Oxygen in the Solar System 9780939950805

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Frontispiece (on facing page): This diagram is a compilation of oxygen isotopic analytical data for meteoritic and lunar samples, collected since about 1973 in the University of Chicago lab of Robert N. Clayton and Toshiko Mayeda. δ18O and δ17O give the deviations, in parts per 1000 (permil; ‰), of the ratios 18O/16O and 17O/16O, respectively, in samples relative to Standard Mean Ocean Water (SMOW). The diagram is, in effect, an isotopic map of solar system bodies. All terrestrial and lunar samples plot on the so-called terrestrial fractionation line (TF), which has a slope of ~½, along which the variation can be explained by simple mass-dependent, physicochemical processes such as evaporation, condensation, and igneous crystallization. Meteoritic samples mostly do not plot on this line, and either plot on separate slope-½ lines parallel to but displaced from TF, or else along linear arrays having slopes closer to 1. The Carbonaceous Chondrite Anhydrous Mineral (CCAM) line is the dominant slope-1 array, whose extreme 16O-rich end is defined by calcium-, aluminum-rich inclusions (CAIs). The CCAM line requires non-mass-dependent isotopic effects or mixing of an additional 16O-rich component. Some CAIs disperse to the right of the CCAM line owing to melt volatilization that resulted in massdependent isotopic fractionation (e.g., the FUN line).

Reviews in Mineralogy and Geochemistry, Volume 68 Oxygen in the Solar System ISSN 1529-6466 ISBN 978-0-939950-80-5

Copyright 2008

The MINERALOGICAL SOCIETY of AMERICA 3635 Concorde Parkway, Suite 500 Chantilly, Virginia, 20151-1125, U.S.A. www.minsocam.org 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.

REVIEWS IN MINERALOGY AND GEOCHEMISTRY Volume 68

2008

OXYGEN IN THE SOLAR SYSTEM SCIENCE EDITOR-IN-CHIEF Glenn J. MacPherson Smithsonian Institution, Washington DC

SCIENCE CO-EDITORS David W. Mittlefehldt and John H. Jones NASA Johnson Space Center, Houston TX

TECHNICAL EDITOR Steven B. Simon University of Chicago, Chicago IL

OXYGEN INITIATIVE JOINT CHAIRS James J. Papike (Univ. of New Mexico, Albuquerque, NM) Stephen Mackwell (Lunar and Planetary Institute, Houston TX)

Series Editor: Jodi J. Rosso MINERALOGICAL SOCIETY OF AMERICA GEOCHEMICAL SOCIETY LUNAR AND PLANETARY INSTITUTE

DEDICATION TO ROBERT N. CLAYTON

It is most fitting that this volume be dedicated to Professor Robert N. Clayton who, among geochemists and cosmochemists, could easily wear the name “Mr. Oxygen.” His 1973 discovery (with coauthors Larry Grossman and Toshiko Mayeda), that calcium-, aluminumrich inclusions (CAIs) in the Allende meteorite have oxygen isotopic (specifically, 16O-rich) compositions that cannot be explained by simple physical processes such as condensation or evaporation, led to the more general discovery of isotope anomalies in CAIs. Working with his long-time associate Tosh Mayeda and many other colleagues over the years, Bob demonstrated the power of 3-isotope oxygen measurements to fingerprint different meteorite classes and, by extension, their different parent bodies. Although the original interpretation of the 16Orich signature in CAIs and other early solar system materials as being presolar in origin has since been abandoned even by Bob himself, the intense interest generated by the 1973 work revolutionized cosmochemistry in a way that continues to this day. Perhaps one way of putting into perspective the giant stature of Bob Clayton within the planetary science and isotope chemistry community is to note the following observation. Isotope geochemists (a group that includes many of us) tend to be a disputatious lot who will readily disagree with one another at meetings over just about anything large or small. Yet Bob is one person with whom most of this community will only rarely disagree: he very commonly is allowed to have The Last Word. In our experience, it is a rare thing for anyone to be so revered. Bob has already received many high honors, not least being the National Medal of Science, election to the U.S. National Academy of Sciences and the Royal society of London, and the Leonard Medal of the Meteoritical Society. The significance of dedicating this volume to him, although seemingly small in comparison with those other exalted tributes, is that this comes directly from a large number of Bob’s friends and colleagues with whom he has worked and collaborated over many years. Thank you, Bob, for your many great accomplishments and for being an inspiration to us all. The authors and editors of this volume are proud to dedicate it in your honor, and hope that it lives up to your high standards.

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FROM THE SERIES EDITOR This volume was jointly published by the Mineralogical Society of America (MSA) and the Lunar and Planetary Institute (LPI). Such huge undertakings rarely go with out hitting a few “bumps in the road,” yet Glenn MacPherson and his team prevailed and produced a finished product all could be proud of. Steve Simon, the behind-the-scenes guy, made my job exceptionally easy and I thank him for all his hard work! Plus, he created the index for this volume—a rarity in this series yet greatly appreciated! You can learn more about this volume and the “Oxygen Inititive” in the Introduction written by Glenn MacPherson. Any supplemental material and errata (if any) can be found at the MSA website www. minsocam.org. A searchable, electronic version of this volume can be found on the GeoScienceWorld website www.geoscienceworld.org. Jodi J. Rosso, Series Editor West Richland, Washington January 2008

