cameroon Line; Rehkämper et al. 1997; Gannoun et al. 2015a). Intraplate volcanism can be associated with convergent (e.g., Samoa; Wright and White 1987) and divergent (e.g., Iceland; Morgan 1971) tectonic settings. In general, intraplate volcanism is controlled by anomalous thermo-chemical and/or tectonic conditions capable of producing large volumes of extrusive products. Many investigations into the HSE of intraplate volcanic rocks have predominantly featured primitive, high-MgO rocks, e.g., komatiites and picrites (e.g., Ireland et al. 2009; Connelly et al. 2011, respectively), because of the compatibility of the HSE during fractional crystallization, and the sensitivity of 187Os/188Os to crustal assimilation processes in more evolved magmas (e.g., Chu et al. 2013). However, evolved potassic and sodic mafic–alkaline volcanic rocks and phonolites, trachytes and rhyolites, which may have experienced extensive fractional crystallization, are also observed and have recently been investigated for their HSE abundances and Re–Os isotope systematics (e.g., Chu et al. 2013; Li et al. 2014; Wang et al. 2014). For this reason, we adopt the same definition used by Day (2013) for intraplate ‘hotspot’ volcanism, i.e., “Volcanic rocks that are unassociated with conventional plate tectonic boundary magmatic processes and that may require anomalous thermo-chemical and/or tectonic conditions to induce small- to large-scale melting.”

Mantle melting processes The composition of the mantle source may be expressed by a variety of end-member compositions based upon its history of prior melt depletion i.e., depleted versus fertile peridotite (e.g., Niu 2004; Godard et al. 2008) and overall lithology, i.e., peridotite versus pyroxenite (Hirschmann and Stolper 1996; Yaxley 2000; Kogiso et al. 2004; Lambart et al. 2012, 2013). In addition, fertile heterogeneities in the mantle nucleate magmatic channels that focus melts up to the surface and hinder their re-equilibration with ambient peridotite (Katz and Weatherley 2012). Therefore, the chemical signature of hybrid melts of peridotite and pyroxenite can be retained in the composition of mantle-derived basalts. Day (2013) discussed the significance of the ‘shape’ that a melting regime can have, discussing two end-member geometries; (i) batch melting of a columnar (cylindrical) region (e.g., Rehkämper et al. 1999), and (ii) regions of adiabatic melting in triangular or corner-flow melting regime (e.g., Plank and Langmuir 1992). Each of these melting regimes aggregate melt pooled from over the melting volume, accounting for the overall composition of the magma generated. Briefly, model (i) is most consistent with an upwelling ‘mantle plume-like’ melting regime. It assumes uniform melting throughout the source region and that the extraction of sulfide-hosted HSE is completely exhausted at 20–25% partial melting. This cylindrical melting model reproduces the HSE abundances of low-degree alkali basalts (e.g., Canary Island lavas; e.g., Day et al. 2009) and high-degree partial melts (e.g., komatiites; e.g., Rehkämper et al. 1999), but the HSE signature of some tholeiitic magmas generated by low degrees of partial melting are not predicted using this cylindrical melt volume (e.g., Momme et al. 2003, 2006). The triangular melting regime (model ii) assumes near-fractional melting in 1% increments with decreasing pressure, i.e., through adiabatic ascent (e.g., Rehkämper et al. 1999; Momme et al. 2003). In this melting regime, S-saturated low-degree partial melts with low HSE-concentrations mix with shallower, higher-degree (and potentially S-undersaturated) partial melt. Refinements to the two general classes of models described above have allowed distinct melt regimes in some continental flood basalt (CFB) provinces to be determined

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(Momme et al. 2006), whereas in the Icelandic rift zones depleted versus enriched mantle components have also been identified (Momme et al. 2003). Moreover, the use of these models has permitted the detection of a pyroxenitic component in primitive lavas from the Canary Islands (Day et al. 2009), and a similar component has been implicated in the generation some Hawaiian lavas (Lassiter et al. 2000; Sobolev et al. 2007). Source compositional estimates become increasingly complicated when the necessity arises to account for the contributions from mixtures of source lithologies (e.g., peridotite and recycled sediment or basalt) and the complex interplay of the HSE that each of these source reservoirs may contribute to a pooled melt (e.g., Hirschmann and Stolper 1996).

Osmium isotopes as tracers of hotspot sources Ocean island basalts. Many intraplate basalts retain HSE signatures of their mantle source region and osmium isotopes, when compared to lithophile element-based radiogenic isotopes, can offer a unique perspective on the petrogenesis of intraplate lavas. The large Re/Os fractionations generated during crust-mantle partitioning make it possible to model 187Os/188Os variations in OIB in the context of variably aged recycled crust and lithosphere (e.g., Hauri and Hart 1993; Marcantonio et al. 1995; Widom et al. 1999; Day et al. 2009; Day 2013). For example, ancient oceanic mantle lithosphere or SCLM has been implicated in the genesis of lavas from the Azores, Iceland and Jan Mayen (Skovgaard et al. 2001; Schaefer et al. 2002; Debaille et al. 2009), where measured unradiogenic 187Os/188Os values cannot be explained by melting exclusively of modern oceanic lithospheric material and thus require a mantle source or sources that have evolved in a low Re/Os environment (cf. unradiogenic abyssal peridotites reported by Snow and Reisberg 1995; Alard et al. 2005; Harvey et al. 2006; Liu et al. 2008; Warren and Shirey 2012; Lassiter et al. 2014). Intraplate basalts and specifically OIB, are generated from mantle sources with distinct long-term time-integrated parent-daughter fractionations of Sr–Nd–Pb–Hf isotopes (e.g., Zindler and Hart 1986; Hofmann 2003; White 2010), and also preserve a large range of 187Os/188Os compositions (e.g., Pegram and Allègre 1992; Hauri and Hart 1993; Reisberg et al. 1993; Marcantonio et al. 1995; Roy-Barman and Allègre 1995; Widom and Shirey 1996; Lassiter and Hauri 1998; Brandon et al. 1999, 2007; Widom et al. 1999; Schiano et al. 2001; Eisele et al. 2002; Schaefer et al. 2002; Lassiter et al. 2003; Workman et al. 2004; Escrig et al. 2005b; Class et al. 2009; Day et al. 2009, 2010b, 2015; Debaille et al. 2009; Ireland et al. 2009; Jackson and Shirey 2011). These signatures are only retained in instances where the melt produced at depth, albeit with ancient timeintegrated compositions, and reflecting the recycling of material back into the convecting mantle (e.g., Zindler and Hart 1986), are not significantly contaminated or overprinted though interaction with the lithosphere through which these basalts necessarily transit en route to the surface. For example, in a recent study of the Louisville Seamount Chain, Tejada et al. (2015) demonstrated that OIB erupted along this chain of volcanoes reach the surface with negligible chemical interaction with the lithospheric mantle that underlies the South Pacific. Moreover, unlike the Hawaiian–Emperor Seamount chain, whose compositions are readily explained by heterogeneous mantle sources (see following section), osmium isotope signatures of these basalts have a very narrow range, consistent with their derivation from a primitive mantle source (cf. Meisel et al. 2001; Becker et al. 2006). Age corrected 187Os/188Os of the Louisville Seamount basalts range from 0.1245–0.1314, similar to other Pacific OIB, such as Rarotonga (0.1249–0.1285, Hauri and Hart (1993); 0.124–0.139, Hanyu et al. (2011) and some Samoan basalts (0.1230–0.1313, Hauri and Hart (1993); Jackson and Shirey (2011)). The age corrected 187Os/188Os for two aggregates of olivine phenocrysts separated from Louisville Seamount basalts (0.1272 and 0.1271–0.1275) agree with whole rocks from the same seamount (0.1253–0.1274; Tejada et al. 2015), supporting the hypothesis that early-crystallizing olivine can preserve the pristine magmatic Os isotopic compositions of their source (cf. Jackson and Shirey 2011; Hanyu et al. 2011) (Fig. 29).

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Os (pg g-1) Figure 29. Plot of Os concentration versus Os isotope ratios for Louiville Seamount Chain basalts (agecorrected) and olivines (present-day). Pacific ocean island basalts (OIB), mid-ocean ridge basalts (MORB), and Ontong Java Plateau basalts (OJP) basalts are shown for comparison. Osmium abundances and isotopic signatures are limited compared to other Pacific OIB. Data sources: Schiano et al. (1997, 2001); Brandon et al. (1999); Eisele et al. (2002); Jackson and Shirey (2011); Hanyu et al. (2011); Tejada et al. (2013). Modified after Tejada et al. (2015).

Figure 29 Studies of HSE abundance complement and extend the knowledge of intraplate magma petrogenesis gleaned from Os isotope systematics. Only lavas with high MgO contents and > 0.05 ng g−1 Os should be considered as potentially being representative of the true HSE characteristics of intraplate magma and its mantle source. Such restrictions on the analysis of intraplate magmas mean that there is still a dearth of high quality HSE data on OIB. Much of what has been elucidated from HSE abundances in OIB comes from studies of Hawaiian lavas (Bennett et al. 2000; Crocket 2002; Jamais et al. 2008; Ireland et al. 2009; Pitcher et al. 2009). These studies support the hypothesis that, in general, high-MgO lavas preserve early-formed Os-rich (+ HSE) phases that become incorporated in early forming phenocrysts such as olivine (e.g., Brandon et al. 1999; Ireland et al. 2009). Removing the effects of mineral fractionation on HSE abundances allowed Day (2013) to directly compare the absolute and relative HSE abundances and calculated Re/Os of parent melts in addition to 187Os/188Os, of Hawaiian, Canary Island and Samoan lavas. Combined with the HIMU type 206Pb/204Pb compositions of Canary Island lavas, this led to the conclusion that, in contrast to Hawaiian and Samoan OIB, and komatiites, whose compositions suggest a relatively high proportion of peridotite in their parental melts, lavas from the Canary Islands, and specifically El Hierro and La Palma, contain recycled oceanic crust in their mantle source. Osmium isotope studies of HIMU-type OIB support and enhance Sr–Nd–Pb isotope and trace element arguments for a recycled oceanic lithosphere component in their mantle source (Hauri and Hart 1993; Marcantonio et al. 1995; Widom et al. 1999; Eisele et al. 2002; Day et al. 2010b). The observed range of 187Os/188Os and 206Pb/204Pb of HIMU basalts (e.g., Becker et al. 2000; Dale et al. 2009a; van Acken et al. 2010) could be produced by direct melting (~ 50% to 90%) of recycled oceanic crust, but would result in melts that contain too much silica and too little magnesium (e.g., Yaxley and Green 1998). Although field evidence suggests that pyroxenites account for ≤ 10% of mantle lithologies (e.g., Pearson et al. 1991; Reisberg et al. 1991), they melt disproportionately to

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peridotite under any P–T conditions (only 1–2% pyroxenite may generate up to 50% of the melt at low degrees of partial melting), thus producing silica-undersaturated, iron-rich melts with high MgO (e.g., Hirschmann et al. 2003). This means that direct melting of recycled oceanic crust and lithosphere is not necessary to produce HIMU OIB. Spinel websterites have been suggested to be geochemically analogous to pyroxenites, at least in terms of their HSE systematics (Marchesi et al. 2014). The Re–Os fractionation generated as a result of peridotite versus pyroxenite (and/or spinel websterites) has been suggested as a likely contributor to the observed 186Os/187Os-rich compositions of some plume basalts (Luguet et al. 2008) previously attributed to interaction between the mantle and outer core (e.g., Walker et al. 1997; Brandon et al. 1998, 2003; Puchtel et al. 2005). Subsequent studies (e.g., Baker and Jensen 2004; Scherstén et al. 2004; Luguet et al. 2008) and, more recently, Marchesi et al. (2014) suggest that such enrichments could be attributed to processes requiring no input from the outer core. However, these models may require unreasonably high contributions from pyroxenitic / spinel websteritic lithologies in the mantle (as high as 90%; van Acken et al. 2010; Marchesi et al. 2014), as a result of the comparatively low Os concentrations in pyroxene-rich lithologies. The enriched mantle (EM) signatures of other OIB has been attributed to the addition of subducted sediment or metasomatized lithosphere into their mantle sources (e.g., Workman et al. 2004). EM-type OIB span a range of compositions in Sr–Nd–Pb isotope space, varying from EMI (e.g., Pitcairn; Woodhead and McCulloch 1998; and the Comores; Class et al. 2009), which exhibit a wide range of Os- and Pb-isotope compositions, but more restricted Sr isotope compositions, to EMII OIB (e.g., Samoa; Wright and White 1987; Workman et al. 2004; Jackson and Shirey 2011). These compositions are consistent with sediment, recycled oceanic crust and peridotite producing EMI-flavoured compositions with more radiogenic 187Os/188Os (Roy-Barman and Allègre 1995; Class et al. 2009), while subducted sediment mixed with ambient peridotite produces enriched EMII compositions with lower 187Os/188Os. Therefore, lithological variations in the mantle source play a key role in the composition of OIB, and HSE abundances combined with 187Re-187Os systematics are critical in the identification of the various components mixed with variably depleted asthenospheric mantle. Continental intraplate volcanism. The heterogeneous mantle sources described above are not restricted to OIB, or oceanic settings in general. These modifiers of magma composition also influence intraplate volcanism associated with continental regions. The main differences between oceanic and continentally erupted intraplate magmas is the greater potential for the latter to be influenced by interaction with the thicker and older overlying sub-continental lithospheric mantle (SCLM) and continental crust, in addition to the potential compositional heterogeneities within the asthenospheric mantle. Recently, Sun et al. (2014) reported Re-Os systematics of ultrapotassic (> 7 wt% K2O) basalts from the Xiaogulihe area of western Heilongjiang Province, NE China. The relatively unradiogenic Os isotope ratios (187Os/188Os = 0.1187–0.1427) contrasted with the similarly potassic basalts from NE China reported by Chu et al. (2013) (187Os/188Os = 0.13–0.17) and were attributed by Sun et al. (2014) to a dominantly peridotitic source, but one that required an unusually high K2O content. In this particular setting, phlogopite-bearing garnet peridotite hosted within the lower part of the SCLM was implicated; its derivation being potassium-rich silicate melts produced by the subduction of ancient continent-derived sediments (> 1.5 Ga). The observation that lherzolite xenoliths from Keluo and Wudalianchi contain phlogopite (Zhang et al. 2000, 2011) supports the hypothesis that SCLM, metasomatized by potassium-rich melts, is present beneath the WEK volcanic field and contributes to the basalts from Xiaogulihe.

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Crustal and lithospheric mantle assimilation/contamination Oceanic intraplate volcanism is often assumed to be immune to lithospheric contamination. Compared to continental intraplate eruptions, OIB do not interact with thermo-chemically complex SCLM. The low Os contents in OIB (typically < 1 ng g−1) makes the Re–Os isotope system a particularly sensitive indicator of lithospheric contamination, and the relatively unradiogenic 187Os/188Os compositions (< 0.18) of OIB relative to local oceanic crustal reservoirs (typically 187Os/188Os > 0.4; Reisberg et al. 1993; Marcantonio et al. 1995; PeuckerEhrenbrink et al. 1995; Widom et al. 1999) make the tracing of assimilation of crustal or lithospheric mantle materials in OIB a straightforward process (e.g., Reisberg et al. 1993; Marcantonio et al. 1995; Lassiter and Hauri 1998; Skovgaard et al. 2001; Gaffney et al. 2005). In particular, at the lowest levels of Os content, OIB are even more vulnerable to crustal contamination (Reisberg et al. 1993), while OIB with Os contents greater than 30 to 50 pg g−1 are typically assumed to be less susceptible to assimilation of lithospheric components (e.g., Reisberg et al. 1993; Eisele et al. 2002; Class et al. 2009). Crustal contamination thus rapidly drives Os isotope ratios to more radiogenic values resulting from the assimilation of oceanic crust with high Re/Os and 187Os/188Os. A consequence of the low HSE abundances of crustal material is that the addition of crust to a primitive melt should result in the dilution of HSE abundances in the resultant magma. Ireland et al. (2009) presented such a model, illustrating the effect of crustal contamination on Hawaiian picrites. Briefly, three end-member scenarios are considered; (i) continental crust addition to komatiite; (ii) oceanic crust addition to tholeiite and, (iii) abyssal peridotite addition to alkali basalt. These models demonstrate that crustal contamination dilutes OIB HSE abundances at ≤20 % crustal or lithospheric assimilation. However, both 187Os/188Os and Re/Os can change dramatically in the evolving liquid, which has implications for the time integrated Os isotope ratio of such contaminated magmas and the effectiveness of using 187 Os/188Os as a tracer for the mantle source of the magma. The effects of assimilation on HSE abundances (absolute or relative) in general, are less well-defined and where this issue has been addressed in the literature the consensus appears to be that fractional crystallization exerts a stronger influence on HSE distributions than contamination factors (e.g., Chazey and Neal 2005; Ireland et al. 2009). However, crustal contamination of continental flood basalts (CFB) can lead to a significant augmentation in the S content of a magma, sometimes resulting in S-saturation and significant HSE fractionation (e.g., Keays and Lightfoot 2007; Lorand and Alard 2010). This may also elevate concentrations of Re and the PPGE relative to Os, Ir, and Ru. Assimilation of mantle lithosphere also has pronounced effects on Re/Os, but requires large additions to generate significant effects on magma HSE abundances. Conversely, Widom et al. (1999) demonstrated that unusually unradiogenic 187Os/188Os in some Canary Island lavas was most likely the result of the assimilation of peridotite xenoliths with sub-chondritic 187Os/188Os and >1 ng g−1 Os, prior to the eruption of the basalt at the surface. More recently, a similar process was described by Gannoun et al (2015a) to account for particularly unradiogenic Os concentrations in basalts from the Cameroon Line (Fig. 30). These simple crustal contamination models can be greatly complicated by the inclusion of fractional crystallization processes, which are often intimately associated with crustal contamination. The combination of these processes will almost inevitably result in the generation of elevated Re, Pt, and Pd abundances compared to Os, Ir, and Ru in melts and crustal rocks, compared with their corresponding mantle residues. However, direct measurement of 187Os/188Os in early formed mineral phases handpicked from intraplate magmas, such as olivine (Debaille et al. 2009; Jackson and Shirey 2011), generally yield more restricted ranges in 187Os/188Os than their associated whole-rocks, and may provide a means of seeing past bulk-rock contamination of OIB. As a result of these potential complications, a common sense approach, based upon a rigorous assessment of local potential contaminants and melt products was advocated by

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Figure 30. 187Os/188Os versus (a) 206Pb/204Pb and (b) 143Nd/144Nd diagrams for Cameroon Volcanic Line (CVL) basalts. Pb and Nd isotope data are from Barfod (1999) and Lee et al. (1996). HIMU, DMM, EM1, and BSE are shown for reference. The average 187Os/188Os ratio for sub-continental lithospheric mantle is from Shirey and Walker (1998). Ultramafic xenoliths beneath the continental part ofFigure the CVL30 are also shown. The increments in the curves are 2%. The grey shaded area indicates the possible compositions for crustally contaminated lavas. The most radiogenic samples from the continental sector can be explained by assimilation of 8 to 16% of continental crust. Assuming for the uncontaminated starting point [Os] = 10 pg g−1, 187Os/188Os = 0.156, 206Pb/204Pb = 20.24, [Pb] = 2 µg g−1; for upper continental crust (UCC) [Os] = 50 pg g−1, 187Os/188Os = 1.4, 206Pb/204Pb = 19.3, [Pb] = 8 µg g−1; and for lower continental crust (LCC) [Os] = 50 pg g−1, 187Os/188Os = 0.8 (Saal et al. 1998), 206Pb/204Pb = 17.5, [Pb] = 8 µg g−1. Curves (1) and (2) describe the possible mixing trajectories between HIMU and DMM. (1) Assimilation of mantle xenocrysts and (2) mixing of lavas derived from DMM and HIMU sources. Modelling parameters are as follows: for (1) DMM mantle [Os] = 3 ng g−1, 187Os/188Os = 0.125, 206Pb/204Pb = 18.5, [Pb] = 0.15 µg g−1 and for (2) DMM melt [Os] = 8 pg g−1, 187Os/188Os = 0.127, 206Pb/204Pb = 18.5, [Pb] = 0.45 µg g−1. For Os–Nd modelling the starting point was chosen to be the closest to HIMU endmember, 143Nd/144Nd = 0.513, [Nd] = 40 µg g−1, 187 Os/188Os = 0.15, [Os] = 10 pg g−1; for UCC 143Nd/144Nd = 0.512, [Nd] = 27 µg g−1 (Rudnick and Fountain, 1995), 187Os/188Os = 1.4, [Os] = 50 pg g−1; and for LCC 143Nd/144Nd = 0.512, [Nd] = 50 µg g−1 (Kwékam et al. 2013), 187Os/188Os=0.8, [Os]=30 pg g−1 (Saal et al. 1998). [Reproduced with permission of Elsevier BV from Gannoun A, Burton KW, Barfod DN, Schiano P, Vlastélic I, Halliday AN (2015a) Resolving mantle and magmatic processes in basalts from the Cameroon volcanic line using the Re–Os isotope system. Lithos Vol. 224–225, p. 1–12.]