1529-6466/08/0068-0000$05.00

DOI: 10.2138/rmg.2008.68.0

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TABLE OF CONTENTS

1

Introduction Glenn J. MacPherson

...................................................................................................................................................1

2

Oxygen Isotopes in the Early Solar System A Historical Perspective Robert N. Clayton

ABSTRACT ..............................................................................................................................5 BEFORE ALLENDE ................................................................................................................5 AFTER ALLENDE...................................................................................................................6 FUN CAIs .................................................................................................................................8 OXYGEN ISOTOPES IN PRESOLAR GRAINS ....................................................................9 CHEMICAL ISOTOPE EFFECTS ...........................................................................................9 PHOTOCHEMICAL EFFECTS .............................................................................................10 INTERNAL ASTEROIDAL PROCESSES ............................................................................10 NITROGEN ............................................................................................................................11 CONCLUSIONS.....................................................................................................................12 ACKNOWLEDGMENT .........................................................................................................12 REFERENCES .......................................................................................................................12

3

Abundance, Notation, and Fractionation of Light Stable Isotopes Robert E. Criss, James Farquhar

ABSTRACT ............................................................................................................................15 INTRODUCTION ..................................................................................................................15 vii

Oxygen in the Solar System ‒ Table of Contents ISOTOPIC ABUNDANCES AND ATOMIC WEIGHTS.......................................................16 NOTATION ............................................................................................................................18 Isotope ratios ...............................................................................................................18 δ-values........................................................................................................................19 Isotopic fractionation factor ........................................................................................19 Big delta and related approximations ..........................................................................19 Capital delta .................................................................................................................20 Capital delta prime and delta prime.............................................................................20 Material balance ..........................................................................................................21 COMMONLY-USED DIAGRAMS .......................................................................................21 δ−δ plot ........................................................................................................................21 Big Δ and Cap Δ plots..................................................................................................21 Three-isotope plot........................................................................................................22 ISOTOPIC FRACTIONATION PROCESSES .......................................................................24 Mass-dependent fractionation ....................................................................................24 Kinetic processes .........................................................................................................26 Non-mass-dependent fractionations ............................................................................27 CONCLUSIONS.....................................................................................................................28 REFERENCES .......................................................................................................................29

4

Nucleosynthesis and Chemical Evolution of Oxygen Bradley S. Meyer, Larry R. Nittler, Ann N. Nguyen, Scott Messenger

ABSTRACT ............................................................................................................................31 INTRODUCTION ..................................................................................................................31 NUCLEOSYNTHESIS OF THE ISOTOPES OF OXYGEN .................................................32 Production of oxygen in mainline stellar burning stages ............................................32 Analysis of the oxygen yields from massive stars .......................................................36 Low-mass stars ............................................................................................................38 Novae and Type Ia supernovae ....................................................................................41 CHEMICAL EVOLUTION OF THE ISOTOPES OF OXYGEN..........................................41 OXYGEN IN PRESOLAR GRAINS .....................................................................................45 Oxygen in carbonaceous grains ...................................................................................46 Presolar oxide and silicate grains ................................................................................47 CONCLUDING REMARKS ..................................................................................................50 ACKNOWLEDGMENTS.......................................................................................................51 REFERENCES .......................................................................................................................51

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Oxygen in the Interstellar Medium Adam G. Jensen, F. Markwick-Kemper, Theodore P. Snow

ABSTRACT ............................................................................................................................55 INTRODUCTION ..................................................................................................................55 Phases in the interstellar medium ................................................................................56 Forms of oxygen in the interstellar medium................................................................56 OXYGEN IN THE GAS PHASE ...........................................................................................56 Measurements of gas-phase oxygen ............................................................................56 Isotope measurements from gas-phase oxygen and carbon monoxide .......................60 Inferring gas-phase depletions of oxygen....................................................................61 OXYGEN IN INTERSTELLAR DUST .................................................................................63 Solar System silicates ..................................................................................................63 Silicates in circumstellar environments of young stars ...............................................63 Dust properties in the interstellar medium ..................................................................64 Dust production by evolved stars ................................................................................65 CONSISTENCY BETWEEN GAS AND SOLID PHASES ..................................................66 Abundance and depletion constraints ..........................................................................66 Transitions between the solid and gas phase in the interstellar medium .....................67 SUMMARY ............................................................................................................................68 REFERENCES .......................................................................................................................68

6

Oxygen in the Sun Andrew M. Davis, Ko Hashizume, Marc Chaussidon, Trevor R. Ireland, Carlos Allende Prieto, David L. Lambert

ABSTRACT ............................................................................................................................73 INTRODUCTION ..................................................................................................................74 THE SOLAR PHOTOSPHERIC ABUNDANCE OF OXYGEN...........................................74 OXYGEN ISOTOPIC COMPOSITION OF THE SUN .........................................................77 Predictions of the isotopic composition of the Sun .....................................................77 Spectroscopic constraints on the oxygen isotopic composition of the Sun .................78 Identification of the solar isotopic composition trapped in lunar samples .................79 Oxygen isotopic composition of the solar wind: direct measurements .......................87 Summary of solar oxygen isotopic composition .........................................................88 ACKNOWLEDGMENTS.......................................................................................................89 REFERENCES .......................................................................................................................89

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Redox Conditions in the Solar Nebula: Observational, Experimental, and Theoretical Constraints Lawrence Grossman, John R. Beckett, Alexei V. Fedkin, Steven B. Simon, Fred J. Ciesla

ABSTRACT ............................................................................................................................93 INTRODUCTION ..................................................................................................................94 OXYGEN FUGACITY DURING CRYSTALLIZATION OF REFRACTORY INCLUSION MELTS.............................................................................94 Experimental technique ...............................................................................................94 Results .........................................................................................................................96 Thermochemistry.........................................................................................................99 Selection of fassaite-melilite pairs.............................................................................103 Oxygen barometry .....................................................................................................105 THE OXIDATION STATE OF IRON IN ORDINARY CHONDRITES .............................109 The problem...............................................................................................................109 Radial transport processes .........................................................................................111 Vertical transport processes .......................................................................................112 Relationship between fO2 of cosmic gases and abundances of C, O and H................114 Condensation of fayalitic olivine...............................................................................115 Change of FeO/(FeO + MgO) during chondrule melting..........................................124 REDOX CONDITIONS INFERRED FROM OTHER IRON-BEARING NEBULAR MATERIALS...............................................................................................126 Amoeboid olivine aggregates ....................................................................................126 Metal grains in CH chondrites...................................................................................127 FORMATION CONDITIONS OF ENSTATITE CHONDRITES........................................127 Mineralogy of EH3 enstatite chondrites ....................................................................127 Condensation at high C/O ratio .................................................................................127 Bulk chemical compositions of EH enstatite chondrites ...........................................130 Condensation of EH enstatite chondrites ..................................................................130 Formation conditions of EH3 enstatite chondrites ....................................................134 CONCLUSIONS...................................................................................................................135 ACKNOWLEDGMENTS.....................................................................................................136 REFERENCES .....................................................................................................................136

8

Oxygen Isotopes of Chondritic Components Hisayoshi Yurimoto, Alexander N. Krot, Byeon-Gak Choi, Jerome Aléon, Takuya Kunihiro, Adrian J. Brearley

ABSTRACT ..........................................................................................................................141 INTRODUCTION ................................................................................................................142 CHONDRITES AND THEIR COMPONENTS ...................................................................144 x

Oxygen in the Solar System ‒ Table of Contents OXYGEN ISOTOPIC COMPOSITIONS OF SECONDARY PHASES .............................145 OXYGEN ISOTOPIC COMPOSITIONS OF REFRACTORY INCLUSIONS...................149 Alteration and secondary minerals of fine grained CAIs (FGIs)...............................149 ORIGINAL OXYGEN ISOTOPIC DISTRIBUTION OF FGIs ...........................................151 FGIs in primitive O chondrites .................................................................................151 FGIs in primitive E chondrites ..................................................................................151 FGIs in CO 3.0 chondrites ........................................................................................151 FGIs in CR chondrites ..............................................................................................152 FGIs in CM chondrites ..............................................................................................154 FGIs in CV chondrites ...............................................................................................154 FGIs in CH chondrites ...............................................................................................155 FGIs in CB chondrites ...............................................................................................155 Chondrule-bearing FGI .............................................................................................155 Summary of oxygen isotopic characteristics of FGIs................................................156 Alteration and secondary minerals of amoeboid olivine aggregates (AOAs) ...........156 ORIGINAL OXYGEN ISOTOPIC DISTRIBUTION OF AOAs .........................................158 AOAs in CV chondrites .............................................................................................158 AOAs in CO, CR, Acfer 094 and CM chondrites......................................................159 Summary of oxygen isotopic characteristics of AOAs ..............................................161 OXYGEN ISOTOPIC DISTRIBUTION OF COARSE-GRAINED CAIs (CGIs) ..............161 7R-19-1, a compact Type A CGI ...............................................................................163 E49, a compact Type A CGI ......................................................................................163 SS-02, a Type B2 CGI ..............................................................................................165 TTV1-01, a Type B2 CGI ..........................................................................................166 1623-2, a compact Type A CGI .................................................................................166 V2-01, a fluffy Type A CGI .......................................................................................167 Chondrule-bearing CGIs ...........................................................................................168 Summary of oxygen isotopic characteristics of CGIs ...............................................169 OXYGEN ISOTOPIC COMPOSITIONS OF CHONDRULES...........................................169 Chondrules in CH chondrites ....................................................................................170 Chondrules in CR chondrites ....................................................................................170 Chondrules in CB chondrites ....................................................................................170 Chondrules in CO and Acfer 094 chondrites.............................................................171 Chondrules in CV chondrites ....................................................................................172 Chondrules in E chondrites .......................................................................................172 Chondrules in ordinary chondrites ............................................................................172 Refractory inclusion-bearing chondrules ..................................................................172 Summary of oxygen isotopic characteristics of chondrules ......................................174 OXYGEN ISOTOPIC COMPOSITIONS OF MATRIX ......................................................175 Existence of submicron silicate grains with extreme non-solar oxygen isotopic compositions ........................................................................................................176 Oxygen isotopic heterogeneity of matrix ..................................................................177 Summary of oxygen isotopic characteristics of matrix .............................................179 IMPLICATIONS FOR ASTROPHYSICAL SETTING OF CHONDRITIC COMPONENT FORMATION........................................................................................179 ACKNOWLEDGMENTS.....................................................................................................181 REFERENCES .....................................................................................................................182