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Day (2013) when applying thresholds for “contaminated” versus “uncontaminated” OIB. Both crustal and SCLM contamination of primitive melts have been reported in the literature (e.g., Ellam et al. 1992; Horan et al. 1995; Molzahn et al. 1996; Chesley and Ruiz 1998; Keays and Lightfoot 2007; Li et al. 2010; Chu et al. 2013). Successful modelling of SCLM or crustal assimilation is dependent upon the accurate determination of likely end-member compositions, ranging from the parental primitive melt to its possible assimilants. Day (2013) successfully demonstrated the effect of contamination of primitive parent melts using North Atlantic Igneous Province (NAIP) picrites (Schaefer et al. 2000; Kent et al. 2004; Dale et al. 2009b) and intrusive rocks from the Rum Intrusion (O’Driscoll et al. 2009).

The origin of Continental Flood Basalts (CFB) and Large Igneous Provinces (LIP) Volcanic rocks from some CFB have been interpreted to have survived the transit from their asthenospheric source to eruption at the surface without any significant interaction with the SCLM or the crust (e.g., Schaefer et al. 2000; Zhang et al. 2008; Dale et al. 2009b; Rogers et al. 2010; Day et al. 2013). Many of these lavas are picritic in composition, have high-MgO (> 13.5 wt%), high Os concentrations, and 187Os/188Os which are, in general, unradiogenic; consistent with their derivation from primitive mantle or a depleted mantle source (e.g., Schaefer et al. 2000; Dale et al. 2009b; Rogers et al. 2010). This chemical and isotopic signature has, in turn, been used to suggest that such CFB may be modern-day equivalents of uncontaminated Archaean komatiites, albeit from a cooler mantle, (cf. Brügmann et al. 1987; Wilson et al. 2003; Puchtel et al. 2009; Connolly et al. 2011). In contrast, several studies have highlighted the importance of an interaction between asthenosphere-derived melts, SCLM and the crust to produce the observed spectrum of CFB compositions (e.g., Ellam et al. 1992; Horan et al. 1995; Molzahn et al. 1996; Chesley and Ruiz 1998; Xu et al. 2007; Li et al. 2010; Heinonen et al. 2014) and the HSE fingerprint of some komatiites (Foster et al. 1996). Osmium isotope systematics, combined with other radiogenic isotope tracers in CFB demonstrate that the interplay between a primary magma and its potential lithospheric contaminants can be complex, as illustrated in a number of localities (e.g., Siberia–Horan et al. 1995; Ethiopia–Rogers et al. 2010; Emeishan, China,–Zhang et al. 2008). Correlations between 187Os/188Os and 87Sr/86Sr (Molzahn et al. 1996) 206Pb/204Pb (Xu et al. 2007), and possibly even 3He/4He (Dale et al. 2009b) illustrate the effects of lithospheric contamination on primary, asthenosphere-derived melts. However, observed variations in Os isotopes are not wholly consistent with SCLM or crustal contamination alone, suggesting that, like many OIB, some inherent heterogeneity within the asthenospheric source is present. For example, the 260 Ma Emeishan province (e.g., Li et al. 2010) requires a more depleted mantle source than 190 Ma Karoo CFB (Ellam et al. 1992). Some CFB provinces may therefore tap mantle sources that contain recycled material, similar to the source of some HIMU and EM flavoured OIB (e.g., Shirey 1997; Dale et al. 2009b), while others are derived from an essentially primitive mantle (see review in Day 2013). Heterogeneity in the composition and distribution of sulfide types within a magma source region in the mantle (e.g., interstitial versus enclosed sulfides; Alard et al. 2002; see also Harvey et al. 2016, this volume) can have a profound influence on the composition of a basaltic melt (e.g., Harvey et al. 2010, 2011). The combination of source heterogeneity and degree of partial melting can therefore account for the observed differences in initial Os isotopic and HSE abundance variations in CFB provinces, that range from depleted DMM-like mantle compositions (e.g., Rogers et al. 2010) through undepleted basalts (e.g., Schaefer et al. 2000), to more radiogenic compositions, which provide strong evidence for recycled components in some CFB provinces (Shirey 1997) Coupled with the effects of adding subducted oceanic lithosphere back into the convecting mantle, i.e., the source of CFB and LIP, and the combinations of pelagic / terrigenous sediments,

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variably altered oceanic crust and serpentinized peridotite (Allègre and Turcotte 1986), unravelling the sources of voluminous basaltic magmatism has somrtimes been demonstrated to be problematic, often requiring both HSE and Re–Os isotope evidence used in concert with more traditional lithophile element-based isotope systems. For example, Heinonen et al. (2014) invoked a mixture of depleted Os-rich peridotite with ∼10–30% of seawater-altered and subduction-modified MORB (with a recycling age of less than 1.0 Ga) as the likely source of the distinctive isotopic fingerprint found in CFB from the Antarctic Karoo province. A specific mixed peridotite-pyroxenite-like source was required to explain the unusual combination of elevated initial 87Sr/86Sr and Pb isotopic ratios, and low initial 187Os/188Os observed in the dykes sampled from around Ahlmannryggen, western Dronning Maud Land. In other words, simple, two component mixing is often not consistent with the observed chemical and isotopic composition of CFB. In the example described by Heinonen et al. (2014), not only was a combination of mixed lithologies in the source, in addition to the inherent differences in their HSE and 187Os/188Os fingerprints required to account for the composition of the Ahlmannryggen dykes, but also a contribution from a seawater-altered subducted component was required. A similar investigation into the nature of the Eastern North America (ENA) Central Atlantic Magmatic Province (CAMP) by Merle et al. (2014) also revealed the complex combination of chemical and isotopic fingerprints that can be preserved in large-volume basaltic eruptions. Although CAMP magmatism in general may have been produced as a result of either heat incubation under thick continental lithosphere (McHone 2000; De Minet al. 2003; Puffer 2003; McHone et al. 2005; Verati et al. 2005; Coltice et al. 2007), or by a plume head under the continental lithosphere (May 1971; Morgan 1983; White and McKenzie 1989; Hill 1991; Wilson 1997; Courtillot et al. 1999; Ernst and Buchan 2002; Cebria et al. 2003), Merle et al. (2014) proposed several increasingly complex scenarios to account for the chemical and isotopic signatures preserved in the ENA CAMP basalts, including (i) direct derivation from a mantle plume (Wilson 1997) or oceanic plateau basalt-type melts (e.g., Kerr and Mahoney 2007); (ii) magmas derived from a mantle plume but contaminated by continental crust en route to the surface (Arndt al. 1993); (iii) mixing between asthenospheric and ultra-alkaline mafic (lamproite, kimberlite, and kamfugite) melts (Arndt and Christensen 1992; Gibson et al. 2006; Heinonen et al. 2010), possibly followed by crustal contamination; (iv) ternary mixing between OIB, MORB and SCLM-related melts, possibly followed by crustal contamination; (v) direct melting of a shallow source enriched in incompatible elements such as metasomatized SCLM or the mantle wedge above subduction zones (Puffer 2001; De Min et al. 2003; Deckart et al. 2005; Dorais and Tubrett 2008). Unfeasibly large degrees of crustal contamination would be required to produce the observed 143Nd/144Nd, 206Pb/204Pb and 208Pb/204Pb isotopic compositions of the ENA CAMP basalts, and crustal contamination, assimilation (of continental crust) with fractional crystallization (DePaolo 1981) and assimilation through turbulent ascent were discounted on the strength of the Re-Os and 187Os/188Os systematics i.e., initial 187Os/188Os higher than 0·15 at Os concentrations lower than 50 ng g−1 (e.g., Widom 1997). Merle et al. (2014) determined that mixing involving either OIB or MORB-like parental melts, followed by crustal contamination, partially reproduces the compositions of the ENA CAMP basalts, but the trends observed in the Nd–Pb and Os–Nd isotopic diagrams require the addition of up to 35% continental crust, yet the assimilation of more than 20% of continental crust is thermodynamically unrealistic (Spera and Bohrson 2001). Consequently, the hypothesis of a magma originating from mixing between OIB and SCLM-related melts and further contaminated by the continental crust was deemed unlikely. Therefore, the continental crust-like characteristics of the ENA CAMP were inferred to be present in the mantle source itself. Recent studies have suggested that such contrasting chemical characteristics may be derived from a metasomatized SCLM-type source (cf. Chu et al. 2013; Sun et al. 2014; Wang et al. 2014), where phlogopite in the SCLM was thought to be derived from the melting of subducted terrigenous sediments. To

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account for the measured Os isotope compositions of the ENA CAMP basalts, the Os isotopic composition of the source needed to be within the range of 187Os/188Os for off-cratonic SCLM (0·1180–0·1290; Carlson 2005), therefore the model favored by Merle et al. (2014) to explain the multi-isotope system fingerprint of the EMA CAMP basalts required a reservoir that experienced progressive incorporation of subducted sediments derived from the local continental crust into a depleted sub-arc mantle wedge above a subduction zone. Recent work has revealed that HSE abundances can be broadly modelled as a function of fractional crystallization in CFB. Day et al. (2013) studied the 1.27 Ga Coppermine CFB in northern Canada, which represents the extrusive manifestation of the Mackenzie large igneous province (LIP), which includes the Mackenzie dyke swarm and the Muskox layered intrusion. These authors reported new HSE abundance and Re-Os isotope data for picrites and basalts from the CFB, as well as a highly unusual andesite glass flow. The glass contained high HSE contents (e.g., 3.8 ng g−1 Os) and mantle-like initial 187Os/188Os (γ1270Ma Os = +2.2), but δ18O, eNdi, and trace element abundances consistent with extensive crustal contamination, implicating a potential origin for this sample (CM19) as a magma mingling product formed within the Muskox Intrusion during chromitite genesis (cf. Day et al. 2008) and direct evidence for the processing of some CFB within upper-crustal magma chambers. These authors also modelled absolute and relative HSE abundances in CFB from the Coppermine, Parana and West Greenland, revealing that HSE concentrations decrease with increasing fractionation for melts with < 8 ± 1 wt% MgO (Fig. 31). The models reveal that significant interelement fractionation between (Re + Pt + Pd)/(Os + Ir + Ru) are generated during magmatic differentiation in response to strongly contrasting partitioning of these two groups of elements

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into sulfides and/or HSE-rich alloys. Furthermore, fractional crystallization has a greater role on absolute and relative HSE abundances than crustal contamination under conditions of CFB petrogenesis due to the dilution effect of low total HSE continental crust. Day et al. (2013) found that picrites (> 13.5 wt% MgO) from CFB (n = 98; 1.97 ± 1.77 ppb) having higher Os abundances than OIB picrites (n = 75; 0.95 ± 0.86 ppb) and interpreted these differences to reflect either higher degrees of partial melting to form CFB, or incorporation of trace sulfide in CFB picrites from magmas that reached S-saturation in shallow-level magma chambers.

Continental intraplate alkaline volcanism Continental intraplate alkaline volcanic rocks (CIAV) comprise a wide spectrum of sodic and potassic compositions ranging from alkali basalts, picrites and basanites through to more evolved eruptive products that include nephelinites, carbonatites, melilitites, and kimberlites. The origin of some of these rock types are not unequivocal, with petrogenetic models ranging from pure incipient rift-related sources (e.g., Thompson et al. 2005), to ‘hotspot’ or ‘plume’ related origins (e.g., Haggerty 1999). Finding a likely source for these volcanic rocks is not made any less ambiguous when experimental and geochemical data are considered as many of these lavas are thought to derive from close to the boundary layer that separates the convecting and conducting mantle (e.g., Foley 1992; Day et al. 2005), i.e., both the asthenosphere and SCLM can be implicated in the genesis of these magmas. Re–Os isotope data are limited for these types of lavas, and instances where this is combined with HSE abundance data are comparatively rare. Examples from the literature when HSE and / or Re–Os isotope data are available are summarized in Day (2013) When elevated osmium contents in basalts clearly exclude the influence of crustal contamination, radiogenic 187Os/188Os (e.g., > 0.15) is often interpreted as being derived from olivine-poor mantle heterogeneities, such as clinopyroxenites (Carlson et al. 1996; Carlson and Nowell 2001; Janney et al. 2002), primarily as a result of their time-integrated ingrowths to high 187Os/188Os (Reisberg et al. 1991; Reisberg and Lorand 1995; Kumar et al. 1996). At the onset of S-saturated melting at depth, these fertile heterogeneities with radiogenic Os isotopic compositions melt preferentially (Hirschmann et al. 2003; Rosenthal et al. 2009). Combined with the Os isotope and HSE signature associated with pyroxenite-dominated melts, high NiO and low MnO concentrations in olivine phenocrysts are also diagnostic of olivine-poor mantle domains such as phlogopite-rich pyroxenites (Prelević et al. 2013). These phlogopitebearing pyroxenites can be derived from the reaction of peridotitic mantle wedge with melts derived from terrigenous sediments, possibly from the uppermost regions of the subducting slab (Prelević et al. 2015). As such, many CIAV appear to have non-peridotitic sources, with some sodic mafic-alkali magmas possessing radiogenic 187Os/188Os compositions, but moderately high Os contents (> 0.5 ng g−1 Os). Extreme Os isotopic compositions could reflect low degrees of partial melting and preferential sampling of more fusible mafic components, such as pyroxenite, in the asthenospheric mantle (cf. CFB above). Alternatively, melting of metasomatized lithosphere during rifting events (e.g., Carlson and Nowell 2001; Thompson et al. 2005) may also be responsible for the HSE abundances and Re–Os systematics of some CIAV, such as the Newer volcanic rocks, Australia (Vogel and Keays 1997). Similarly, carbonatites may also ultimately originate from mafic as opposed to ultramafic sources due to their close association with other ultrapotassic rocks (e.g., Gudfinnsson and Presnall 2005). For example, young (> 20 Ma) carbonatites from Fuerteventura, Canary Islands, possess low Os abundances (5–15 pg g−1) and highly radiogenic 187Os/188Os that extend to values in excess of 0.6 (Widom et al. 1999). Conversely, the high Os abundance and unradiogenic Os isotope signatures of some kimberlites and katungites are consistent with a petrogenesis involving the assimilation or derivation from the SCLM (Pearson et al. 1995; Carlson et al. 1996; Araujo et al. 2001; Carlson and Nowell 2001; Pearson et al. 2008). More recently, Chalapathi Rao et al (2013) provided strong evidence for contrasting mantle sources for kimberlites and lamproites in the

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Eastern Dharwar craton, southern India. Re–Os isotope of orangeites from the Bastar craton and Mesoproterozoic kimberlites and lamproites contrasted with an unradiogenic Re-depleted kimberlite sample with present-day 187Os/188Os (0.1109) and a Re–Os isotopic fingerprint characteristic of Proterozoic lithosphere, with the positive γOs (2.9–3.6; where γOs refers to the percentage deviation at the time of emplacement of the 187Os/188Os from Primitive Upper Mantle with a 187Os/188Os of 0.1296; Becker et al. 2006)) of two kimberlites from Raichur and Narayanpet (Eastern Dharwar craton) that retained both both plume and subduction-related source signatures (cf. Heinonen et al. 2014 for the petrogenesis of continental flood basalts from the Antarctic province of the Karoo). The enriched Re/Os mantle sources for the nearby Kodomali orangeite (γOs = +3) and the Krishna lamproites, with very radiogenic (γOs of + 56 to + 355), similar to those displayed by the lamproites of the Italian peninsula (Conticelli et al. 2007), suggest a subducted component for the latter ultra-potassic rocks, demonstrating the complex interplay of likely sources contributing to magma genesis around the Eastern Dharwar craton in both time and space (Chalapathi Rao et al. 2013). The low Os concentrations of primary low-degree potassic and sodic mafic–alkali volcanic rocks, combined with the high Os abundance of mantle and crustal xenoliths in some kimberlites, alnoites and melnoites make these volcanic rocks highly susceptible to contamination as they pass through and interact with the SCLM and overlying crust. Evolved magmas of this type may also be susceptible to the effects of S-saturation prior to eruption (Vogel and Keays 1997), i.e., they may have experienced the prior precipitation of sulfide and concomitant harvesting of HSE from the S-saturated magma. Despite these caveats, some continental intraplate magmas still retain unique information on the composition of their mantle source. In particular, early Cretaceous alkaline picrites and basalts from the North China craton have petrological and Os–Sr–Nd isotope compositions consistent with contributions from recycled and foundered eclogitic lower continental crust (Gao et al. 2008). More recently, Chu et al. (2013) examined a suite of highly potassic basalts from Wudalianchi-Erkeshan, NE China and, despite the incorporation of modest amounts of continental crust and the potential of sulfide contamination derived from the SCLM, traced the source of the basalts back to the asthenosphere. Their findings suggested a complex interaction between crust and SCLM with highly potassic melts generated at least partly from SCLM containing phlogopite, itself with an ancient terrigenous sediment signature (Sun et al. 2014). In contrast to a predominantly peridotitic phlogopite-bearing source for continental volcanism reported by Sun et al. (2014), Miocene ultrapotassic rocks within the Sailipu area of the western Lhasa terrane, southern Tibet, were variously attributed to the interaction of both spinel- and garnet-lherzolite derived melt with a phlogopite-bearing pyroxenite source (Wang et al. 2014). Although the latter study postulated that the observed chemistry of the ultramafic melts could be attributed to crustal contamination, unfeasibly large-scale assimilation of continental crust would be necessary to account for the nature of the Sailipu basalts. While the lithophile element-based isotope systems are relatively insensitive to crustal contamination, mixing calculations using HSE concentrations and 187Os/188Os of primitive arc compositions (Os = 0.2 ng g−1; 187Os/188Os = 0.125; Shirey and Walker 1998; Suzuki et al. 2011), continental crust (Os = 0.01 ng g−1; 187Os/188Os = 1.10; Shirey and Walker 1998) and depleted mantle material (Os = 0.405 ng g−1; 187Os/188Os = 0.10815; Shirey and Walker 1998) demonstrated that the composition of the samples from western Lhasa (Wang et al. 2014) would require an unreasonably high degree of crustal contamination (> 80%) (Fig. 32). Two other studies of ultrapotassic rocks from Italy and the Balkans (Conticelli et al. 2007; Prelević et al. 2015, respectively) attributed a similar combination of mantle sources (as opposed to crustal contamination) as being primarily responsible for the observed chemical and isotopic compositions. The recent study of Chu et al. (2013) also discussed the complex chemical and isotopic signatures preserved in the Wudalianchi-Erkeshan highly potassic basalts in the context of crustal and lithospheric contamination. Here, the range of 187Os/188Os in basalts

698

Gannoun, Burton, Day, Harvey, Schiano & Parkinson 0.20

>8 % crustal contamination

0.19

187

Os/188Os

0.18

lower continental crust

metasomatic sulfide

0.17

Wudalianchi-Erkeshan highly potassic basalts Keluo peridotite xenoliths

3.5 - 8 % crustal contamination

0.16

2 - 3.5 % crustal contamination

0.15 0.14

< 2% crustal contamination

0.13 fractional crystallization

0.12 0.11

addition of primary sulfide 0

10

20

1/[Os]

30

40

Figure 32. 187Os/188Os versus 1/Os for the Wudalianchi-Erkeshan highly potassic basalts, NE China. The solid lines represent binary mixing lines modeled as follows: Fields of crust addition to the intraplate basalts of 2%, 3.5%, and 8% lower continental crust are calculated using the values of Saal et al. (1988; 187 Os/188Os = 0.8 and Os concentration = 49 pg g−1). Metasomatic compositions are based upon mean values from Alard et al. (2002) and Sen et al. (2011). Mean primary sulfide compositions are taken from Alard et al. (2000, 2002), Pearson et al. (2002), Harvey et al. (2006, 2010, 2011), Lorand et al. (2013)—see also the supplementary information from Harvey et al. (2016, this volume). Modified after Chu et al. (2013).

Figure 32 (187Os/188Os = 0.1187–0.17) was partially attributed to 2–8 % crustal contamination; a degree of assimilation that otherwise would be difficult to detect using lithophile element isotope systems. In fact, Gannoun et al. (2015a) suggested that degrees of crustal assimilation of up to 15 % would have no measureable effect on Nd and Pb isotope ratios of basalts, while Li et al. (2014) commented that lithophile element-based isotope systems may be opaque to as much as 18 % crustal contamination. In the latter study, high NiO and SiO2 contents, but low MnO, CaO, MgO, and Pb contents, in addition to radiogenic 187Os/188Os, low Os abundances (5–43 ng g−1) and high, but variable, Re/Os (3–126) of intra-continental OIB-like basalts from West Qinling, central China, were attributed to crustal contamination on the strength of the sensitivity of Os isotope systematics to the incorporation of continental crust.

In contrast, the most unradiogenic Os isotope signatures observed in CIAV may have been affected by the assimilation of xenocryst-hosted primary sulfide. The often unradiogenic 187 Os/188Os and high (> µg g−1) Os content of sulfides enclosed within olivine xenocrysts (Alard et al. 2002) are prime candidates for the source of a possible “nugget effect”. For example, a 20 μg mantle sulfide with an Os concentration of 20 µg g−1 (see Alard et al. 2000, 2002; Pearson et al. 2002; Harvey et al. 2006, 2010, 2011; Lorand et al. 2013; Harvey et al. 2016, this volume for typical sulfides) contains twice as much Os as 2 g of basalt with an Os concentration of 100 pg g−1. This type of nugget effect was attributed by Chu et al. (2013) as being responsible for the poor reproducibility of 187Os/188Os in two WudalianchiErkeshan basalts (LHS-6 and HSS-6). In this instance, the heterogeneous distribution of a component that contains anomalously high Os (+ PGE) abundances throughout the sampled rock powder could account for the observed heterogeneities in replicate basalt analyses. A similar source of heterogeneity was suggested by Gannoun et al. (2015a) to account for comparable unradiogenic 187Os/188Os signatures in some Cameroon Line basalts.

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Processes affecting the HSE compositions of sub-aerial volcanism The previous sections demonstrate that it is essential to consider the many possible source and contamination factors that may influence the ultimate composition of intraplate magmas. Irrespective of the tectonic setting in which an erupted magma was generated, sub-aerially erupted lavas may be subject to an additional group of processes whose affects need to be assessed prior to interpretations concerning magma sources and potential contaminants. These processes, including post-emplacement alteration and magmatic degassing, were reviewed comprehensively in Day (2013). While there has been a dearth of new data in the intervening period, one study in particular merits attention; the recent examination of Os loss through magmatic degassing at Piton de la Fournaise, Réunion Island (Gannoun et al. 2015b). Oceanic island basalts have lower Re concentrations than MORB. This is anomalous considering the incompatible behavior of Re during basalt petrogenesis (Hauri and Hart 1997). This apparent quirk has been attributed to two possible causes; (i) the presence of garnet and/ or sulfide in their mantle source (Righter and Hauri 1998), or (ii) magmatic degassing of Re (Bennett et al. 2000; Lassiter 2003; Norman et al. 2004). Several lines of evidence support the idea that Re loss is a late and shallow stage process, which favors process (i) above. For example, an increase in oxygen fugacity promotes the loss of Re from Re metal (Borisov and Jones 1999), suggesting that at the oxidation state relevant to OIB (FMQ), the rate of Re loss from a magma will increase by an order of magnitude per log unit of fO2 increase. Sub-aerial eruptions from Réunion and Hawaii preserve evidence for an increase in fO2 in the lavas during emplacement, from FMQ − 1.8 close to eruption vents, to up to FMQ + 3 in lava samples that have travelled several km and cooled slowly (Rhodes and Vollinger 2005; Boivin and Bachélery 2009). Although Re and Os have the highest elemental condensation temperature (1821 and 1812 K, respectively; Lodders 2003), these elements are commonly enriched in volcanic gas sublimates and aerosols (Crocket 2000; Yudovskaya et al. 2008; Mather et al. 2012). However, the relative and absolute volatilities of Re and Os, and hence the degree of degassing from sub-aerial lavas, are not well constrained. The propensity for an elemental species to be volatilized post-eruption can be described in terms of an emanation coefficient, (Ex), where Ex = (Ci – Cf) / Ci, (Ci = concentration of element x in the magma and Cf = concentration of element x in the magma after degassing; Gill et al. 1985; Lambert et al. 1986). The emanation coefficient of Re ranges from 0.12 (Rubin 1997) to as high as 0.74 Norman et al. (2004). The difficulties associated with the analysis of pg g−1 quantities of Os in basalts make the emanation coefficient of Os even less well known. In their recent study, Gannoun et al. (2015b) investigated the Re–Os isotope and elemental systematics of basaltic lavas and gas condensates (a range of Na–K–Ca–Cu sulfates, Ca–Mg– Al–Fe fluorides, and native sulfur) produced during eruption and degassing at Piton de la Fournaise, Réunion Island, in order to examine the geochemical behavior of these two elements during magma degassing. High temperature (> 350 °C) deposits were enriched in Re (24–79 ng g−1), almost two order of magnitude higher than the corresponding lavas (0.130– 0.137), while the Os abundances of the high temperature condensates were similar to those of the lavas (14–132 pg g−1). The highest temperature condensates (Na–K sulfates; 384 to 400 °C), yielded 187 Os/188Os that were significantly lower (i.e., 0.124–0.129) than their corresponding lava. These unradiogenic osmium isotope ratios were attributed by Gannoun et al. (2015b) to the volatilization of Os originally contained in old, unradiogenic mantle sulfides. Sulfides associated with earlier volcanic eruptions at Réunion Island (< 7 Ma) were deemed too young to provide the distinctive unradiogenic Os fingerprint of the volcanic gas, leading Gannoun et al. (2015b) to infer that the observed unradiogenic Os was ultimately derived from a mantle source. In the context of osmium mantle geochemistry, loss of unradiogenic Os during magmas degassing could help to explain osmium isotope disequilibrium between lavas and melting residues.

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This contrasted with the 187Re–187Os systematics of the low-to-medium temperature condensates, which contained the highest Os abundances (13–77 ng g−1) with unfractionated 187 Os/188Os (0.130–0.135), which are indistinguishable from the April 2007 lava flow and the historical lavas of Piton de la Fournaise (i.e., 187Os/188Os = 0.130–0.137; Schiano et al. 2012). In addition, very high concentrations of iridium (1–8 ng g−1) reported for hieratite condensates (K2SiF6) suggested that Ir was also transported in volatile emissions as gaseous IrF6 (cf. Toutain and Meyer 1989). The selective enrichment of HSE demonstrates their potential for transport as metallic hexafluorides (Molski and Seppelt 2009; Craciun et al. 2010; Gannoun et al. 2015b; see also review in Day 2013). The absence of isotopic fractionation between gas deposits and lavas also indicates that external components (such as seawater, rainwater or air), which all possess particularly radiogenic 187Os/188Os (Levasseur et al. 1998, 1999; Gannoun et al. 2006; Chen et al. 2009) have no significant influence on the Os budget of volcanic gases.

HIGHLY SIDEROPHILE ELEMENT SYSTEMATICS OF ARCS Highly siderophile element abundance studies have been applied to arc volcanism to understand both subduction processes and the generation of economic deposits of precious metals within arc settings. A critical question has regarded the potential mobility of the HSE in subduction zone environments and the collateral effects such processes have regarding the siderophile element budget of the mantle. Fractionation of Re and Pt from Os in subduction zone environments could have a potentially significant effect both on Os isotope signatures at arcs (e.g., Brandon et al. 1996), but also on the long-term Re/Os and Pt/Os fractionations observed in OIB, MORB and mantle rocks. In addition, the potential mobility of HSE in subduction zone environments has important implications regarding the formation of economic PGE ore deposits such as major epithermal gold deposits associated with some volcanic arcs (e.g., McInnes et al. 1999). For the purpose of this review, we focus on the petrogenetic implications of arc volcanism. Arc magmatism includes tholeiitic to calc-alkaline compositions and dominantly involves the generation of basalt-andesites, andesites and more evolved magma-types. Only a limited number of arc volcanoes are known to erupt lavas approaching basaltic or picritic compositions. Because the HSE are typically compatible during mantle melting, as well as during fractional crystallization, this means that Os concentrations in arc volcanic rocks are typically very low, resulting in increased susceptibility of arc lavas to crustal contamination (e.g., Lassiter and Luhr 2001; Hart et al. 2002; Righter et al. 2002; Turner et al. 2009; Bezard et al. 2015). High 187 Os/188Os in arc lavas has therefore been attributed to assimilation of arc crust during magmatic ascent, but also due to enrichment in radiogenic Os due to contamination of the mantle wedge by slab-derived fluids/melts (e.g., Alves et al. 1999, 2002; Borg et al. 2000), or a combination of these processes (Suzuki et al. 2011). In this section, we review the work done so far in arcs, using both lavas, as well as mantle-derived xenoliths erupted in association with active arcs. Since the behavior of the HSE are reviewed extensively elsewhere in this volume, the focus of this section is largely on the information that can be obtained from the HSE regarding arc processes.

HSE and 187Os/188Os in arc lavas The majority of arc related volcanism is located around the Pacific ‘Ring of Fire’, extending from the southern tip of Chile, up much of South and North America, into the Aleutians and Kamchatka, through Japan and down as far as the Tonga Trench and New Zealand. Other significant arcs include the Lesser Antilles Arc and the Scotia Arc (Fig. 33). Despite the extensive distribution of arc volcanoes, limited work has been conducted on ReOs isotopes in arc volcanic rocks, primarily due to the limited availability of high MgO lavas, which are normally favored for study by Os isotope and HSE abundance studies. High MgO lavas do occur in some arc settings, most notably Grenada, south Lesser Antilles Arc, and as boninite occurrences. These lavas are discussed in detail, below.

HSE and Os-isotope Systematics in Volcanic Rocks

701

Hawaii Cameroon Volcanic Line

Figure 33. World map showing locations of major convergent margin settings (stippled lines) and divergent boundary settings (grey lines) mentioned in the text.

Work on arcs has shown that arc volcanic rocks typically contain between 0.00005 and Figure 33 1 ng g−1 Os and 0.01–1 ng g−1 Re (Fig. 34). Rhenium concentrations generally increase with decreasing MgO in arc lavas, consistent with moderate incompatibility of Re. However, Re can also behave as a volatile element during oxidizing conditions in arc lavas, and for this reason it is likely that low concentrations could reflect loss of Re by this process (e.g., Righter et al. 2008). Positive correlation between Os and MgO is consistent with strong compatibility of Os during fractional crystallization of arc lavas. The low MgO and HSE contents in arc lavas can make them potentially highly susceptible to crustal contamination effects (cf. Lassiter and Luhr 2001). Osmium isotopic ratios in recently erupted arc lavas can span an extreme range, from high MgO lavas with 187Os/188Os (~0.1268–0.128) similar to typical mantle estimates, to andesites, rhyolites and dacites with 187Os/188Os > 1. There is an overall relationship of increasing 187Os/188Os with decreasing Os content, although more than one trend has been recognized in plots of reciprocal Os versus 187Os/188Os (Fig. 35). Alves et al. (2002) pointed out that initial Os isotopic ratios are positively and systematically correlated on 187Os/188Os versus reciprocal Os plots, reflecting binary mixing processes, with a common end-member represented by upper mantle peridotite compositions To date, no study has found clear associations of Re or Os contents and 187Os/188Os with arc basement type, convergence rate or sediment supply (Table 1). This may be partly due to the lack of available high MgO rocks with which to make cross-comparison of ‘primary magmatic composition’. For example, Lassen Peak lavas with 8–11.1 wt% MgO have up to 0.37 ng g−1 Os and span a range of 187Os/188Os from 0.1289–0.235 (Borg et al. 2002). It has been suggested that these lavas contain a contribution of radiogenic Os from the subducting slab. Conversely, Grenada picrites and basalts (10.5–17.4 wt% MgO) contain up to 0.36 ng g−1 Os and have a slightly more restricted range of 187Os/188Os (0.1268–0.1644), yet these lavas are not considered to have a contribution from the slab, but instead have experienced various levels of crustal assimilation (Woodland et al. 2002; Bezard et al. 2015). Likewise, boninite (13 wt%) and some low MgO lavas (< 1.5 wt%) from the Tonga-Kermadec arc have 187Os/188Os of 0.1275–0.1283, indicating that more radiogenic values for lavas in this arc are consistent with localized arc contamination (Turner et al. 2009). Unique to that study is that the sample with the least radiogenic Os signature is a dacite, suggesting that evolved magmas can develop by fractionation from mantle-derived magma with minimal interaction with high Re/Os arc crust.

Gannoun, Burton, Day, Harvey, Schiano & Parkinson 702

25

Crustal thickness

0.0003–0.362

0.0002–0.03

Os (ppb)

0.051–1.62

0.01–0.43

0.067–0.457

Re (ppb)

0.131–2.11

0.157–3.15

0.1268–0.811

0.192–1.45

Os/188Os

10

30

31

18

n

Table 1. Example of Re–Os isotopes and geophysical characteristics at selected subduction zones. (Data sources provided in the text.). Crustal basement

30–35

0.009–94.1

187

Oceanic

Oceanic

0.004–0.52

0.0002–0.01

(km)

235

–

(m)

1750

30

Sediment thickness above AOC

86

Oceanic

Transitional

(Ma)

100–150

300

Age of downgoing AOC

1.4

1500

1

Convergence Rate

1.4 50

138

18

(cm/yr)

South

North

9–14

7.6–7.9

Subduction Zone

Java

0.258

Lesser Antilles

Papua New Guinea

0.14–1.181

22

0.399

0.133–0.246

7

0.13–0.73

0.074–0.92

0.1378–0.319

1

29

0.0008

0.0003–0.36

0.024–0.73

0.279

0.1277–0.371

0.00005–0.003 25–45

0.001–0.023

0.187

0.082–0.527

–

18–25

0.0017

0.006–0.93

15 to 30 Accreted arc terrane

30

Transitional

Oceanic

66

120

364

Continental

600

350

Continental

50

90

170

146

54

270

9

8.9–9.2

15

6.7–9.6

Kamchatka 7–8.7

15

Izu-Bonin

Aleutians

8.4–8.9

5.7 –8.5

Phillipines

Mexico

Oceanic/ Continental

Columbia

40–70

6

Continental

0.133–1.524

125

0.1–0.23

26–82

0.0012–0.021

10.3–10.8

Peru-Chile

North Lesser Antilles = Saba, Redonda, Monserrat, Guadeloupe, Dominica South Lesser Antilles = Martinique, St. Lucia, St. Vincent, Grenada

HSE and Os-isotope Systematics in Volcanic Rocks

703

Os (ng g-1)

1 0.1 0.01 0.001 0.0001 10

Re (ng g-1)

1 0.1 0.01 0.001

0

2

4

6

8

10

12

14

16

18

MgO (wt. %) Figure 34. Plots of Os and Re versus MgO content for convergent margin picrites, basalts and evolved rocks. Data sources: Brandon et al. (1996); Alves et al. (1999, 2002); Borg et al. (2000); Woodland et al. (2002); Woodhead and Brauns (2004); Turner et al. (2009); Bezard et al. (2015).