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Mass-independent Oxygen Isotope Variation in the Solar Nebula Edward D. Young, Kyoshi Kuramoto, Rudolph A. Marcus, Hisayoshi Yurimoto, Stein B. Jacobsen

ABSTRACT ..........................................................................................................................187 INTRODUCTION ................................................................................................................188 GALACTIC OXYGEN ISOTOPE EVOLUTION–A NON-CHEMICAL PATH TO MASS INDEPENDENCE .............................................................................189 Testing the hypothesis – the oxygen isotopic composition of the Sun ......................191 CHEMICAL MASS-INDEPENDENT OXYGEN ISOTOPE FRACTIONATION ............192 The MIF in ozone formation .....................................................................................193 Conditions for a chemical MIF in the formation of CAIs .........................................195 A possible chemical mechanism for MIF in CAIs ....................................................195 Consequences of chemical mechanism for MIF in the early water...........................196 Testing the hypothesis: experiment to test gas phase MIF at high temperature........197 PHOTOCHEMICAL MASS-INDEPENDENT OXYGEN ISOTOPE FRACTIONATION: CO SELF-SHIELDING ................................................................198 CO photodissociation and self-shielding ...................................................................198 Astronomical observations of oxygen isotope fractionation by CO self-shielding ...200 The pivotal role of H2O .............................................................................................201 CO self-shielding at the inner annulus of the solar circumstellar disk ......................203 CO self-shielding at the surfaces of the solar circumstellar disk ..............................204 CO self-shielding in molecular clouds and inheritance in the Solar System ............210 Testing the hypotheses: predictions of the CO self-shielding models .......................212 SUMMARY ..........................................................................................................................213 REFERENCES .....................................................................................................................214

10

Oxygen and Other Volatiles in the Giant Planets and their Satellites Michael H. Wong, Jonathan I. Lunine, Sushil K. Atreya, Torrence Johnson, Paul R. Mahaffy, Tobias C. Owen, Thérèse Encrenaz

ABSTRACT ..........................................................................................................................219 INTRODUCTION ................................................................................................................220 Oxygen-based insights from the outer planets and their moons................................220 The protosolar abundances ........................................................................................221 MEASURING OXYGEN IN JUPITER’S ATMOSPHERE .................................................222 Structure of the cloud layers ......................................................................................222 Galileo Probe Mass Spectrometer water mixing ratio measurements .......................223 The probe entry site: A 5-μm hot spot.......................................................................225 Spectroscopic measurements of Jovian water ...........................................................226 xii

Oxygen in the Solar System ‒ Table of Contents Lightning on Jupiter ..................................................................................................227 Oxygen isotopes in Jupiter ........................................................................................228 Summary of Jovian oxygen .......................................................................................228 OUTER PLANET VOLATILE GASES ...............................................................................229 Oxygen and other heavy element enrichments in Jupiter ..........................................229 Volatile enrichments in the other outer planets .........................................................231 OXYGEN IN OUTER PLANET SATELLITES ..................................................................232 Jupiter’s satellites ......................................................................................................234 Saturn’s satellites .......................................................................................................234 Outer Solar System satellites and Kuiper Belt Objects .............................................235 FORMATION OF THE OUTER PLANETS........................................................................236 Volatile enrichment by icy planetesimals ..................................................................237 Volatile enrichment by carbonaceous planetesimals .................................................238 Volatile enrichment by disk evolution .......................................................................239 CONCLUSIONS...................................................................................................................240 ACKNOWLEDGMENTS.....................................................................................................241 REFERENCES .....................................................................................................................241

11

Oxygen in Comets and Interplanetary Dust Particles Scott A. Sandford, Scott Messenger, Michael DiSanti, Lindsay Keller, Kathrin Altwegg

ABSTRACT ..........................................................................................................................247 INTRODUCTION ................................................................................................................247 THE CHEMICAL FORM OF OXYGEN IN THE INTERSTELLAR MEDIUM, “COMETARY” INTERPLANETARY DUST PARTICLES, AND COMETS ...............249 Oxygen carried by carbonaceous materials in the interstellar medium, meteorites, cosmic dust, and cometary samples.....................................................................249 Direct detection of oxygen-bearing volatiles in comets ............................................253 The oxygen-bearing minerals in “cometary” IDPs and samples from comet 81P/Wild 2 ...........................................................................................................258 OXYGEN ISOTOPES IN THE INTERSTELLAR MEDIUM, COMETS, COMETARY SAMPLES, AND “COMETARY” IDPS..................................................260 Oxygen isotopes in interstellar materials ..................................................................261 In situ measurement of the oxygen isotopes in the volatile material of comet Halley ...................................................................................................261 Oxygen isotopic compositions of meteorites, “cometary” IDPs and samples from comet 81P/Wild ............................................................................262 FUTURE IN SITU MEASUREMENTS OF ISOTOPIC RATIOS IN COMETS.................264 CONCLUSIONS...................................................................................................................265 ACKNOWLEDGMENTS.....................................................................................................265 REFERENCES .....................................................................................................................265

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Oxygen and Asteroids Thomas H. Burbine, Andrew S. Rivkin, Sarah K. Noble, Thais Mothé-Diniz, William F. Bottke, Timothy J. McCoy, M. Darby Dyar, Cristina A. Thomas