Figure 34

Contents of the HSE in arc-related lavas have been reported for Grenada basalts and picrites, Izu-Bonin lavas (Woodland et al. 2002) and Lihir lavas (McInnes et al. 1999) (Fig. 36). These generally high MgO lavas show similar Re and PPGE enrichment over the IPGE, to many intraplate tholeiites and alkali basalts (e.g., Day 2013). However, despite the picritic (MgO > 13·5 wt%) nature of Grenada lavas, they contain low concentrations of the HSE (< 0·2 ng g−1 Ir, 1–4 ng g−1 Pd) compared with lavas from other settings of similar MgO content (see section on continental flood basalts). Woodland et al. (2002) argued that this was probably due to a combination of lower degrees of partial mantle melting and early removal of PGE with cumulus phases such as olivine, magnetite and sulfide. Comparison of alkali Grenada lavas with boninitic Izu–Bonin lavas illustrates that although the major element chemistries of Grenada and Izu–Bonin are different, relative and absolute abundances of the IPGE and PPGE are similar. Rhenium, however, is markedly depleted in the Grenada picrites compared with the Izu–Bonin boninites, suggesting either retention of Re by residual garnet in the Grenada sub-arc mantle wedge (Woodland et al. 2002) or volatile-loss of Re. In both cases, their generation above a subduction zone did not appear to have any significant systematic effect on the HSE signatures of resultant lavas.

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Gannoun, Burton, Day, Harvey, Schiano & Parkinson 3.5

187

Os/188Os

3.0 2.5 2.0 1.5 1.0 0.5 0.00001

0.0001

0.001

0.01

0.1

1

20000

25000

Os (ng g-1)

187

Os/188Os

3.5 3.0 2.5 2.0 1.5 1.0 0.5 0

5000

10000

15000

1/[Os] Figure 35. Plots of Os and reciprocal Os (1/Os) versus 187Os/188Os for convergent margin picrites, basalts and evolved rocks. Data sources: Brandon et al. (1996); Alves et al. (1999, 2002); Borg et al. (2000), Woodland et al. (2002); Woodhead and Brauns (2004); Turner et al. (2009); Bezard et al. (2015).

HSE and 187Os/188Os in arc xenoliths Studies of mantle xenoliths from arc settings have provided the opportunity to document the behavior of the HSE during slab fluid-induced metasomatism of the mantle wedge, with spinel Figure harzburgite, websterite and pyroxenite mantle xenoliths occurring in back-arc environments in a number of arcs. Relatively radiogenic Os isotope signatures in mantle xenoliths and mantle rocks from arc settings, including the Cascades, Canadian Cordillera, Japan, Lihir, Papua New Guinea, Kamchatka, and the Catalina Schist have been documented, and attributed to the mobility of Os in slab fluids (Brandon et al. 1996, 1999; McInnes et al. 1999; Peslier et al. 2000; Widom et al. 2003). For example Simcoe xenoliths, which represent fragments of mantle lithosphere from the back-arc of the Cascade arc front, have been metasomatized by silica-rich

35

HSE and Os-isotope Systematics in Volcanic Rocks 1000

100

Grenada picrites & basalts

[ Sample × 1000] / CI-chondrite

10

Kamchatka xenoliths and arc sediments

10

1

1 0.1

0.1

0.01

0.01 1000 100

100

705

Lihir mantle xenoliths and lavas

0.001 1000 100

Izu-Bonin lavas

10

10 1

1

0.1

0.1

0.01

0.01

0.001

0.001

Re Pd Pt Ru Ir

Os

Re Pd Pt Ru Ir

Os

Figure 36. CI-chondrite normalized HSE diagrams for Lihir mantle xenoliths and lavas (McInnes et al. 1999), Grenada picrites and basalts and Izu-Bonin lavas (Woodland et al. 2000), and Kamchatka xenoliths (Widom Figure 36 et al. 2003). Note that PPGE > IPGE lavas from the Izu-Bonin and Grenada. CI chondrite normalization from Horan et al. (2003). HSE are plotted with more compatible elements to the right (see Day, 2013 for discussion).

fluids or hydrous melts leading to higher fO2 leading to radiogenic Os isotopic compositions being imparted to these peridotites (Brandon et al. 1996, 1999). These features are consistent with part or the entire metasomatic agent being derived from the Juan de Fuca slab. Studies of Kamchatka peridotites also indicate metasomatism of the Kamchatka sub-arc mantle wedge by radiogenic slab-derived fluids and melts (Widom et al. 2003). The HSE patterns of the arc-related mantle xenoliths are broadly similar to typical oceanic mantle xenoliths (Fig. 36), but the xenoliths can often exhibit elevated 187Os/188Os, with Simcoe xenoliths ranging from 0.1226–0.1566 and Kamchatka xenoliths ranging from 0.1232–0.1484. The regional variations in Re–Os isotope signatures are consistent with previous petrographic and geochemical studies of the Kamchatka mantle xenoliths that reveal multistage metasomatic histories resulting from interaction of the mantle wedge with a variety of slab-derived fluids and melts, including silicic slab–melt metasomatism associated with subduction of relatively hot, young (∼15–25 Ma) oceanic crust in the northern arc front, hydrous slab-fluid metasomatism associated with subduction of colder, old (∼100 Ma) oceanic crust in the southern arc front, and carbonate-rich slab-melt metasomatism in the southern segment behind the arc front, where the slab is deeper. Similar ranges of Re–Os isotope signatures in peridotites from Avachinsky, Japan and Lihir, and from Valovayam and the Cascades, respectively, suggest that the age (temperature) and depth of subducting oceanic crust influences the Re–Os composition of metasomatized sub-arc mantle.

Radiogenic Os from slab components or from crustal contamination A continuing debate exists over the influence of slab-derived 187Os/188Os to arcs, versus the potential for crustal or seawater contamination of magmas with low Os abundances. From Lassen lavas, Borg et al. (2000) showed that crustal contamination could only explain the ReOs isotope systematics if distribution coefficients for Re in sulfide were ~40–1100 times higher than published estimates, and instead argued for contributions from a highly radiogenic Os slab component (187Os/188Os up to 1.4). Alves et al. (2002) also favoured slab components adding radiogenic Os to arcs, citing evidence from arcs worldwide for different mixing systematics between mantle peridotite and variably radiogenic Os slab contributions. Conversely, Bezard et

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al. (2015) have shown that Grenada picrites with radiogenic 87Sr/86Sr (0.705) have 187Os/188Os (0.127) that overlap with the mantle range and that assimilation and fractional crystallization can explain compositions of Lesser Antilles lavas, without the requirement of a slab input (Fig. 37). Dreher et al. (2005) studied Os isotopes in Mindanao adakites, showing that the majority of these rocks had unradiogenic Os isotopes, inconsistent with the idea that adakites with high Sr/Y and low Y and heavy rare earth element concentrations, reflect melting of young subducted crust in subduction zones. On the other hand, the range in Os isotopes in Mexican Volcanic Belt rocks, which represent subduction-related calc-alkaline and lamprophyric rocks in which high fO2 precludes sulfide fractionation, could be explained by up to 12% assimilation and fractional crystallization (Lassiter and Luhr 2001). To obviate potential issues of shallow-level crustal contamination, Suzuki et al. (2011) examined Cr-spinel from beach sands in the Bonin Islands, reasoning that Cr-spinel is an early-formed mineral in most magmas and an indicator of primitive magma Os compositions. They found unradiogenic Os in Cr-spinel with boninitic affinity, versus a potential slab component reflected in spinel with tholeiitic affinity. These authors also argued that oxidative conditions in the mantle can lead to radiogenic Os mobilization in the arc. Ultimately, the most convincing arguments for or against radiogenic Os from the slab comes from high-MgO Grenada picrites. These samples have been shown to have less-radiogenic Os signatures in more mafic lavas, with an increasing influence of crustal contamination in more evolved melts (Woodland et al. 2002; Bezard et al. 2015). Combined with evidence for the potential influence of subduction zone fluids on the composition of arc xenoliths, these results suggest that some contribution from the slab can be exhibited in arc lavas, but that the role of crustal contamination of melts within the arc itself can obfuscate original mantle-derived signatures. 0.22

187

Os/188Os

0.20

Model 2 (D Sr = 1.8/Os = 12) Model 1 (D Sr = 1.4/Os = 16)

0.18 0.16 0.14 0.12 0.703

0.704

0.705

0.706

Sr/86Sr

87

Figure 37. Assimilation accompanied by fractional crystallization (AFC) models of 87Sr/86Sr versus 187 Os/188Os for Lesser Antilles primitive lavas Parameters for Models 1 and 2 are shown in the figure and in Table 3 of Bezard et al. (2015).

Mechanical mixing processes The debate as to whether slab-derived signatures are evident in HSE and Os isotopes Figure 37mechanical within arc volcanic rocks has recently been enhanced by the recognition that mixing between peridotite mantle and recycled ocean rocks is likely an important process in modifying HSE contents at subduction zones. Studies of HSE contents and Os isotope compositions of mélange mafic metamorphic blocks at Catalina Island and the Franciscan

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Sample/PM

1

707

Os/188Os

187

0.1240.192

0.1

0.01 0.214.75

Average Core Average Rind

0.001

Re

Pd

Pt

Ru

Ir

Os

Figure 38. Primitive mantle normalized HSE diagram for average core and rind compositions of Mèlange metamorphic mafic blocks from the Catalina Schist, Franciscan Complex and Samana Metamorphic Complex. Cores are consistent with dominantly reflecting basaltic/sedimentary protoliths with radiogenic Os and rinds represent as much as 70% peridotite HSE contributions. Data are from Penniston-Dorland et al. (2012, 2014), with primitive mantle normalization from Becker et al. (2006).

Complex (California) and at the Samana Metamorphic Complex (Dominican Republic) have shown significant differences between block cores and block rims (Penniston-Dorland et al. 2012, 2014). In particular, while the cores of the blocks have enhanced PPGE compared with rocks, 38 or some IPGE and radiogenic 187Os/188Os, mimicking patterns for evolved basalticFigure sedimentary protoliths, the rims approach HSE contents expected in some mantle peridotites, with less radiogenic 187Os/188Os than the cores (Fig. 38). Penniston-Dorland et al. (2014) have demonstrated that mélange mechanical mixing occurs across a range of temperatures (≤ 200 to ~ 600 °C) during subduction leading to a hybrid rock composition of peridotite, basaltic materials and sediments. Measurements of the HSE in arc volcanics suggest variable amounts of peridotitic mantle with radiogenic Os components (e.g., Alves et al. 1999, 2002; Borg et al. 2000) and mechanical mixing may play a major role in this process.

CONCLUSIONS AND PERSPECTIVES The highly siderophile elements are expected to be strongly incorporated into Earth’s metallic core, but their abundance in the upper mantle appears to have been set by the late addition of meteoritic material after core formation was complete. Partial melting of the mantle since that time has resulted in a significant fractionation of the HSE. The platinum-PGE, Re and Au, can behave as moderately compatible or incompatible elements during melting, and may be variably enriched in melts,while the Iridium-PGE behave as highly compatible elements. Sulfide appears to be a major host for HSE in mantle rocks, despite its relatively low abundance (between 0.04 and 0.08%). However, sulfide cannot account for the fractionation of HSE that occurs during the melting that generates MORB, which generally possess very low Os–Ir–Ru contents, and relatively high Re–Pd and Pt. Rather this fractionation appears to result from the crystallization of Os–Ir–Ru alloy phases in refractory mantle rocks, accompanying the exhaustion of sulfide by melting. The HSE content of MORB is further modified by the segregation of sulfide during fractional crystallization in the magmatic environment, where the HSEs are quantitatively removed into sulfide, leaving the residual melt depleted in these elements. The fractionation of Re and Os accompanying the generation of MORB, intraplate lavas and those produced at convergent margins is one of the key processes controlling the distribution of these elements between Earth’s mantle and crust. Therefore, decay of 187Re to 187Os provides an

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exceptional tracer of recycled crustal materials in Earth’s mantle. This is because oceanic and continental crust possess high Re/Os ratios, and develop radiogenic Os isotope compositions over time, which in turn can be readily traced as recycled material if mixed back into the convective mantle. However, while MORB glass commonly preserves a radiogenic 187Os/188Os composition, this is most readily explained by seawater-derived contamination of the melt that occurs during magma ascent through the oceanic crust. Although reliable data for MORB glass remain limited these observations suggest that to a greater or lesser extent all MORB glass has been affected seawater contamination. This then also implies that other elements may have been affected by such contamination, most likely dependent upon their relative concentration in MORB glass and seawater. Sulfide, although demonstrably affected by the same seawater contamination, provides a more reliable record of the primary 187Os/188Os isotope composition of MORB, particularly those sulfides with high Os concentrations (i.e., > 100 ppb). These high-Os sulfides preserve relatively unradiogenic 187Os/188Os isotope compositions pointing to a mantle source that has experienced long term depletion of Re, similar to abyssal peridotites, with no evidence for the presence of recycled crust. In addition to the effects of seawater contamination observed in MORB, intraplate lavas and those generated at convergent margins may interact with sub-continental lithospheric mantle, itself variably contaminated by multiple metasomatic events since it became isolated from the convecting mantle, and incorporate additional complications from the overlying crust. At convergent margins there is the additional complication of fluxes generated as a result of the subduction of the down-going slab with the potential for overprinting pre-existing Re–Os isotope and HSE fingerprints. While the HSE and its isotope systems offer some unique perspectives on mantle processes and the generation of a wide range of magmas, their application needs to be exercised with care—the geochemical context provided by other isotope systems and trace element signatures should be considered and the specific set of local conditions, both physical and chemical, taken into account in addition to the use of these invaluable tools.

ACKNOWLEDGMENTS We thank Jean-Louis Birck, Olivier Alard, Christian Pin, Ivan Vlastélic, Anthony Cohen, Ali Bouhifd for valuable insight and discussions over the years. AG would like to thank Jean-Luc Devidal for assistance with the electron microprobe measurements at Blaise Pascal University. The authors thank Chris W. Dale for a review that greatly improved the manuscript. This research was partially financed by the French Government Laboratory of Excellence initiative n°ANR-10-LABX-0006, the Région Auvergne and the European Regional Development Fund. JD acknowledges the support of NSF (NSF-EAR grant 1447130). JH was supported by a NERC Advanced Research Fellowship (NE/J017981/1), a Blaustein Visiting Professorship at Stanford University, and a visiting investigator appointment at the Carnegie Institution, Washington. This is Laboratory of Excellence ClerVolc contribution number 178.

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Fresenius J Anal Chem 341:537–41 Walker RJ (2009) Highly siderophile elements in the Earth, Moon and Mars: update and implications for planetary accretion and differentiation. Chemie der Erde-Geochem 69:101–125 Walker RJ, Fassett JD (1986) Isotopic measurement of sub-nanogram quantities of rhenium and osmium by resonance ionization mass spectrometry. Anal Chem 58:2923–2927 Walker RJ, Carlson RW, Shirey SB, Boyd FR (1989) Os, Sr, Nd, and Pb isotope systematics of Southern African peridotite xenoliths; implications for the chemical evolution of subcontinental mantle. Geochim Cosmochim Acta 53:1583–95 Walker RJ, Morgan JW, Beary E, Smoliar MI, Czamanske GK, Horan MF (1997) Applications of the 190Pt-186Os isotope system to geochemistry and cosmochemistry. Geochim Cosmochim Acta 61:4799–4808 Walker RJ, Storey M, Kerr AC, Tarney J, Arndt NT (1999) Implications of 187Os isotopic heterogeneities in a mantle plume: evidence from Gorgona Island and Curaçao. Geochim Cosmochim Acta 63:713–728 Wallace P, Carmichael ISE (1992) Sulfur in basaltic magmas. Geochim Cosmochim Acta 56:1863–1874 Wang BD, Chen JL, Xu JF, Wang LQ (2014) Geochemical and Sr–Nd–Pb–Os isotopic compositions of Miocene ultrapotassic rocks in southern Tibet: Petrogenesis and implications for the regional tectonic history. Lithos 208–209:237–250 Warren JM, Shirey SB (2012) Lead and osmium isotopic constraints on the oceanic mantle from single abyssal peridotite sulfides. Earth Planet Sci Lett 359–360:279–293 Watson EB, Baker DR (1985) Chemical diffusion in magmas: an overview of experimental results and geochemical applications. Adv Phys Geochim 9:120–151 Watson EB, Ben Othman D, Luck J-M, Hofmann AW (1987) Partitioning of U, Pb, Cs, Yb, Hf, Re and Os between chromian diopsidic pyroxene and haplobasaltic liquid. Chem Geol 62:191–208 White WM (2010) Oceanic island basalts and mantle plumes: the geochemical perspective. Annu Rev Earth Planet Sci 38:133–160 White R, McKenzie D (1989) Magmatism at rift zones: the generation of volcanic continental margins and flood basalts. J Geophys Res 94:7685–7729 White WM, Schilling JG (1978) The nature and origin of geochemical variation in Mid-Atlantic Ridge basalts from the Central North Atlantic. Geochim Cosmochim Acta 42:1501–1516 Widom E (1997) Sources of ocean island basalts: a review of the osmium isotope evidence. Physica A 244:484–496 Widom E, Shirey SB (1996) Os isotope systematics in the Azores: implications for mantle plume sources. Earth Planet Sci Lett 142:451–465 Widom E, Hoernle KA, Shirey SB, Schmincke H-U (1999) Os isotope systematics in the Canary Islands and Madiera: lithospheric contamination and mantle plume signatures. J Petrol 40:279–296 Widom E, Kepezhinskas P, Defant M (2003) The nature of metasomatism in the sub-arc mantle wedge: evidence from Re–Os isotopes in Kamchatka peridotite xenoliths. Chem Geol 196:282–306 Wilson JT (1963) A possible origin of the Hawaiian Islands. Can J Phys 41:863–870 Wilson M (1997) Thermal evolution of the Central Atlantic passive margins: continental break-up above a Mesozoic superplume. J Geol Soc, London 154:491–495 Wilson AH, Shirey SB, Carlson RW (2003) Archaean ultra-depleted Komatiites formed by hydrous melting of cratonic mantle. Nature 423:858–860 Woodhead J, Brauns M (2004) Current limitations to the understanding of Re–Os behavior in subduction systems, with an example from New Britain. Earth Planet Sci Lett 221:309–323