ABSTRACT ..........................................................................................................................273 INTRODUCTION ................................................................................................................273 DYNAMICAL STRUCTURE OF THE ASTEROID BELT ................................................275 ASTRONOMICAL TECHNIQUES .....................................................................................276 Brightness ..................................................................................................................276 Reflectance spectroscopy...........................................................................................276 Spectral data ..............................................................................................................278 Corrections ................................................................................................................279 Interaction of photons with a surface ........................................................................281 ABSORPTION BANDS .......................................................................................................281 Electronic absorption features ...................................................................................281 Vibrational absorption features..................................................................................288 SPACE WEATHERING........................................................................................................291 Effect of space weathering on reflectance spectra .....................................................291 Space weathering environment of asteroids ..............................................................293 Experimental studies .................................................................................................293 Evidence of space weathering on asteroids ...............................................................294 Implications for visible/near-IR remote sensing .......................................................294 ORDINARY CHONDRITES, LODRANITES/ACAPULCOITES, AND UREILITES ......295 DETERMINING MINERAL CHEMISTRIES ....................................................................296 Determining the ratio of olivine to pyroxene ............................................................296 Modified Gaussian Modeling ....................................................................................297 ASTEROID TAXONOMY ...................................................................................................297 A-types ......................................................................................................................301 C-complex .................................................................................................................304 D- and P-types ...........................................................................................................306 E- and Xe-types ........................................................................................................307 K- and L-types ...........................................................................................................308 M-types ......................................................................................................................308 O-types ......................................................................................................................310 Q-types ......................................................................................................................310 R-types.......................................................................................................................310 S-complex ..................................................................................................................310 T-types .......................................................................................................................312 V-types .......................................................................................................................312 HELIOCENTRIC DISTRIBUTIONS OF TAXONOMIC CLASSES IN THE MAIN BELT......................................................................................................314 DISTRIBUTION OF HYDRATED ASTEROIDS IN THE MAIN BELT ...........................321 NEAR-EARTH ASTEROIDS...............................................................................................323 SPACECRAFT MISSIONS ..................................................................................................324 COLLISIONAL AND DYNAMICAL EVOLUTION OF ASTEROIDS .............................326 xiv

Oxygen in the Solar System ‒ Table of Contents DELIVERY OF METEOROIDS TO EARTH ......................................................................327 THE EFFECTS OF PLANETARY EMBRYOS AND RADIAL MIXING IN THE MAIN BELT......................................................................................................329 COULD IRON METEORITES HAVE COME FROM THE TERRESTRIAL PLANET REGION?........................................................................................................330 SUMMARY ..........................................................................................................................331 ACKNOWLEDGMENTS.....................................................................................................331 REFERENCES .....................................................................................................................331

13

Oxygen Isotopes in Asteroidal Materials Ian A. Franchi

ABSTRACT ..........................................................................................................................345 INTRODUCTION ................................................................................................................346 ORDINARY CHONDRITES................................................................................................349 Introduction ...............................................................................................................349 Ordinary chondrites – whole-rock .............................................................................349 Ordinary chondrites – components............................................................................351 R CHONDRITES..................................................................................................................356 ENSTATITE METEORITES ................................................................................................358 Introduction ...............................................................................................................358 EH and EL chondrites ...............................................................................................358 Aubrites .....................................................................................................................360 CARBONACEOUS CHONDRITES ....................................................................................361 Introduction ...............................................................................................................361 CV chondrites ............................................................................................................361 CK chondrites ............................................................................................................365 CO chondrites ............................................................................................................366 CM chondrites ...........................................................................................................368 CI chondrites .............................................................................................................371 CR chondrites ............................................................................................................372 CH chondrites ............................................................................................................373 CB chondrites ............................................................................................................374 PRIMITIVE ACHONDRITES .............................................................................................375 Introduction ...............................................................................................................375 Acapulcoites and lodranites.......................................................................................376 Brachinites .................................................................................................................377 Winonaites .................................................................................................................377 Ureilites .....................................................................................................................378 BASALTIC ACHONDRITES...............................................................................................379 Introduction ...............................................................................................................379 Howardites, eucrites and diogenites ..........................................................................380 Angrites .....................................................................................................................381 Basaltic inclusions .....................................................................................................382 IRONS AND STONY-IRONS ..............................................................................................382 Introduction ...............................................................................................................382 xv

Oxygen in the Solar System ‒ Table of Contents IAB Complex.............................................................................................................383 IIAB ...........................................................................................................................384 IIE ..............................................................................................................................384 IIIAB..........................................................................................................................385 IVA ............................................................................................................................386 Mesosiderites .............................................................................................................387 Pallasites ....................................................................................................................387 Ungrouped irons ........................................................................................................388 CONCLUSIONS...................................................................................................................389 REFERENCES .....................................................................................................................390

14

Oxygen Isotopic Composition and Chemical Correlations in Meteorites and the Terrestrial Planets David W. Mittlefehldt, Robert N. Clayton, Michael J. Drake, Kevin Righter

ABSTRACT ..........................................................................................................................399 INTRODUCTION ................................................................................................................400 BACKGROUND ...................................................................................................................400 Nebular element fractionations..................................................................................400 Oxygen isotope anomalies.........................................................................................403 Mechanisms of non-mass-dependent isotope fractionation ......................................405 CHONDRITIC METEORITES ............................................................................................407 Micro- and meso-scale correlations...........................................................................407 Correlations among chondrite groups .......................................................................408 UREILITES ..........................................................................................................................417 TERRESTRIAL PLANETS .................................................................................................419 SUMMARY ..........................................................................................................................422 ACKNOWLEDGMENTS.....................................................................................................423 REFERENCES .....................................................................................................................423