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Woodhead JD, McCulloch MT (1998) Ancient seafloor signals in Pitcairn Island lavas and evidence for large amplitude, small length-scale mantle heterogeneities Earth and Planetary Science Letters 94(3–4):257–273 Woodland SJ (2000) Development of ICP MS isotope dilution preconcentration techniques for determination of platinum group elements in volcanic rocks. PhD thesis University of Durham Woodland SJ, Pearson DG, Thirlwall MF (2002) A platinum group element and Re–Os isotope investigation of siderophile element recycling in subduction zones: comparison of Grenada, Lesser Antilles Arc, and the IzuBonin arc. J Petrol 43:171–198 Workman RK, Hart SR, Jackson MG, Regelous M, Farley KA, Blusztajn J, Kurz M (2004) Recycled metasomatised lithosphere as the origin of the Enriched Mantle II (EM2) end-member: evidence from the Samoan Volcanic Chain. Geochem Geophys Geosys 5:Q04008 http://dx.doi.org/10.1029/2003GC000623 Workman RK, Hart SR (2005) Major and trace element composition of the depleted MORB mantle (DMM). Earth Planet Sci Lett 231:53–72 Wright E, White WM (1987) The origin of Samoa: new evidence from Sr, Nd and Pb isotopes. Earth Planet Sci Lett 81:151–162 Xu JF, Suzuki K, Xu YG, Mei HJ, Li J (2007) Os, Pb, and Nd isotope geochemistry of the Permian Emeishan continental flood basalts: insights into the source of a large igneous province. Geochim Cosmochim Acta 71:2104–2119 Yang AY, Zhao T-P, Zhou M-F, Deng X-G, Wang G-Q, Li J (2013) Os isotopic compositions of MORBs from the ultra-slow spreading Southwest Indian Ridge: Constraints on the assimilation and fractional crystallization (AFC) processes. Lithos 179:28–35 Yang AY, Zhou M-F, Zhao T-P, Deng X-G, Qi L, Xu J-F (2014) Chalcophile elemental compositions of MORBs from the ultraslow-spreading Southwest Indian Ridge and controls of lithospheric structure on S-saturated differentiation. Chem Geol 382:1–13 Yaxley GM (2000) Experimental study of the phase and melting relations of homogeneous basalt + peridotite mixtures and implications for the petrogenesis of flood basalts. Contrib Mineral Petrol 139:326–338 Yaxley GM, Green DH (1998) Reactions between eclogite and peridotite: mantle refertilization by subducted oceanic crust. Schweiz Mineral Petrogr Mitt 78:243–255 Yokoyama T, Walker D, Walker RJ (2009) Low osmium solubility in silicate at high pressures and temperatures. Earth Planet Sci Lett 279:165–173 You CF, Spivack AJ, Gieskes JM, Martin JB, Davisson ML (1996) Boron contents and isotopic compositions in pore waters: a new approach to determine temperature induced artefacts-geochemical implications. Mar Geol 129:351–361 Yudovskaya MA, Tessalini S, Distler VV, Chaplygin IV, Chugaev AV, Dikov YP (2008) Behaviour of highly siderophile elements during magma degassing: a case study at the Kudryavy volcano. Chem Geol 248:318–341 Zhang M, Suddaby P, O’Reilly SY, Norman M, Qiu J (2000) Nature of the lithospheric mantle beneath the eastern part of the Central Asian fold belt: mantle xenolith evidence. Tectonophysics 328:131–156 Zhang YL, Liu CZ, Ge WC, Wu FY, Chu ZY (2011) Ancient sub-continental lithospheric mantle (SCLM) beneath the eastern part of the Central Asian Orogenic Belt (CAOB): implications for crust–mantle decoupling. Lithos 126:233–247 Zhang Z, Zhi X, Chen L, Saunders AD, Reichow MK (2008) Re-Os isotopic compositions of picrites from the Emeishan flood basalt province, China. Earth and Planetary Science Letters 276:30–39 Zindler A, Hart SR (1986) Chemical geodynamics. Annu Rev Earth Planet Sci 14:493–571 Zindler A, Staudigel H, Batiza R (1984) Isotope and trace element geochemistry of young Pacific seamounts: implications for the scale of upper mantle heterogeneity. Earth Planet Sci Lett 70:175–195

12

Reviews in Mineralogy & Geochemistry Vol. 81 pp. 725-774, 2016 Copyright © Mineralogical Society of America

Highly Siderophile and Strongly Chalcophile Elements in Magmatic Ore Deposits Sarah-Jane Barnes Sciences de la Terre Université du Québec à Chicoutimi Chicoutimi, Québec, G7H 2B1 Canada [email protected]

Edward M. Ripley Department of Geological Sciences Indiana University Bloomington Indiana, 47405 USA [email protected]

INTRODUCTION An ore deposit by definition must be economically viable, that is to say it must contain sufficient material at high enough grade to make it possible to mine and process it at a profit (Bates and Jackson 1987). This requires the elements to be collected and concentrated by some phase and for them to be deposited close to the surface of the earth. At the oxygen fugacities found in the crust, native Fe is not normally stable and thus the highly siderophile elements (defined as Ru, Rh, Pd, Re, Os, Ir, Pt, and Au) cannot behave as siderophile elements except in rare cases such as on Disko Island (Klöck et al. 1986) where the magma is sufficiently reduced for native Fe to be present. However, if mafic magmas become saturated in a base-metal-sulfide liquid, the highly siderophile elements behave as highly chalcophile elements (Table 1). Thus these elements are generally found in association with base-metal-sulfide minerals which crystallized from a magmatic sulfide liquid, namely pyrrhotite, pentlandite, chalcopyrite, cubanite ± pyrite. An exception to this is Au. Although Au is strongly chalcophile and is produced as a by-product from many platinum-group element (PGE) deposits (Table 2), most primary Au deposits consist of native Au (Groves et al. 1998). These will not be discussed in this chapter. There are many PGE-deposits (i.e., accumulations of PGE minerals and base metal sulfides containing PGE; Bates and Jackson 1987) around the world, but most of these do not constitute PGE ore deposits, because they are either too small or their grade is too low, or other political or infrastructure factors prevent the economic exploitation of the deposit (Bates and Jackson 1987) For the purpose of this work we have defined PGE ore deposits as those which have significant production (> 2% of the annual world production) of Pt or Pd (Fig. 1, data from Mudd 2012; Cowley 2013). Aside from these deposits, Pt and Pd are also produced as by-products from many magmatic Ni-deposits. But the amount from each deposit is small and, in total, for all deposits is ~ 2% of annual world production (Fig. 1). They will not be discussed further, but the behavior of the PGE in these types of deposits is similar to their behavior in the Noril’sk and Sudbury deposits and more details on them can be found in Naldrett (2011). Platinum deposits in zoned complexes, also known as Alaskan 1529-6466/16/0081-0012$10.00

http://dx.doi.org/10.2138/rmg.2016.81.12

Barnes & Ripley

726

Table 1. Partition Coefficients Sulfide/silicate liquid Experimental Min

Max

MSS/sulfide liquid

Empirical MORB

Experimental Min

ISS/sulfide liquid

Empirical

Max

Ag

300

970

1138

0.01

0.11

As

0.3

15

25

0.02

0.5

Au

2360

11200

967

0.0038

0.09

Bi

130

1130

316

0.003

0.0074

Experimental Min

0.38 0.1

Refs

Max

0.19

1.2

0.11

0.24

0.21

1

1,2,3,4,10, 24,25

0.026

0.13

1,10,11,12, 24,25

0.3–0.5

1,2,10,11,24,25 1,10,11,24,25

Cd

60

112

107

Co

20

114

45

0.92

1.6

1.6

2,25

Cu

330

2130

1334

0.06

0.36

0.07

1.00

2

1,2,4,9,10,12, 15,16,17,18, 19,20,21,23, 24

Ir

48000

1900000

11600

2.3

14.7

3.8-13

0.05

0.22

4,5,10,11,15, 16,18,19,20, 21,22,24,25

Mo

0.1

3.46

1.2

2.1

2.9

Ni

250

1700

776

0.36

1.72

1.1

0.1

0.9

Os

10000

1140000

24800

2

23

3.4–11

0.06

0.53

1,2,10,17, 22,25

1

0.05

1,2,10,15,16,17, 18,19,20,21, 22,23,24 4,5,10,11, 16,23,24,25

Pb

10

93

57

0.001

0.049

Pd

57000

536000

16735

0.06

0.24

0.13

0.3

0.7

Pt

4830

3450000

37300

0.04

0.03

0.004

0.125

0.487

Re

20

200000

870

1.6

8.5

4.3–9

0.054

0.11

4, 8,10,11,22

Rh

25000

591000

24300

1

11

2.7–8.3

0.055

0.15

4,10,11,15, 16,21,22,23

Ru

97700

485000

15500

1

19

0.96

0.083

0.84

4,10,11,15, 16,21,22,23

Sb

1.4

67

3.6

0.002

0.017

0.029

0.142

1,2,9,10,11,12

0.83

1.2

9,10,11,12,23

Se

226

2339

345

0.5

0.75

Sn

2.7

8.6

11

< 0.03

0.009

0.16

Te

1005

8789

4478

0.015

0.07

0.31

3.5

0.36

0.62

Zn

0.4

0.02

3.9

1,2,9,10,11 10,11,15,16, 21,22,23 4,6,7,10,11,15, 16,21,22,23

1,10,11 0.822

9,10,11,12 10,11

References: 1. Li and Audétat (2012); 2 Kiseeva and Wood (2013); 3. Li and Audétat (2013); 4 Mungall and Brenan (2014); 5 Fonseca et al. (2011); 6. Fonseca et al. (2009); 7. Pruseth and Palme (2004); 8. Brenan (2008); 9 Brenan (2015); 10 Patten et al. (2013); 11. Liu and Brenan (2015); 12. Helmy et al. (2010); 13. Helmy et al. (2013a); 14. Helmy et al. (2013b); 15. Mungall et al. (2005); 16. Fleet et al. (1993); 17. Sinyakova and Kosyakov (2012); 18. Sinyakova and Kosyakov (2007); 19. Sinyakova and Kosyakov (2009) 20. Sinyakova and Kosyakov (2014); 21. Barnes et al. (1997a); 22. Ballhaus et al. (2001); 23. Brenan (2002); 24. Thériault and Barnes (1998); 25. Barnes et al. (2006). Note: Bold type indicates coefficients calculated using average MORB to estimate the silicate liquid composition because the concentrations in the glass were less than detection limit. Italics minimum partition coefficients, calculated using whole rock values to estimate the silicate liquid composition.

Merensky UG2 Platreef Main Sulfide Zone JM reef JM reef Roby + Offset Roby + Offset All deposits All deposits All deposits

Bushveld

Bushveld

Bushveld

Great Dyke

Sillwater

Sillwater

Lac des Iles

Lac des Iles

Noril'sk-Talnakh

Noril'sk-Talnakh

Sudbury

Cu %

0.07

0.08 0.07

0.06

0.07

0.11

~ 0.12 0.1

0.11

~0.01–0.05

~ 0.1

0.18

~ 0.1

~ 0.2

PGE dominated deposits

Ni %

0.18

0.21

2.6

3.7

2.7

2.8

2.7

2.9

Pt ppm

1648

458

1665

1.2

0.89

0.73

1.08

1.84

1.36

0.5

1

1.36

Ni-Cu deposit with PGE by-product

16

58

46

149

2136

3418

6636

3733

Tons ×106

0.58

4.4

3.6

2.8

2.1

10.7

12.9

1.8

3.4

1.8

1.4

Pd ppm

0.05

0.49

0.5

0.2

Rh ppm

0.18

0.26

0.21

0.18

0.15

0.059

0.059

0.26

Au ppm

6

5

4

5

4

3

2

1

1

1

1

Type

13

12

12

11

11

10

9

7

8

7

7

Ref

Notes: 1. Measured, indicated and inferred; 2. Reserves and mineralized material; 3. Measured; 4. Measured and indicated; 5. Inferred; 6. Mined and reserves; 7. PGE+Au form D. Causey as quoted in Zientek (2012); Ni, Cu from Mudd (2012); 8. Mudd (2012); 9. Abott et al. as quoted in Zientek (2012); 10. Calculated using Sillwater web page www.stillwatermining.com - Investor presentation May 2015; 11. NAP web page Dec 2013 www.napalladium.com; 12. Norilsk Nickel web page for Dec 2013 www.nornick.ru; 13. Naldrett (2011)

Deposit Name

Host Intrusion

Table 2. The World’s Major Platinum-group Element Resources

Magmatic Ore Deposits 727

or Uralian complexes are not currently mined and will not be covered here. Information on these types of deposits is provided by Augé et al. (2005) and Anikina et al. (2014). Finally a small quantity of Pt (~ 2%, Cowley 2013) is produced from alluvial deposits, information on these can be found in Weiser (2002) and Tolstykh et al. (2004).

The majority of the world’s Pt and much of its Pd are produced from the Bushveld Complex of South Africa (Figs.1 and 2). These resources are present in three deposits; the UG2 reef, the Merensky reef and the Platreef (Table 2). The next most important source of Pd and Pt (Fig. 1) is as a by-product of nickel mining in the Noril’sk area of Russia. These Ni deposits occur in three sub-volcanic intrusions: Noril’sk 1, Talnakh, and Kharaelakh (Fig. 3). In fact the Noril’sk deposits produce most of the world’s Pd, although the actual Pd resources are less than those of the Bushveld Complex (Fig. 1, Table 2).

Barnes & Ripley

728 80 70

World Production of Pt and Pd in 2013

72

60

Pt Pd

50

40

37

% 40 30 20

7 5

10

2

0

6

14 3

0.3 2

Bushveld Great Dyke Stillwater Lac des Iles Norilsk

6

Sudbury

2 2 Other

Figure 1. Main producers of Pt and Pd in 2013, calculated from Cowley (2013).

Granophyres + Granites Rustenburg Layered Suite Rooiberg Group Pretoria Group Chuniseport Group

ZIMBABWE

Northern Limb

Younger Sedimentary Cover & Intrusions Bushveld Complex

Transvaal Supergroup

Natal

South Africa

Archaean Basement Pre-Transvaal Formations PGE deposits

f

Transvaal Transvaal Basin Basin

NAMIBIA

500 Miles

Potgietersrus Thabazimbi

Faults

MOZAMBIQUE

Fig. 1

f

f

Far Western Limb Western Limb

f

f

Eastern Limb

f

f f

Pretoria

26°S

Johannesburg 26°

27°

28°

0

25

50

75

100

km 29°

Bethal

30°

Figure 2. Geology of the Bushveld Complex showing the location of the 3 main deposits, Platreef, Merensky reef, and UG2 reef. (Modified after Barnes and Maier 2002a). S.J. Barnes, 01-fig_1

In addition to the Bushveld Complex there are three intrusions that are primary producers of Pt and Pd: the Great Dyke of Zimbabwe, the Stillwater Complex of the United States, and the Lac des Iles Complex of Canada (Figs. 4–6). The ore deposits in these intrusions are the Main Sulphide Zone, the JM reef, and the Roby and Offset Zones respectively (Table 2). In the past, the Sudbury Igneous Complex, Canada, (Fig. 7) produced a significant amount of Pt and Pd as a by-product of Ni mining (Table 2). The Sudbury Igneous Complex (Fig. 7) is a unique structure formed as a result of a meteorite impact, which flash melted a mixture of Archean and early Proterozoic crust (Dietz 1964). However, current production is low (Fig. 1) and these resources are now largely mined out. Nonetheless, an example of the Sudbury ores will be discussed because in terms of resources it originally represented a major source of PGE (Table 2).

Magmatic Ore Deposits

729

0 Map area

40

Cover Great Dyke Granites Greenstones Archean Gneisses

Talnakh 18O

Pfunz

i Belt

lt

Kharaelakh

16O

Be

Russia

Ma go na

Moscow

Fau lt

Noril'sk I

laekh

20O

- Kh

ara

Approximate outline of the Zimbabwe craton 100 km

50 km

22O 26O

Flood basalts (Triassic-Permian) Clastics (Permian-Carboniferous) Carbonates + Evaporites (Paleozoic)

East Boulder Mine Me

tam

orp

hic

roc

ks

0 2 4 6 8 km

J-M Reef Paleozoic and Mesozoic Sedimentary rocks

Stillwater Mine

Granite

6

5

4

3

2 1

Upper Banded

Stillwater Complex

(b) km series

Middle Banded

CANADA USA

Lower Banded

MONTANA

30O

32O

Figure 4. Simplified geology of Zimbabwe showing the location of the Great Dyke. Modified after Prendergast and Wilson (1989).

Figure 3. Geology of the Noril’sk-Talnakh area showing the location of the Noril’sk 1, Talnakh and Kharaelakh intrusions. Modified after Zientek et al. (1994).

(a)

28O

Basal Ultramafic

Nor il'sk

N

elt

oB

pop

Lim

zones Gabbronorite III Olivine-bearing V Anorthosite II

Picket Pin

Olivine-bearing IV Olivine-bearing III Anorthosite I Olivine-bearing II Gabbronorite II Norite II Olivine-bearing I Gabbronorite I Norite I

J-M Reef

Bronzitite Peridotite Basal Bronzitite Basal Norite

0 Figure 5. Geology of the Stillwater Complex showing the position of the JM reef. Modified after Godel and Barnes (2008).

Piña, R., Gervilla, F., Barnes, S. J., Ortega, L., & Lunar, R. (2014)

Barnes & Ripley

730

Canada

N

Yuk.

BC

North Lac des Iles intrusion

Mine Block intrusion

NT

Nanuvut

Alb.

NL Sas.

Man.

Quebec Ontario

P.E.I. NB

NS

Ottawa

Lac des Iles

East gabbro

High-grade Zone Camp Lake intrusion

1 km

Roby Zone

Diabase

500 m

Offset Zone

Felsic intrusive

Twilight Zone

Magnetite gabbro Leucogabbro

Hornblende Gabbro

Heterolitic gabbro breccia

Sheriff Zone

Pyroxenite Websterite

Variatextured gabbro Gabbronorite breccia Gabbronorite

Figure 6. Geology of the Lac des Iles Complex modified after Djon and Barnes (2012 ).

Ontario Quebec

Onaping-Levack Embayment Broken Hammer

U.S.A.

nge

Coleman

350 km

a th R Nor McCreedy East

Strathcona McCreedy West

46o30’

SRSZ

Fraser

Craig

Z

SRS

Sou

Gertrude

Totten

ge

an th R

Creighton

Whitewater Group Sedimentary rocks Onaping Formation Fallback Breccia Sudbury Igneous Complex Granophyre Norite Sublayer Norite Diorite ‘offset’ dykes Ni-Cu-PGE deposit

Garson

Palaeoproterozoic Metavolcanic and Metasedimentary rocks

Frood Sudbury Creighton Fault Copper Cliff South Granite plutons Kelly Lake N Archean and Proterozoic 81o00’

10 km

Gneiss and granite Fault

Figure 7. Geology of Sudbury Igneous Complex showing location of the McCreedy East deposit, modified after Dare et al. (2014).

Magmatic Ore Deposits

731

CLASSIFICATION OF THE DEPOSITS Broadly speaking the deposits may be classified based on their location in an intrusion and the amount of base metal sulfide minerals (BMS) present. There are three broad groups; stratiform or reef deposits, contact deposits and Ni-sulfide deposits.

Reef Fig. 8or stratiform deposits Bushveld 7 km

W&E limb

N limb

Stillwater

Plat Great

S MR UG2

+ reef Dyke ++ ++ ++ N

MSZ

JM

0 km unconformity

orthopyroxenite

diorite, gabbronorite, magnetitite gabbronorite, anorthosite, troctolite

orthopyroxenite, chromitite, harzburgite, dunite

norite, orthopyroxenite, chromitite

PGE reef

marginal zone norite, pyroxentite dolomite + ++ granite + ++

Figure 8. Stratigraphic columns of Bushveld, Great Dyke and Stillwater showing the positions of the reefs. MR = Merensky, MSZ = Main sulphide zone, modified from Figures of Prendergast and Wilson (1989), Barnes and Maier (2002a), Zientek et al. (2002).