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Record of Low-Temperature Alteration in Asteroids Michael E. Zolensky, Alexander N. Krot, Gretchen Benedix

ABSTRACT ..........................................................................................................................429 INTRODUCTION ................................................................................................................429 C, P AND D ASTEROIDS – CARBONACEOUS CHONDRITES .....................................430 Aqueous activity on the CI parent asteroid(s) and its oxygen isotope record ...........430 Aqueous activity on the CM parent asteroid(s) and its oxygen isotope record .........434 Oxygen isotopic compositions of secondary minerals in the ungrouped carbonaceous chondrite Tagish Lake ....................................................................437 Aqueous alteration of CR chondrites and their oxygen isotope record .....................438 Hydrous and anhydrous alteration of CV chondrites and their oxygen isotope records ......................................................................................................439 Low-temperature aqueous alteration of CO chondrites.............................................448 Veritas asteroids – hydrous chondritic interplanetary dust particles .........................448 S ASTEROIDS – ORDINARY AND R CHONDRITES......................................................451 Aqueous alteration of ordinary chondrites and its oxygen isotope record ................451 Aqueous alteration of R-chondrites and its oxygen isotope record...........................452 M AND E ASTEROIDS – INCLUDING ENSTATITE CHONDRITES .............................452 OXYGEN ISOTOPIC COMPOSITION OF ASTEROIDAL WATER AND EVOLUTION OF OXYGEN ISOTOPIC COMPOSITION OF THE INNER PROTOPLANETARY DISK ..........................................................................................454 SUMMARY AND FUTURE WORK ...................................................................................455 ACKNOWLEDGMENTS.....................................................................................................456 REFERENCES .....................................................................................................................456

16

The Oxygen Cycle of the Terrestrial Planets: Insights into the Processing and History of Oxygen in Surface Environments James Farquhar, David T. Johnston

ABSTRACT ..........................................................................................................................463 INTRODUCTION ................................................................................................................463 ISOTOPIC VARIATIONS AMONG TERRESTRIAL MATERIALS..................................464 Historical account of oxygen isotopic variations of terrestrial reservoirs .................465 Molecular oxygen ......................................................................................................471 Ozone.........................................................................................................................471 Other oxygen-bearing atmospheric species with nonzero Δ17O ................................473 Multiply substituted molecular species .....................................................................473 EVOLUTION OF OXYGEN IN EARTH’S SURFACE ENVIRONMENTS .....................474 Planetary processing of oxygen .................................................................................475 OBSERVATIONS RELEVANT TO THE EVOLUTION OF OXYGEN IN THE ATMOSPHERE AND OCEANS ...................................................................................476 xvii

Oxygen in the Solar System ‒ Table of Contents Hypotheses about the levels of oxygen in Earth’s early environments ....................476 The transition from a low-oxygen atmosphere to a high oxygen atmosphere...........480 Into the Paleoproterozoic and Mesoproterozoic .......................................................481 Oxygen and Proterozoic carbon cycle .......................................................................482 Oxygen concentration variations since the end of the Proterozoic ...........................483 Conceptual model for oxygenation of Earth surface environments ..........................484 NEW FRONTIERS ..............................................................................................................485 CONCLUDING STATEMENTS ..........................................................................................486 ACKNOWLEDGMENTS.....................................................................................................487 REFERENCES .....................................................................................................................487

17

Redox Conditions on Small Bodies, the Moon and Mars Meenakshi Wadhwa

ABSTRACT ..........................................................................................................................493 INTRODUCTION ................................................................................................................493 SMALL BODIES..................................................................................................................494 Brachinites and other primitive achondrites ..............................................................494 Ureilites .....................................................................................................................495 Aubrites .....................................................................................................................496 Angrites .....................................................................................................................496 Eucrites ......................................................................................................................496 THE MOON..........................................................................................................................497 MARS ...................................................................................................................................499 OTHER TERRESTRIAL PLANETS ...................................................................................503 SUMMARY AND CONCLUSIONS ....................................................................................505 ACKNOWLEDGMENTS.....................................................................................................506 REFERENCES .....................................................................................................................506

18

Terrestrial Oxygen Isotope Variations and Their Implications for Planetary Lithospheres Robert E. Criss

ABSTRACT ..........................................................................................................................511 INTRODUCTION ................................................................................................................511 OXYGEN ISOTOPE GEOCHEMISTRY OF TERRESTRIAL ROCKS ............................512 Earth’s primordial δ18O value ....................................................................................512 Oxygen isotope variations of terrestrial rocks ...........................................................513 ISOTOPIC FRACTIONATION PROCESSES ....................................................................514 Isotopic fractionation factors ....................................................................................514 Fractional crystallization and AFC ...........................................................................515 Real magmas ............................................................................................................516 Subsolidus fractionation processes ...........................................................................517 xviii

Oxygen in the Solar System ‒ Table of Contents OXYGEN ISOTOPE ZONATION AND HETEROGENEITY IN PLANETARY LITHOSPHERES ..........................................................................................................519 Processes producing 18O zonation ............................................................................519 Processes producing 18O heterogeneity ....................................................................521 Bulk 18O composition of the continents ....................................................................523 Isotopic changes over geologic time ........................................................................524 CONCLUSIONS ..................................................................................................................525 REFERENCES ....................................................................................................................525