Most primary PGE-deposits (the UG2, the Merensky, the JM reefs and the Main Sulphide Zone) take the form of laterally extensive narrow (stratiform) layers, of 1–3 m thickness, that contain 3–15 ppm Pt + Pd and occur within the Bushveld, Stillwater and Great Dyke layered intrusions respectively (Fig. 8). (Workers on economic deposits express the grade of a deposit in g.tonne−1 or as ppm: 1 g.tonne−1 = 1 ppm = 1 mg.g−1, this chapter will use ppm). In some places the PGE enriched layers widen to 10–20 m in structures which are termed pot-holes at the Bushveld reef (Viljoen 1999) and ballrooms in the JM reef (Zientek et al. 2002). With the notable exception of the UG2 reef, the reefs contain a small amount (0.5–3 wt%) of BMS. Base metal sulfides are present in the UG2 reef, but the amount is so low (< 0.1 wt%) that there are no visible BMS in hand specimen. The reefs are present in a variety of rock types. In the case of the UG2 the host is a massive chromitite with interstitial orthopyroxene and plagioclase (Barnes and Maier 2002a; Mathez and Mey 2005). The rock types in the Merensky reef vary from thin chromitite seams, through coarse-grained melanorite to anorthosite (Kruger and Marsh 1985; Barnes and Maier 2002b). The JM reef comprises troctolite, olivine gabbronorite and anorthosite (Barnes and Naldrett 1985; Zientek et al. 2002; Godel and Barnes 2008). The host rock of the Main Sulphide Zone is described as a plagioclase orthopyroxenite or bronzitite (Oberthür 2002; Wilson and Brown 2005).

Contact deposits The term contact deposit has been used to refer to deposits of variable width that are composed of disseminated BMS found at the margins of intrusions. The Platreef, which occurs along the northern edge of the Bushveld Complex (Figs. 2 and 8), falls into this group. Strictly speaking, the deposit should not be referred to as a reef because it is not a narrow zone, but rather 50–100 m thick with variable distribution of PGE across the zone. There is a great deal of variation in the rock types from pyroxenite to gabbronorite with minor anorthosite and peridotites, and included in these are xenoliths of the country rocks (Kinnaird 2005; Maier et al. 2008; McDonald and Holwell 2011). The combination of extreme variations in grain size and textures, the mixture of rock types, and presence of xenoliths with reaction rims around them, produced a heterogeneous zone referred to as varitextured.

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Eckstrand (2005) suggested that the Roby Zone of the Lac des Iles Complex should be classified with the Platreef. The Roby, Offset, and Twilight zones consist of varitextured gabbronorite containing disseminated BMS and they occur at the contact between the margins of the intrusion and the homogeneous east gabbro (Fig. 6). Other authors classify the Lac des Iles deposits as a magmatic breccia; e.g., Lavigne and Michaud (2001).

Ni-sulfide deposits In the Noril’sk area the ores take a number of forms: disseminated BMS, massive BMS, vein/ breccia/cupriferous BMS, and low-S–high-Pd–Pt ore (Distler 1994; Torgashin 1994; Sluzhenikin et al. 2014). Each type will now be briefly described. The disseminated ore is the main ore type of the Noril’sk 1 intrusion (Sluzhenikin et al. 2014). It occurs in the lower parts of the intrusion (Fig. 9a) in a varitextured gabbronorite. Russian geologists use the term taxitic for this type of texture. The BMS occur both as interstitial amorphous patches and as globules (referred to as droplet ore). The droplets are 1–4 cm in size and zoned with pyrrhotite-rich bases and chalcopyriteor cubanite (CuFe2S3)-rich tops (Czamanske et al. 1992; Barnes et al. 2006) (Fig. 10). The main ore type in the Talnakh and Kharaelakh intrusions is massive BMS, which occurs at the lower contact between the intrusions and the country rock (Fig. 9b). In many cases, there is a narrow zone of hornfels between the massive BMS and the intrusion. The massive BMS show a mineralogical zonation similar to that observed in the droplet ore (Fig. 9b). The lower parts and the margins are rich in pyrrhotite and the top or central parts are rich in a number of Cu-rich sulfide minerals from chalcopyrite to talnakhite [Cu9(Fe, Ni)8S16] (Torgashin 1994). Fig. 9 Cross-section of Noril'sk 1 intrusion

W a)

E

40m

mafic tuff/basalt

clastic sediments

disseminated sulfide in varitextured ol gabbronorite

disseminated sulfide in tuff/basalt

ol gabbronorite & gabbronorite

low-S high-Pd-Pt

varitextured leucogabbronorite & hornfels

Oktryabr'syk N-S Section

S b)

50 m

basalt

N

200 m Hornfels/ anhydrite

Hornfels/ argillite

Gabbro norite+/-ol

Cu-poor massive sulfide

Cu-rich massive sulfide

breccia/vein sulfide

Figure 9. Cross sections of the a) Noril’sk 1 and b) Oktrabr’sky deposits redrawn and simplified from Figures in Torgashin (1994) and Sluzhenikin et al. (2014).

Magmatic Ore Deposits Photomicrograph Pn Cb

2 mm

Ni X-ray map

Cu

Po X-ray map

Re 185 Laser map Cb

Fe

733

X-ray map

Os 189

Ir 193

Ru 101

Pd 108

Te 128

Cd 111

Pn

Po Rh 103

Pd to come Cu Figure 10. Zoned sulfide droplet from Noril’sk 1 (this work). Distribution of Fe, Cu, Ni determined by micro-XRF. Distribution of PGE determined by laser ablation ICP-MS. 108Pd has been corrected for Cd and Zn interferences. 103Rh has been corrected for Cu interference. 101 Ru has been corrected to Ni interference. Cb = Cubanite, Cp = Chalcopyrite. Pn = Pentlandite, Po = pyrrhotite.

Vein/breccia/cupriferous ore occurs in the country rock both above and below the intrusions and it also occurs cross-cutting the intrusions (Fig. 9b, Torgashin 1994; Sluzhenikin et al. 2014). The vein ore consists mainly of Cu-rich BMS and takes two forms. In some cases it Fig 10 consists of veins 2–10 cm wide containing massive BMS with sharp contacts with the host rock. In other cases, such as at the Oktrabr’ysk deposit (Fig. 9b), the vein ore occurs towards the top of the intrusion and in the overlying country rocks as a network of fine anastomosing veins. This ore is referred to as breccia or cupriferous ore. The low-S–high-Pd–Pt ore is not common. It occurs as narrow discontinuous layers towards the top of some intrusions (Distler 1994; Sluzhenikin 2011; Sluzhenikin et al. 2014). The host rock type is gabbronorite with disseminated chromite (Fig. 9a). Generally, the rocks contain ~ 1 wt% BMS. Traditionally, the Sudbury ores have been divided into contact ores and offset ores (Ames et al. 2007; Farrow and Lightfoot 2002). The contact ores consist of massive BMS and disseminated BMS found at the contact between the intrusion and the country rocks and include the chalcopyrite-rich massive BMS veins which occur in fractures immediately beneath the contact ores. The offset ores are found in quartz diorite dikes (Grant and Bite 1984; Lightfoot et al. 1997). These dikes are 50–100 m wide and extend for several kilometers into the country rocks. The dikes are thought to represent injections of the differentiated impact melt into the country rocks. Massive BMS occur both at the center and on the margins of the dikes, and occur where the dikes widen. More recently a new category of deposits has been described in the country rocks; these are low-S–high-Pd–Pt deposits (Pentek et al. 2008; Tuba et al. 2014).

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MINERALS HOSTING THE PLATINUM-GROUP ELEMENTS Base metal sulfides The highly chalcophile nature of the PGE suggests that they should partition into a sulfide liquid and therefore could be present in the BMS (pyrrhotite, pentlandite, chalcopyrite, pyrite) formed from this liquid. It has long been established that Pd, Rh, and Ru can be present in pentlandite (Cabri 1992 and references therein). However, until the advent of laser ablation inductively coupled plasma mass spectrometry (LA ICP-MS) analysis it was difficult to obtain in situ concentrations of all the PGE in BMS. As discussed below, there is now a rapidly growing literature on PGE contents of the BMS. We have chosen to represent examples where both whole rock and mineral analyses are available from the same samples. This section is not intended as a complete inventory of all the analyses available. In order to facilitate comparison of the mineral data with the whole rock data, the mineral data will be presented in the same fashion as the whole-rock PGE data, namely normalized to primitive mantle and plotted in order of compatibility. More details on the reason for this choice of this are outlined below in the section on whole-rock geochemistry. Pentlandite is the BMS with the highest PGE and Re content (Table 3). All of the pentlandites show similar patterns with an increase in concentration from Re to Pd and strong, negative Pt anomalies (Fig. 11). Pentlandites from the JM reef are richest in all of the PGE. The other reefs and Noril’sk pentlandites contain an order of magnitude less PGE. Pentlandites from McCreedy East and the Roby Zone contain 3 orders of magnitude less Os, Ir, Ru, and Rh than the JM reef pentlandites (Fig. 11, Table 3). The whole-rock concentrations of these elements are low at McCreedy and Roby Zone (Table 4). The magmas that formed the Sudbury deposits and Roby Zone are thought to have been andesitic in composition and thus contained less Os, Ir, Ru, and Rh than the picritic to basaltic magmas that formed the other deposits. Thus there are less of these elements in the sulfides. Pyrrhotite is the next most important host of PGE and Re. Pyrrhotites from the reefs and Noril’sk 1 have higher concentrations of all the PGE than those from McCreedy East or Roby Zone and, once again, reflect the difference in whole-rock concentrations (Table 4, Fig. 12). Rhenium, Os, Ir, Ru, and Rh are present at approximately Fig. 11 7

10

6

10

10

Pentlandite

10

5

10

10

10

10

10

10

10

10

10

10

10

10

4 3

mineral/mantle

2 1 0

10 10 10 10 10 10 10

5 4

Ni Re Os Ir Ru Rh Pt Pd Au Cu

6 5 4

Chalcopyrite

3 2 1 0 -1 6

Ni Re Os Ir Ru Rh Pt Pd Au Cu

10

Pyrrhotite

5

10

Pyrite

4

3

10

2

10

1

10

3 2 1

10

0

0

10

-1

-1

Ni Re Os Ir Ru Rh Pt Pd Au Cu

Merensky Plat reef

JM

10

Main Sulfide Zone altered Py Roby

Ni Re Os Ir Ru Rh Pt Pd Au Cu Noril'sk 1

Roby

McCreedy

altered Py McCreedy

Figure 11. Mantle-normalized concentrations in pentlandite, pyrrhotite, chalcopyrite and pyrite data sources listed in Table 3.

Magmatic Ore Deposits

735

Table 3. Metal Concentrations in the Base Metal Sulfide Minerals Deposit

Ni

Cu

Re

Os

Ir

Ru

Rh

Pt

Pd

Au

%

%

ppm

ppm

ppm

ppm

ppm

ppm

ppm

ppm

Ref

Pentlandite Merensky

32.41

0.11

0.27

3.11

Platreef

34.40

0.02

4.66

10.80

38.03

16.25

223

0.05

1

0.59

1.20

5.16

15.00

0.45

119

0.01

2

MSZ

34.00

0.00

0.07

0.44

1.09

JM

29.56

0.25

0.11

8.69

10.31

4.39

5.72

0.05

24

0.02

3

33.64

134.2

1.54

8782

0.01

Roby & Twilight

4

34.28

0.20

0.10

0.08

0.03

0.12

0.29

0.25

522

0.00

5

Noril'sk 1

31.20

0.05

0.27

0.82

McCreedy

34.32

0.07

0.05

0.02

2.52

8.50

28.55

11.88

296

0.50

6

0.02

0.25

0.04

0.01

6

0.01

7

Merensky

0.30

0.02

0.23

3.82

Platreef

0.51

0.00

0.93

3.34

5.21

1.13

0.98

1.33

0.06

1

1.80

7.81

0.45

0.61

0.06

0.01

2

Pyrrhotite

MSZ

0.28

0.00

0.13

0.39

1.22

4.14

2.25

0.86

1.24

0.01

3

JM

0.29

0.06

0.04

0.21

0.72

1.32

0.36

0.03

24.04

0.01

4

Roby & Twilight

0.82

0.14

0.12

0.11

0.03

0.23

0.45

0.28

0.78

0.01

5

Noril'sk1

0.14

0.00

0.40

1.26

4.08

11.30

46.91

10.24

0.36

0.06

6

McCreedy

0.66

0.00

0.06

0.02

0.02

0.04

0.02

0.01

0.02

0.00

7

0.21

0.70

6.51

0.04

1

1.31

0.54

0.01

2

1.74

1.04

0.01

3

Chalcopyrite/Cubanite Merensky

0.19

33.84

0.06

0.30

0.30

3,000. The mechanism generally invoked to attain a high R-factor is to argue for an extremely dynamic system. In the case of Noril’sk 1, Talnakh and Kharaelakh, the intrusions that contained the BMS can be thought of as conduits through which many pulses of magma may have been transported. In addition the sulfide droplets themselves may have been formed in a lower chamber and been transported; during transport they could have collected some metals from the transporting magma. They could also have been kept in suspension after they were emplaced and collected some metals from the various pulses of magma that passed through the system. Rice and Moore (2001) carried out finite element modelling to demonstrate how sulfide droplets can be held in suspension above an embayment in the footwall of an intrusion. In the case of the Platreef it is argued that sulfide saturation took place at depth and the BMS liquid were entrained in the magma and emplaced at the margins of the Bushveld Complex (McDonald and Holwell 2007). During transport, BMS could have interacted with large volumes of magma. Some authors (Maier et al. 2013 and references therein) suggest that the Platreef changes character down dip and into the intrusion and merges with the Merensky reef away from the margins of the intrusion. We suggest that possibly the magma with entrained sulfide droplets froze rapidly to form the Platreef at the margins of the intrusion; however, further into the intrusion the incoming magma with entrained droplets had more time to interact with the resident magma and thus the sulfide droplets had the opportunity to collect more PGE, and thus the Merensky reef has a higher PGE grade than the Platreef.

752

Barnes & Ripley

The UG2 and JM reef both present a challenge to a simple sulfide collection model. Both require extreme R-factors in the 105–106 range (Fig. 18), which one of the authors (ER) regards as unreasonable. One possible solution is that the sulfide droplets come into contact with sulfide-undersaturated magma and party dissolve (Kerr and Leitch 2005). The sulfides could, for example, form at the top of intrusion and settle downwards into sulfide-undersaturated magma (Holwell and Keays 2014). Due to their very high partition coefficients into sulfide liquid, the PGE would be preferentially retained in the sulfide liquid. Other processes that could have affected the JM and UG2 reefs are considered in the crystallization of sulfide liquids and sub solidus sections below.

Crystallization of a sulfide liquid Experimental work has shown that the first phase to crystallize from a sulfide liquid is an Fe-rich monosulfide solid solution (MSS) and the remaining sulfide liquid becomes enriched in Cu (Kullerud et al. 1969; Dutrizac 1976; Ebel and Naldrett 1997; Sinyakova and Kosyakov 2009). When the temperature decreases sufficiently (< 900 oC), the Cu-rich liquid crystallizes as intermediate solid solution (ISS). The partition coefficient of Ni between MSS and sulfide liquid is dependent on temperature and fS2 of the system (Li et al. 1996; Barnes et al. 1997a; Makovicky 2002; Liu and Brenan 2015). At fS2 reflecting crustal conditions, the partition coefficient increases from ~ 0.2 at 1100 oC to ~ 1 at 900 oC. Thus, the evolved MSS is richer in Ni than the first MSS. Natural examples of the crystallization history of sulfide liquids can be observed in sulfide droplets from MORB. Some have chilled sufficiently rapidly for a quench texture of the BMS to be preserved showing an intergrowth of MSS and ISS, with a Ni-rich phase between the MSS and ISS (Fig. 20a). If the crystallization of the liquid proceeds more slowly, then the Curich liquid separates from MSS in some cases. Indeed, both in MORB droplets and at Noril’sk 1 there are sulfide droplets with Cu-rich tops and Cu-poor bottoms preserving the crystallization history of the sulfide liquid (Figs. 10 and 20b). In the case of some of the massive ores, after initial crystallization of Fe-rich MSS the Cu-rich fractionated liquid migrates away either into the footwall or into the overlying rocks as seen in the veins surrounding both the Noril’sk 1 intrusion and the Sudbury Igneous Complex. This fractionated liquid crystallizes as an ISS cumulate and can be seen in some MORB droplets (Fig. 20c) Rhenium, Os, Ir, Ru, and Rh substitute into MSS, whereas most other elements do not (Table 1). Thus, the MSS cumulates (Cu-poor massive sulfides) are enriched in IPGE and Re relative to the original sulfide liquid, and have flatter mantle-normalized metal patterns than the disseminated sulfides (Fig. 12b). These elements are also concentrated in the lower parts of the Noril’sk sulfide droplets (Fig. 10). Most of the elements aside from Cu, Cd, and Zn are also incompatible with the ISS (Table 1) therefore Pd, Pt, and TABS partition into the last sulfide liquid. Because the major elements have all been consumed there is very little of the late liquid and it is generally trapped in the ISS cumulate (Cu-rich ore). Consequently, the Cu-rich ores are normally enriched in Pt, Pd, and TABS, and have mantle-normalized metal patterns that are much steeper than the disseminated ore (Figs. 12b,c and Figs. 15c,d). The fractionated interstitial liquid will be enriched in Pt, Pd, and TABS. It is possible that the liquid becomes saturated in Pd- and Pt-bismuthotellurides, antinomides and stannides. Composite, rounded grains of mixtures of bismuthotellurides and stannides are found interstitial to chalcopyrite in the Cu-rich ore at McCreedy East (Figs. 21 a,b). The composite nature of these grains suggests that they cannot be exsolutions and they have been interpreted as the product of crystallization from the late liquid (Dare et al. 2014). In some cases the late liquid escapes the ISS cumulate, producing a Cu-rich ore depleted in Pd, Pt, and TABS, and a separate Pd–Pt–TABS-rich liquid. This Pd–Pt–TABS-rich liquid could potentially migrate into the rocks around the sulfide deposits and form low-S–highPd–Pt deposits such as those found around the Sudbury intrusion (Pentek et al. 2008; Tuba

Magmatic Ore Deposits

753

a)

b)

c)

Figure 20. Photomicrographs of sulfide droplets found in MORB glass (modified after Patten et al. 2012) a) Droplet that crystallized rapidly showing MSS and ISS intergrowth with Ni-rich phase between the two phases. b) A droplet that crystallized slightly more slowly resulting in solid MSS at the base and quenched textured ISS in the top left hand corner. c) A droplet that crystallized slowly with MSS at the margins and solid ISS in the center.