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Basalts as Probes of Planetary Interior Redox State Christopher D. K. Herd

ABSTRACT ..........................................................................................................................527 INTRODUCTION ................................................................................................................527 THE OXIDATION STATE OF THE EARTH’S MANTLE .................................................528 The lower mantle .......................................................................................................530 The upper mantle .......................................................................................................531 OXYBAROMETERS APPLICABLE TO BASALTIC ROCKS ..........................................533 Oxygen fugacity from mineral equilibria ..................................................................535 Multivalent trace elements.........................................................................................541 Oxygen fugacity from multivalent trace elements.....................................................541 THE BASALT-MANTLE SOURCE REDOX RELATIONSHIP ........................................546 Is basalt oxygen fugacity reflective of the redox state of its mantle source? ............546 Implications for understanding the redox states of planetary interiors .....................548 ACKNOWLEDGMENTS.....................................................................................................549 REFERENCES .....................................................................................................................549

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Rheological Consequences of Redox State Stephen Mackwell

ABSTRACT ..........................................................................................................................555 INTRODUCTION ................................................................................................................555 DEFORMATION OF OLIVINE...........................................................................................556 Olivine single crystal studies .....................................................................................556 Olivine aggregate studies...........................................................................................558 How does oxygen fugacity affect creep of olivine? ..................................................562 DEFORMATION OF OTHER SILICATES .........................................................................564 How does oxygen fugacity affect creep of other silicates? ......................................566 SUMMARY ..........................................................................................................................567 REFERENCES .....................................................................................................................568

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Oxygen in the Solar System ‒ Table of Contents

Appendix: Meteorites – A Brief Tutorial David W. Mittlefehldt ABSTRACT ..........................................................................................................................571 INTRODUCTION ................................................................................................................571 CHONDRITES .....................................................................................................................572 Carbonaceous chondrites ...........................................................................................574 Ordinary chondrites ...................................................................................................575 Enstatite chondrites ...................................................................................................576 Rumuruti-like and Kakangari-like chondrites ...........................................................576 ACHONDRITES ..................................................................................................................576 Acapulcoite-lodranite clan.........................................................................................577 Winonaites and silicate inclusions from IAB (and possibly IIICD) irons .................579 Angrites .....................................................................................................................579 Aubrites .....................................................................................................................579 Brachinites .................................................................................................................580 Howardite-eucrite-diogenite clan ..............................................................................580 Ureilites .....................................................................................................................581 IRONS ...................................................................................................................................581 Magmatic iron meteorite groups ...............................................................................583 Non-magmatic iron meteorite groups........................................................................585 STONY IRONS ....................................................................................................................585 Main-group and Eagle Station grouplet pallasites.....................................................585 Mesosiderites .............................................................................................................586 ACKNOWLEDGMENTS.....................................................................................................587 REFERENCES .....................................................................................................................587

Subject Index ...................................................................................................................591 Meteorite Index ..............................................................................................................597

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 1-3, 2008 Copyright © Mineralogical Society of America

Introduction Glenn J. MacPherson Smithsonian Institution PO Box 37012, MRC 119 Washington, D.C. 20013-7012, U.S.A. [email protected]

Oxygen (O) Atomic No. = 8 Atomic Wt. = 15.9994 Solar System Abundance (relative to 106 Si atoms)1 = 1.413 × 107 Three stable isotopes: 16O, 17O, 18O Relative isotopic abundances on Earth2: 99.762% 16O, 0.038% 17O, 0.200% 18O First isolated as an element in the late 18th century by Scheele and Priestley (independently), oxygen is the third most abundant element in the universe. It is abundant on Earth as a colorless elemental gas in the atmosphere, in combination with hydrogen in water, and in combination with silicon and other metals in the silicates and oxides that make up Earth’s crust and mantle. Oxygen is chemically very electronegative, and only a few metals (mainly, the platinum group metals, gold, silver, mercury, and copper) persist in their elemental form on Earth’s surface in the presence of the oxygenated atmosphere. The abundance of elemental oxygen in Earth’s atmosphere is more or less steady-state, being produced and sustained by the action of photosynthetic plants against the constant removal by physical and biogenic oxidation processes.

Hydrogen may be the most abundant element in the universe, but in science and in nature oxygen has an importance that is disproportionate to its abundance. Human beings tend to take it for granted because it is all around us and we breathe it, but consider the fact that oxygen is so reactive that in a planetary setting it is largely unstable in its elemental state. Were it not for the constant activity of photosynthetic plants and a minor amount of photo dissociation in the upper atmosphere, we would not have an oxygen-bearing atmosphere and we would not be here. Equally, the most important compound of oxygen is water, without which life (in the sense that we know it) could not exist. The role of water in virtually all geologic processes is profound, from formation of ore deposits to igneous petrogenesis to metamorphism to erosion and sedimentation. In planetary science, oxygen has a dual importance. First and foremost is its critical role in so many fundamental Solar System processes. The very nature of the terrestrial planets in our own Solar System would be much different had the oxygen to carbon ratio in the early solar nebula been somewhat lower than it was, because elements such as calcium and iron and 1 2

Abundance from Lodders (2003) From Böhlke et al. (2005)