Fig. 20

et al. 2014) and in the upper parts of the Noril’sk intrusions (Sluzhenikin 2011). The Pd–Ptbismuthotellurides which could crystallize from these liquids solidify at fairly low temperatures (< 600 oC, Elliott 1965; Hoffman and MacLean 1976; Savitsky et al. 1978; Moffatt 1979a,b). Therefore, when there is a structural adjustment of a cooling cumulate (or possibly even during later metamorphism) the bismuthotellurides could be readily mobilized and move into

Barnes & Ripley

754 Cp

PdBi Te Ag2Te

a) Pd2Sn

PtSnTe PtTe2

PdBi2

BiTe

Cu-FeSn-S

Ag2Te Cp

PbS

30 m

40 m c)

Pn CuRhPtS

50 m

PtAs2

Silicates

d) PdTe

PdAs PdAs

PdTe

e Po

70 m

Cp

Pl

Po

Amph

b)

PbS

PdBiTe (michenerite)

50 m

Figure 21. Back scatter electron images of PGM: a) and b) PGM that have crystallized from the fractionated interstitial liquid (modified from (Dare et al. 2014); c) CuPtRhS exsolution in pyrrhotite and pentlandite from the Merensky reef (modified from Prichard et al. 2004); c) Residual PGM after the replacement of pentlandite, Lac des Iles (modified from Djon and Barnes 2012); e) Remobilized Pdbismuthotellluride, (modified from Dare et al. 2010).

Fig. 21

fractures and faults and may occur as isolated grains as seen in the Creighton deposit, Sudbury (Dare et al. 2010; Fig. 21e). Possibly an extremely fractionated liquid also formed the High Grade zone at Lac des Iles (Djon and Barnes 2012). One could even suggest that addition of Pd–Pt–TABS liquid to a pre-existing disseminated sulfide layer formed the JM reef, thus accounting for the extremely low Cu/Pd ratio observed in the JM reef, by addition of Pd.

Late magmatic fluids A number of authors (Boudreau et al. 1986, 2014; Pentek et al. 2008; Tuba et al. 2013) favor a role for late magmatic fluids in the formation of PGE-reefs, the Lac des Iles deposits and the low-S–high-Pd–Pt deposits. In this model late magmatic fluid partly dissolves magmatic BMS which have already crystallized. Palladium and, to a lesser extent, Pt dissolve into the fluid. The fluid migrates away from the residual BMS and, when the physico-chemical conditions change, the fluid deposits the Pd and Pd. In some cases the deposition is in stratiform layers within the intrusion: these are reefs. In some cases the Pd and Pt are transported out of the intrusion and may form low-S–high-Pd–Pt zones surrounding the intrusions. Boudreau and Meurer (1999) Hanley et al. (2008) Godel and Barnes, (2008) have suggested that Pd and, to a lesser extent, Pt have been added to the JM reef by late fluids. Pentek et al. (2008) Tuba et al. (2013) argue that the low-S–high-Pd–Pt zones formed around the Sudbury deposits formed in this way. Fluids derived from late cooling or metamorphism also result in a change in the BMS present. At Lac des Iles, pyrrhotite and pentlandite are replaced with pyrite and millerite (Djon

Magmatic Ore Deposits

755

and Barnes 2012). The pyrite inherits the PGE that were present in the pyrrhotite and thus contains Re, IPGE, and Rh (Fig. 11). Millerite contains some IPGE but does not accept the Pd; thus, Pd combines with Te, As, and Sb to form isolated PGM grains (Fig. 21d). Similar replacements have been observed in Grasvalley prospect in the Bushveld Complex (Smith et al. 2014). In the case of the Platreef, the PGM assemblage in the altered ores is richer in arsenides and antinomides (Hutchinson and Kinnaird 2005).

Subsolidus events At temperatures < 600 oC, MSS exsolves to form pyrrhotite and pentlandite ± pyrite, and ISS exsolves to form chalcopyrite ± cubanite ± pyrite. The sulfide minerals that exsolve from MSS (pyrrhotite and pentlandite) inherit the IPGE and Re which were in the MSS (Fig. 11), and thus the mass balance indicates that these elements are concentrated in pyrrhotite and pentlandite (Fig. 13). All of the PGE, Re, and Au are incompatible with ISS (Table 1), and thus the chalcopyrite and cubanite that exsolve from the ISS are poor in highly siderophile elements (Fig. 11). Significant amounts of Pd are found in pentlandite, which at first glance is surprising because the partition coefficient of Pd into MSS is 0.1–0.2 (Table 1) and thus one would not expect pentlandite which is derived from MSS to contain Pd. Nonetheless, some Pd will partition into MSS. Dare et al. (2010) postulated that during exsolution of MSS to pentlandite and pyrrhotite the Pd diffuses along with the Ni into pentlandite. Dare et al. (2010) and Piña et al. (2012) showed large granular pentlandites, which form at high temperatures, contain more Pd than small flame pentlandites. They interpret their observation to be the result of the larger grains, which formed at higher temperature, as having depleting the MSS in Pd before the flame pentlandite formed. The pentlandite hosts much of the Pd in massive ores because these are MSS cumulate and thus the contribution of Pd from the Cu-liquid is limited to the trapped liquid fraction. Consider for example the case where the original sulfide liquid contains 10 ppm Pd, then the MSS should contain 1 ppm. A MSS cumulate with 90% MSS will then contain 2 ppm Pd. One ppm of this in the MSS and 1 ppm in the trapped liquid fraction. The Pd in the MSS will diffuse into pentlandite during exsolution of the MSS, and thus 50% of the Pd will be present in pentlandite. The diffusion of Pd into pentlandite during exsolution could explain the high-Pd content of the pentlandites formed in MSS cumulates, but it will not explain the tendency for almost 100% of Pd to be present in pentlandite in BMS thought to represent the sulfide liquid compositions such as the Noril’sk 1 sulfide droplets (Fig. 13). In this case the pentlandite should contain only 10–20% of the Pd. It is possible that not all pentlandite forms by exsolution. Coarse-grained pentlandite occurs between the pyrrhotite and chalcopyrite or cubanite in the droplets (Fig. 10). Makovicky (2002) describes experiments reported by Distler et al. (1977) where the texture was reproduced in experiments by cooling a sulfide liquid. The “pentlandite” formed by peritectic reaction between the fractionated liquid and the MSS. They referred to the pentlandite as “high-temperature” pentlandite. If this is the origin of the coarse pentlandite in the droplets then Pd could have partitioned into the “high temperature” pentlandite when it formed from the fractionated liquid. Not all of the PGE are present in BMS (Fig. 13). During cooling and exsolution of the MSS, the solubility of the PGE in the sulfide minerals decreases due to the fall in temperature and the change in S content of the BMS (Makovicky et al. 1986; Li et al. 1996). Therefore, as the BMS cool, PGM exsolve forming elongate grains within the BMS. For example CuRhPtS exsolutions develop in pyrrhotite and pentlandite from the Merensky reef (Fig. 21c). The formation of PGM can be enhanced by the circulation of late magmatic fluids or metamorphic fluids which dissolves some S, resulting in a further fall in PGE solubility in the BMS, especially at the grain boundaries (Godel et al. 2007). Subsolidus processes have also been called upon to explain the low-S and -Cu contents of the UG2 reef (Naldrett 2011). Naldrett (2011) proposed that originally more BMS were present

Barnes & Ripley

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than are now observed and as the cumulates cooled there was a diffusion of Fe from the BMS into chromite which destabilized the sulfide minerals, and resulted in S and Cu loss from the assemblage and in exsolution of laurite from the BMS (Junge et al. 2014). The postcumulate loss of Cu from the UG2 reef could provide a more reasonable explanation for the low Cu/Pd than the very high R-factors suggested by a simple sulfide collection model. Prichard et al. (2004) also suggest S loss explains the PGM mineralogy in the chromite layers of the Merensky reef.

UTILIZATION OF THE Re–Os ISOTOPE SYSTEM IN STUDIES OF MAGMATIC Ni–Cu–PGE ORE GENESIS Background Rhenium and Os may both be concentrated in magmatic sulfide liquids, sulfide minerals, or minerals that are strongly associated with sulfides (e.g., arsenides, bismuthinites, tellurides), and hence the Re–Os isotope system has been extensively utilized in studies of sulfide ore genesis. Deposits studied include relatively low-temperature, black-shale-hosted, Ni and base metal deposits (e.g., Horan et al. 1994; Pasava et al. 2010), hydrothermal Au and related sulfide mineralization (e.g., Selby and Zhao 2012), higher-T porphyry Cu–Mo systems (e.g., Barra et al. 2003; Drobe et al. 2013), and magmatic Ni–Cu–PGE deposits (see reviews by Lambert et al. 1998a,b, 1999a). We will briefly review the application of the Re–Os system to magmatic Ni–Cu–PGE deposits, but stress that the system has applicability far beyond the high-T realm covered here. All aspects of the Re–Os system that are of importance for geochronology have been covered in other chapters of this book and will not be repeated here (Harvey and Day 2016, this volume; Harvey et al. 2016, this volume). The direct dating of molybdenite using the Re–Os system is reviewed by Stein et al. (2001). We will review geochronological applications for several deposits, as appropriate, below. Our emphasis will be on the application of Re–Os isotope measurements as indicators of the source of Os in magmatic ore deposits and the degree of crustal contamination that source magmas have undergone. To this end, it is important to review the partitioning behavior of Re and Os into sulfide and silicate melts, silicate and oxide minerals, and metallic alloys. One of the basic tenets of the Re–Os system is that during partial melting of the mantle, Re is strongly partitioned into the melt, whereas Os is retained in the mantle (e.g., Morgan et al. 1981; Hauri 2002). The high Re/Os ratios developed in the partial melts lead to the production of radiogenic Os in the crust, and much higher 187Os/188Os ratios than in the residual mantle. Mantle-derived magmas that then interact with crust during or after emplacement may be characterized by elevated 187Os/188Os ratios. Although the distribution of Re and Os appears to be well-constrained, the carriers of Re and Os in the mantle remain controversial. Within the mantle both Re and Os appear to be largely controlled by chalcophile tendencies. Studies by Mallmann and O’Neill (2007), Brenan et al. (2003), and Righter and Hauri (1998), indicate that Re may be partitioned into silicate and oxide minerals, but is much more strongly partitioned into sulfide. Experimental studies by Brenan (2008) and Fonseca et al. (2011) indicate that DOs(sulfide melt–silicate melt) is in the order of 104–105, and DRe(sulfide melt–silicate melt) for basaltic systems is ~ 400–800 (see also Brenan et al. 2016, this volume). If sulfide minerals in the mantle are exhausted during partial melting, the high Re/Os values of partial melts can only be explained if Os is retained in the mantle in residual phases such as Os–Ir alloys (e.g., Fonseca et al. 2011). If sulfide minerals are retained, then the results of Brenan (2008) suggest that Os may be considerably more compatible than Re in residual sulfide, and observed Re/Os ratios in basaltic partial melts can be produced. Additional studies are needed to better constrain the sites of Re and Os in the mantle, and particularly the partitioning of Os into residual phases such as alloys during high-degree partial melting of the mantle, where sulfide minerals are completely dissolved in the partial melt. Many studies of magmatic Ni–Cu–PGE deposits are

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consistent with Re-enrichment in the crust. Whether this is completely a function of melting where sulfide is retained in the mantle, the retention of Os in alloys, or local interaction between mantle magmas and crustal sedimentary rocks where Re may be strongly enriched over Os (e.g., organic–rich shales) must be determined on a deposit to deposit basis. Examples of the application of the Re–Os isotope system to several Ni–Cu–PGE deposits are presented below. However, we first review the basic tenets of how the system has been applied in the evaluation of contamination of mafic and ultramafic magmas. The gOs terminology has been defined in other chapters (Harvey et al. 2016, this volume), but, in brief, is the percentage deviation in 187Os/188Os from Bulk Silicate Earth (BSE), taken as either Primitive Mantle or Chondrite, at the time of interest: gOs = ((187Os/188Os(sample) − 187Os/188Os (PM or CH))/ 187Os/188Os(PM or CH) − 1) × 100 (see Shirey and Walker 1998; Carlson 2005). When isochronous behavior is demonstrated, the initial 187Os/188Os value is utilized in the gOs calculation. Gamma-Os values that show significant deviation from 0 typically results from either crustal contamination or hydrothermal processes involving the transport of Re, Os, or both. Derivation of magma from recycled crust in the mantle or metasomatically enriched lithosphere may also produce negative or positive deviation from 0, although the number of intrusions with gOs values near those of BSE suggests that processes occurring in the crust are of more significance in controlling gOs. An isochron diagram is useful for illustrating the effects of these processes. Figure 22 shows a mantle-derived magma with a low 187Re/188Os ratio (mafic magmas may have a wide range of Re/Os ratios but here we take a value less than 150 as being representative of basaltic magmas), Figure 22. 187Os/188Os vs. 187Re/188Os isochron diagram illustrating mixing at 1.1 Ga between an uncontaminated mantle magma and a Re-rich sedimentary rock that had evolved between 1.85–1.1 Ga. The contaminated magma (M) crystallizes at ~ 1.1 Ga and evolves to the position of the contaminated bulk rock, far above the 187 Os/188Os value that the uncontaminated rock would have attained. Minerals A and B may crystallize from the contaminated magma, and together with the bulk rock value define an isochron with an elevated gOs(i) value.

which interacts with a crustal sedimentary rock with an elevated 187Re/188Os ratio. We utilize a sedimentary contaminant with an age of 1.85 Ga, and mixing between the mantle magma and the sedimentary contaminant at 1.1 Ga. The choice of initial 187Os/188Os ratio is arbitrary, but we use a chondritic initial for both the mantle-derived magma and the sedimentary contaminant (e.g., mantle-derived detritus, or seawater controlled by leaching of chondritic oceanic crust). The isochron, which represents growth of the sedimentary rock from 1.85 Ga until 1.1 Ga, is shown for reference, as is the isochron for uncontaminated mantle. At 1.1 Ga, a mantle-derived magma is emplaced and is contaminated with Os derived from the sedimentary rock (point M). Crystallization of the contaminated magma and decay from the time of mixing (1.1 Ga) until the present will produce an isochron yielding the age of the mixing event with an elevated initial 187 Os/188Os, and an elevated gOs value. If multiple pulses of magma are involved, which may have interacted to different degrees with the sedimentary country rocks (different f values), then a knowledge of the age of the intrusion (e.g., from U–Pb isotope measurements) is required to assess the range of gOs values. The two-component isotopic mixing equation has the form:

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758  187 Os   188   Os mix,o

  Os   ( ) ( f ) + C (1- f )   Re  + (e − e )  ( )        Os  ( )( f )  Os     ( C ) (1- f )   Os    Re   + ( C ) (1- f )   + (e − e )      ( C )( f )  Os Os

 COss =   COss 

c

Os

187

187

188

188

s,o

c

+

Os



187

c

Os

Os

λT

 

188

187

 c,o 

λt 2

s,t 2

188

λt1

 c,t2

λt 2



where f = mass fraction of the mantle-derived source magma, COss and COsc = concentration of Os in the source magma and contaminant, respectively, subscripts (mix,o), (s,o) and (c,o) refer to the initial 187Os/188Os ratio of the mixture, source magma, and contaminant, subscripts (c, t2) and (s, t2) refer to the 187 Re/188Os ratios of contaminant and source magma at the time of mixing, T = 4.557 x 109, t1 = age of the contaminant, and t2 = time of mixing. We illustrate in Figure 23 that in addition to multiple magma pulses and variable degrees of contamination, a suite of non-linear 187Os/188 Os values may also indicate either Re loss/gain or Os gain/loss via lower-temperature interaction with fluids. Xiong and Wood (2001) have shown that Re may be soluble as a hydroxyl species in low-T fluids, and recent studies by Foustoukos et al. (2015) further quantify the extent to which Re and Os may be transported via hydrothermal fluids. Figure 23 illustrates a case of crystallization from 1.1 Ga until 0.5 Ga, when hydrothermal alteration changed the Re/Os ratio of the system. In the case of Re loss or Os gain, the resultant 187 Os/188 Os values will plot above the isochron for the uncontaminated system, and these points may be difficult to distinguish from those produced by variable degrees of crustal contamination of multiple pulses of magma. Sampling at small spatial intervals may be required to resolve the interpretation; evidence of high degrees of hydrothermal alteration would be expected. In the case of Re gain or Os loss, the resultant 187Os/188Os values will fall below the expected isochron for uncontaminated mantle and will produce negative gOs values. Scatter in 187Os/188 Os values must be carefully evaluated as it may not be indicative of contamination of mantle-derived magma, but perturbation of the Re–Os system by interaction between igneous rock and low-temperature fluids

Figure 23. Isochron diagram illustrating the potential effects of Re loss/gain and Os loss/gain on today’s isotope ratios. In this example a rock derived from a mantle-derived magma at 1.1 Ga evolves to 0.5 Ga and then the system is opened via interaction with a fluid. In the case of Re loss/Os gain, the 187Re/ 88Os values are reduced and the resultant values may evolve to 187Os/188Os values that lie above the expected isochron for unperturbed samples, and may be difficult to distinguish from samples with elevated gOs values due to crustal contamination. In the case of relative Re gain/Os loss the resultant values are below those produced from uncontaminated magma and are characterized by negative gOs values.

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The “R-factor” and its application to Re–Os isotopes The “R-Factor” (ratio of sulfide liquid to silicate magma) has been discussed above, and is often mentioned in conjunction with Re–Os isotope studies (e.g., Lambert et al. 1998a,b, 1999a). If S in the ore-forming system has been primarily derived from country rocks, and contamination has involved only S, then the R-factor directly relates to f values where f(silicate) = R/(1 + R). In the more general case, where a sulfide-undersaturated magma assimilates country rock, including S, utilization of the “R-factor” is appropriate after the separation of immiscible sulfide liquid. Although both Re and Os may be strongly partitioned into the sulfide liquid (see above), no isotopic fractionation accompanies this process. If Os is more strongly partitioned into the sulfide liquid than is Re, then the Re/Os ratio of the sulfide liquid may be lower than that of the coexisting silicate magma, but crystallization products of both liquids should lie along the same isochron provided closed system conditions have prevailed. This is an important point in the evaluation of massive sulfides that are suspected of separating from a parent liquid that is now represented by a silicate assemblage and contains only disseminated sulfides. If the systems share a common origin then they should fall along the same isochron.

Examples of the application of the Re–Os isotope system to magmatic ore deposits The Sudbury Igneous Complex. The Sudbury Igneous Complex (SIC) was the world’s largest producer of Ni, with significant production of PGE as by-products. Walker et al. (1991) presented results of one of the earliest applications of the Re–Os system to magmatic ore deposits, in this case the Ni-Cu ores of the SIC. The work by Walker et al. (1991) was followed by studies by Dickin et al. (1992) and Morgan et al. (2002) who obtained greater precision in measurements using negative thermal ionization mass spectrometry. All of these studies confirmed the presence of crustal Os in the ores. Samples collected from the McCreedy West and Falconbridge mines produced isochrons of 1835 ± 70 Ma and 1825 ± 340 Ma, respectively (Morgan et al. 2002; Fig. 24). The ages agreed within uncertainty with that of 1850 ± 1 Ma determined using U–Pb methods by Krogh et al. (1984). Gamma-Os values for the McCreedy West data averaged 346 ± 10 and that from Falconbridge was 375 ± 12. The 187Os/ 188Os values could be modelled as resulting from mixtures of Proterozoic and Archean country rocks, However, Morgan et al. (2002) also used 186Os/188Os ratios to indicate that in addition to metasedimentary rocks a third component was necessary to explain the Re–Os systematics of the Falconbridge and McCreedy West samples. The third component was thought to be mafic rocks of Archean or Proterozoic age. It is now thought that the Sudbury melt sheet included Ni–Cu- and PGE-bearing sulfides from Proterozoic mafic rocks in the target area (Keays and Lightfoot 2004). Of particular significance is the conclusion that the Re–Os isotope systematics of the Sudbury ores can only be explained if a significant portion of Os was derived from crustal metasedimentary rocks, in agreement with the origin of the SIC as an impact melt sheet.

Figure 24. Isochron diagram for samples from the McCreedy West mine, Sudbury (modified from Morgan et al. 2002).

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Ni–Cu–PGE deposits of the Noril’sk region, Siberia. Walker et al. (1994) and Malitch and Latypov (2011) have presented Re–Os isotopic data from the Noril’sk, Talnakh, and Kharaelakh ore-bearing intrusions. Data from both studies provide isochrons that are consistent with the U–Pb ages of the intrusions (~ 247 Ma) and gOs values that show slight differences between intrusions but range from only 0.4 to 12.9 (Figure 25). Both Walker et al. (1994) and Malitch and Latypov (2011) suggested that the Re–Os isotopic values may be indicative of their mantle source and in particular deep mantle that had incorporated ancient Re-enriched crust. Alternatively, the gOs values could be indicative of very minor amounts of crustal contamination in the lithosphere. The very low gOs values do not indicate the extensive interaction with crustal material that is suggested by the elevated sulfur isotope compositions of the ores (e.g., Grinenko 1966; Ripley et al. 2003; Malitch et al. 2014) and the anomalous Sr-isotope compositions reported by Arndt et al. (2003). This may reflect the derivation of S and Sr from evaporites, which would tend to have low concentrations of Re and Os. Alternatively, Lesher and Burnham (2001) suggested that due to mass balance constraints, Re–Os isotopic values may be “re-set” to near chondritic values and S isotope values little affected as uncontaminated mantle-derived magmas exchanged with sulfides in a conduit environment. A similar process has been proposed to explain the Re–Os isotopic systematics of the Eagle deposit (see below).

Figure 25. Isochron diagram for samples from the Noril’sk 1 and Talnakh intrusions (modified from Walker et al. 1994).

The Voisey’s Bay Ni–Cu–Co deposit, Labrador. Lambert at al. (1999b, 2000) have utilized the Re–Os isotope system in the interpretation of the genesis of the Voisey’s Bay Ni–Cu–Co deposit. The deposit is known as a “conduit-type” because of the clear occurrence of sulfide mineralization in dike-like bodies that connect larger intrusions (e.g., Naldrett and Li 2007; Lightfoot et al. 2012). Massive sulfide mineralization occurs in a widened portion of the dike that connects the western deeps intrusion and the upper Eastern deeps intrusion. Mining of the massive sulfide, or “Ovoid”, is currently in progress. Lambert et al. (1999b, 2000) identified two scales of Re–Os isotope development, and a multi-stage process of mineralization. Whole-rock samples of the massive sulfide assemblage remained closed over a large spatial interval, and an imprecise 1323 ± 135 Ma isochron was generated; this age was within error of the 1332.7 ±1 Ma badelleyite U–Pb age for the Voisey’s Bay intrusion of Amelin et al. (1999). Of particular significance is the initial 187Os/ 188Os value determined from the isochron plot (Figure 26). This initial ratio yielded a gOs value of 1040 ± 200, signifying extensive contamination with old crustal material. The Proterozoic Tasiuyak Gneiss was taken to be a likely source, as other petrogenetic indicators also suggested that assimilation of the Tasiuyak Gneiss had been extensive (e.g., Li and Naldrett 2000). Archean rocks of the Nain Province could not be entirely eliminated, and Lambert et al. (2000) suggested that a twostage process of magma contamination had occurred, with high-MgO basaltic magmas first contaminated in the lower crust. The results from the study of the whole-rock massive sulfides

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provide an excellent example of how the Re–Os system may provide both geochronologic data and information on magma–country-rock interaction. In this case the Re–Os system of the sulfides was homogenized prior to crystallization of the sulfide assemblage. Lambert et al. (2000) also showed that disseminated sulfide-bearing intrusive rocks showed a great deal of scatter on an isochron plot, and suggested that variable degrees of mixing with country rocks were responsible. In addition, sulfide mineral separates from the massive sulfides defined an isochron with an age of 1004 ± 20 Ma, consistent with re-setting of the sulfide system on the mineral scale as a result of hydrothermal alteration associated with the Grenville orogeny.

Figure 26. Isochron diagram for samples from the Voisey’s Bay deposit (modified from Lambert et al. 1999b).

The Eagle Deposit, Midcontinent Rift System, Michigan. The Eagle deposit is a newly discovered Ni–Cu–PGE sulfide deposit in mafic rocks associated with the 1.1 Ga Midcontinent Rift System of Minnesota, Michigan and Canada. Mineralization occurs within irregularities in a dike-like body, which is part of the Marquette-Baraga Dike Swarm (Ding et al. 2010). It bears similarities to the Voisey’s Bay deposit in terms of its conduit environment. Ding et al. (2012) present Re–Os isotopic data from disseminated, net-textured, and massive sulfides. Their results produced an isochron of 1106 ± 34 Ma (Fig. 27), in good agreement with the U–Pb age of the intrusion of 1107 ± 2 Ma reported by Ding et al. (2010). The gOs value corresponding to the initial 187Os/188Os ratio of 0.1607 was 34, and could be explained by less than 3% bulk contamination of a mantle-derived picritic magma with Proterozoic country rocks. The apparent low degree of contamination

Figure 27. Isochron diagram for samples from the Eagle deposit (modified from Ding et al. 2012).

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is contrary to models for ore genesis that favor extensive crustal contamination. Ripley and Li (2003) and Ding et al. (2012) showed how exchange reactions in magma conduits involving early-formed sulfide and uncontaminated magmas could produce isotopic values approaching those of the uncontaminated magma and mask evidence for earlier high degrees of contamination (Fig. 28).

Figure 28. Plot showing the decrease in Os and S isotope ratios due to exchange between sulfide in a conduit system and batches of uncontaminated, mantle derived magma. The equation of Ripley and Li (2003) was used to calculate the trends: d34S sul,exchanged, = (d34S sul,initial + R*( d34S sil, initial + ∆))/(1 + R*). For Os the 187 Os/188Os ratio is used rather than the d34S value. In this example the initial 187Os/188Os ratio of the sulfide was set at 1, the 187Os/188Os ratio of the mantle-derived magma was 0.15, the concentration of Os in the silicate magma was set at 1 ppb, the d34S values of the magma and sulfide were 0 and 12 ‰, and the concentrations of S in the sulfide and magma were 30 wt.% and 1000 ppm, respectively. R* is (Csil/Csulf ) × R, where concentrations are S or Os and R is the silicate liquid/sulfide liquid mass ratio. N is the integrated silicate/sulfide mass ratio where in this example each pass of uncontaminated magma interacted with an equal mass of sulfide (R = 1). ∆ is the isotopic difference between S or Os in the sulfide mass and that in the silicate melt (at magmatic temperatures this value is ~ 0). If the R value for each increment of exchange is larger, then the rate of S or Os isotope decline is greater.

The Stillwater Complex, Montana. As described above, the 2700 Ma Stillwater Complex hosts one of the world’s major PGE reef-style deposits. The JM Reef and host rocks, as well as PGE-bearing chromitites in the Ultramafic Series, have been studied using the Re–Os isotope system by a number of researchers (e.g., Lambert et al. 1989, 1994; Martin 1989; Marcantonio et al. 1993; Horan et al 2001). Lambert et al. (1989, 1994), Martin (1989), and Horan et al. (2001) have all concluded that two sources of Os were involved in the generation of the Complex. Lambert et al. (1994) reported that PGE-enriched units of the Complex (B chromitite seam and the J-M Reef) were isochronous and characterized by elevated gOs values of 12–34. PGEpoor chromitites (G and H) were found to be near-chondritic with gOs values of −2 to +4. The Re–Os data were consistent with other petrogenetic data that indicated that two magma types were involved—one an ultramafic magma with little sign of crustal contamination but possible derivation from the subcontinental lithospheric mantle, and the other a more mafic to anorthositic magma that had been contaminated with Archean crust prior to emplacement. The mixing of these two magma types was proposed to account for the elevated gOs values and was thought to be instrumental in the formation of the JM Reef. Horan et al (2001) separated chromite from chromite seams and silicate–rich layers of the Ultramafic Series. They found variable gOs values of 2–16.4, and also suggested that two magmas were involved in the formation of the Ultramafic Series, one which was near-chrondritic and the other which was characterized by elevated gOs values and had been contaminated by sedimentary rocks beneath the Complex. Martin (1989) concluded that a chondritic mantle melt was contaminated with mafic-to-intermediate crust at the time of the Ultramafic-Banded Series transition and at the time of JM Reef formation.

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Marcantonio et al. (1993) indicated that chromitites from the Ultramafic Series were derived from a mantle magma that had been affected by little or no interaction with continental crust. The minor scatter in the chromitite Os-isotope values was attributed to hydrothermal processes. Molybdenite (Re-rich and Os-poor) was separated from the G-chromitite and yielded an age of 2740, in agreement with the accepted age of the Complex. Marcantonio et al. (1993) suggested that all elevated gOs values found in the Complex could have resulted from interaction with hydrothermal fluids of the type which produced the molybdenite The Great Dyke, Zimbabwe. Schoenberg et al. (2003) studied chromite separates from chromitites of the 2576 Ma Great Dyke. Results from their study show that most chromitites form an imprecise isochron of 2580 ± 500 Ma (Fig. 29). Mixing calculations involving Archean country rocks suggest that 0–33% of the Os in the chromitites may have been of crustal origin. When combined with other isotopic and petrologic data the authors concluded that the source magmas underwent no interaction with subcontinental lithospheric mantle, and that the parent magma of the Great Dyke was plume-derived with a heterogeneous Os isotope signature, not unlike some oceanic basalts where the incorporation of recycled oceanic crust has been proposed (see for example Day 2013; Gannoun et al. 2016, this volume). Significant contamination by Archean granitoid crust was ruled out, based on the low-Os concentrations and unradiogenic Os-isotope ratios of representative country-rock samples.

Figure 29. Isochron diagram for samples from the Great Dyke (modified from Schoenberg et al. 2003).

The Bushveld Complex, South Africa. Re–Os isotope studies of chromitites, pyroxenites, sulfides, and alloys have been conducted as an aid in evaluating the genesis of PGE enrichment in the UG2 chromitite (McCandless and Ruiz 1991; Schoenberg et al. 1999), the Merensky Reef (Hart and Kinloch 1989; McCandless and Ruiz 1991; Schoenberg et al. 1999), and the Platreef (Reisberg et al. 2011). Results of these studies have been summarized by Reisberg et al. (2011), and are shown here in Figure 30. The results from all of these studies indicate that the Bushveld parent magmas underwent variable degrees of contamination by crust, in agreement with other isotopic indicators of crustal contamination. Schoenberg et al. (1999) noted that chromitites show increasing 187Os/ 188Os values upward, from near chondritic values in the lower chromitites to strongly enriched values in the upper chromitites. The enrichment was thought to be a result of enhanced degrees of crustal contamination. Hart and Kinloch (1989) analyzed erlichmanite and laurite grains from the Merensky Reef, and found that erlichmanite had Os-isotope ratios that were near to chondritic, with 187Os/ 188Os values similar to those of the lower chromitites.

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Figure 30. 187Os/ 188 Os values from the Bushveld Complex (modified from Reisberg et al. 2011). The Sandsloot pyroxenites are from the Platreef and have values very similar to those of the Merensky Reef (Reisberg et al. 2011).

Laurite grains were strongly radiogenic, with 187Os/188Os values similar to those of sulfides and whole rocks from the Reef reported by both McCandless and Ruiz (1991) and Schoenberg et al. (1999). McCandless and Ruiz (1991) suggested that the contaminant was Re-rich black shales that occur below the Complex. Schoenberg et al. (1999) proposed that primitive mantle-derived magma mixed with highly contaminated granophyric roof melt near the top of the chamber. In both scenarios, the Merensky Reef records a substantial contribution of crustally derived Os. The chondritic erlichmanite grains found in the Reef by Hart and Kinloch (1989) suggest that uncontaminated mantle-derived magma was first emplaced at the level of the Reef, but that highly crustally-contaminated magma followed. It is also possible that the erlichmanite grains represent mantle xenocrysts and that crustally contaminated magmas from chambers at depth entrained them. The radiogenic signals from pyroxenites of the Platreef also led Reisberg et al. (2011) to propose that the Platreef parental magma was crustally contaminated by black shales, and well-mixed prior to emplacement. They note that the rapid return to less radiogenic Osisotope compositions above the Merensky Reef cannot mark a change from one magma type to another, but must represent the input of a large quantity of radiogenic Os at the level of the reef. The spike in 187Os/188Os values at the level of the Merensky Reef and the Platreef was thought to indicate the contamination of magma by a crustal lithology unusually rich in radiogenic Os—most likely black shales. The Reisberg et al. (2011) model differs somewhat from the McCandless and Ruiz (1991) model in that Reisberg et al. (2011) propose contamination in magma chambers at intermediate crustal depths, rather than immediately below the current

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base of the Complex. Using representative concentrations and 187Os/188Os ratios Reisberg et al. (2011) determined that less than 1% bulk assimilation of black shale was necessary to account for the enriched signals in the Platreef. They also suggested that minor additional crustal contamination may have occurred at the level of emplacement of the Platreef. Summary. The Re–Os isotope system has obvious utility in the age-dating of magmatic ore deposits. In addition, the system may be a sensitive indicator of the importance of crustal contamination in ore formation. Results range from those where there is very little Os isotope indication of crustal contamination being involved in ore genesis (e.g., Noril’sk, Great Dyke) to those where the addition of crustal Os is strongly indicated (e.g., Sudbury, Voisey’s Bay, Bushveld). In conduit deposits such as Noril’sk and Eagle where other isotopic systems (e.g., S, Rb-Sr) support a strong role for crustal contamination. A process such as isotopic exchange and mixing involving sulfide-rich magma in the conduit and uncontaminated magmatic pulses may explain the decoupling between the isotopic systems. Another end-member interpretation is that in some environments country rock contaminants may contain insufficient Re and Os to perturb the mantle characteristics of the source magma, whereas concentrations and isotopic values are sufficient to readily perturb other isotopic systems such as S. A third interpretation would be that in systems where no Os isotope evidence for contamination exists, the magma was derived from a mantle source that had been modified by the addition of crustal components such a S (e.g., addition of altered crust via subduction) and as a result was characterized by anomalous S isotopic ratios, but the amount of radiogenic Os added was too small to perturb the mantle Os signal. Subcontinental lithospheric mantle is thought to be Re-depleted and hence melts derived from that source should be expected to have negative gOs values. The Re–Os isotopic system will no doubt continue to be a major tool for the evaluation of the source of Os and the significance of crustal contamination in magmatic Ni–Cu–PGE deposits. Care must be taken to ensure that isotope perturbations have not been caused by hydrothermal activity and the mobility of Re and Os in relatively low-temperature fluids.

CONCLUSIONS Most PGE ore deposits are found within or at the contact with mafic–ultramafic intrusions, and are associated with BMS minerals such as pyrrhotite, pentlandite, and chalcopyrite. In addition to BMS, the PGE are present in PGM found among the BMS and at the contact between BMS and chromite or silicates. The standard model for the formation of the deposits is that a mafic to ultramafic magma became saturated in a base-metal sulfide liquid and the PGE partitioned into this sulfide liquid. The PGE content of the liquid is dependent on the composition of the silicate magma, but also on the amount of magma with which it interacted. In order to form a PGE ore deposit, the sulfide liquid must interact with a very large amount of silicate magma. Some deposits such as the droplet ore of Noril’sk 1 or Merensky reef and Main sulfide zone may have formed in this manner. Other deposits require additional processes to explain the variations in PGE content of the ores. Fractional crystallization of the sulfide liquid is an important process that leads to the formation of an MSS cumulate depleted in Cu but enriched in Re, Os, Ir, Ru, and Rh. The fractionated liquid is enriched in Cu, Pt, Pd, Au, and Te, As, Bi, Sb (TABS). Intermediate solid solution sulfide crystallizes from the fractionated liquid but Pt, Pd, Au, and TABS do not partition into ISS, and they concentrate into the very last sulfide liquid. In most cases this liquid is trapped with ISS cumulate and as a result the Cu-rich ore is also Pt-, Pd-, Au- and TABSrich. The variations in PGE contents in some ores such as the McCreedy East deposit of Sudbury or the Talnakh and Kharaelakh deposits of the Noril’sk area are the result of this crystal fractionation. The Pt–Pd–Au–TABS-rich liquid would not crystallize until relatively low temperatures (500–600 oC). Therefore, it is possible that the liquid could be injected into

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the fractures around or within the intrusion during some structural disturbance. Also even once crystallized, it could melt at lower amphibolite faces conditions and migrate during metamorphism. Possibly the low-S, high-Pt–Pd deposits such as the High Grade ore of Lac des Iles or Broken Hammer of Sudbury formed by mobilization of Pd–Pt–TABS liquid. Because of the chalcophile nature of both Re and Os, the Re–Os isotopic system is widely applied in the evaluation of ore genesis. The system has wide applicability in the geochronology of magmatic ore systems. In addition, the long-term enrichment of Re in the crust has made the system a very sensitive one in the evaluation of the degree of interaction between mantlederived magmas and continental crust. Osmium-isotope ratios may provide an indication of the source of Os in an ore deposit, as well as the extent of country rock contamination that may have been important in the ore-forming process.

ACKNOWLEDGMENTS The research of SJB has been supported by Natural Sciences and Engineering Research Council of Canada for the past 29 years and by the Canada Research Chair in Magmatic Metallogeny for 10 years. Both sources of funds are for fundamental research and the continuity of this support has been essential to the development of SJB’s research program. The research of EMR on the genesis of magmatic ore deposits has been supported by the United States National Science Foundation for the past 38 years; appreciation is expressed for both basic research support and analytical instrumentation. We also thank all those members of the magmatic ore deposits community who organized and participated in conferences and field trips to the world’s Ni and PGE deposits; seeing these deposits and discussions with colleagues on these occasions greatly helped to clarify our ideas. We thank the editors Dr. Harvey and Dr. Day for the invitation to write this chapter and the reviewers Dr. D. Holwell and Dr B. O’Driscoll for their time and effort in helping to improve the manuscript.

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