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titanium would have been locked up during condensation as carbides, sulfides and nitrides and even (in the case of silicon) partly as metals rather than silicates and oxides. Equally, the role of water ice in the evolution of our Solar System is important in the early accretion and growth of the giant planets and especially Jupiter, which exerted a major control over how most of the other planets formed. On a smaller scale, oxygen plays a critical role in the diverse kinds of physical evolution of large rocky planets, because the internal oxidation state strongly influences the formation and evolution of the core, mantle and crust of differentiated planets such as the Earth. Consider that basaltic volcanism may be a nearly universal phenomenon among the evolved terrestrial planets, yet there are basalts and basalts. The basalts of Earth (mostly), Earth’s Moon, Vesta (as represented by the HED meteorites) and Mars are all broadly tholeiitic and yet very different from one another, and one of the primary differences is in their relative oxidation states (for that matter, consider the differences between tholeiitic and calc-alkaline magma series on Earth). But there is another way that oxygen has proven to be hugely important in planetary science, and that is as a critical scientific clue to processes and conditions and even sources of materials. Understanding the formation and evolution of our Solar System involves reconstructing processes and events that occurred more than 4.5 Ga ago, and for which the only contemporary examples are occurring hundreds of light years away. It is a detective story in which most of the clues come from the laboratory analysis of the products of those ancient processes and events, especially those that have been preserved nearly unchanged since their formation at the Solar System’s birth: meteorites; comets; and interplanetary dust particles. For example, the oxidation state of diverse early Solar System materials ranges from highly oxidized (ferric iron) to so reducing that some silicon exists in the metallic state and refractory lithophile elements such as calcium exist occur in sulfides rather than in silicates or carbonates. These variations reflect highly different environments that existed in different places and at different times. Even more crucial has been the use of oxygen 3-isotope variations, which began almost accidentally in 1973 with an attempt to do oxygen isotope thermometry on high-temperature solar nebula grains (Ca-, Al-rich inclusions) but ended with the remarkable discovery (see Clayton 2008) of non-mass-dependent oxygen isotope variations in hightemperature materials from the earliest Solar System. The presolar nebula was found to be very heterogeneous in its isotopic composition, and virtually every different planet and asteroid for which we have samples has a unique oxygen-isotopic fingerprint. The idea for this book originated with Jim Papike, who suggested the idea of a study initiative (and, ultimately, a published volume) focused on the element that is so critically important in so many ways to planetary science. He recognized that oxygen is such a constant theme through all aspects of planetary science that the proposed initiative would serve to bring together scientists from a wide range of disciplines for the kind of cross-cutting dialogue that occurs all too rarely these days. In this sense the Oxygen Initiative is modeled on the Basaltic Volcanism Study Project, which culminated in what remains to this day a hugely important reference volume (Basaltic Volcanism Study Project 1981). After obtaining community input and feedback, primarily through the Curation and Analysis Planning Team for Extraterrestrial Materials (CAPTEM) and the Management Operations Working Group for NASA’s Cosmochemistry Program, a team of scientists was assembled who would serve as chapter writing leads, and the initiative was formally proposed to and accepted by the Lunar and Planetary Institute (LPI; Dr. Stephen Mackwell, Director) for sponsorship. A formal proposal was then submitted to and approved by the Mineralogical Society of America to publish the resulting volume in the Reviews in Mineralogy and Geochemistry (RiMG) series. Three open workshops were held as preludes to the book: Oxygen in the Terrestrial Planets, held in Santa Fe, NM July 20-23, 2004; Oxygen in Asteroids and Meteorites, held in Flagstaff, AZ June 2-3, 2005; and Oxygen in Earliest Solar System Materials and Processes (and including the outer planets and comets), held in Gatlinburg, TN September 19-22, 2005. The workshops were each organized around a small number of sessions (typically 4-6), each focusing on a

Introduction

3

particular topic and consisting of invited talks, shorter contributed talks, and ample time for discussion after each talk. In all of the meetings, the extended discussion periods were lively and animated, often bubbling over into the breaks and later social events. As a consequence of the cross-cutting approach, the final book spans a wide range of fields relating to oxygen, from the stellar nucleosynthesis of oxygen, to its occurrence in the interstellar medium, to the oxidation and isotopic record preserved in 4.56 Ga grains formed at the Solar System’s birth, to its abundance and speciation in planets large and small, to its role in the petrologic and physical evolution of the terrestrial planets. Thanks are due to many people and organizations, without whose help and support neither the workshops nor this volume would have happened. Dr. Steve Mackwell immediately recognized the importance of the initiative and pledged LPI logistical and financial support for the workshops and publication. In particular, Sue McCown and Kimberly Taylor of the LPI provided terrific logistical and on-site support for the three workshops. Financial support for the workshops and this book was also provided by NASA’s Cosmochemistry Program and its Discipline Scientist, Dr. David Lindstrom. I particularly wish to acknowledge the invaluable contributions my co-editors, Dave (Duck) Mittlefehldt and John Jones, and of the associate editors who handled papers coming out of the third workshop on Earliest Solar System Materials and processes: Drs. Andy Davis, Sasha Krot, Larry Nittler, Ed Scott, Sara Russell, and Ed Young. Dr. Steve Simon undertook what probably was the most Herculean task of all, acting as technical editor for all of the book chapters to ensure that journal style and internal consistency were maintained. He did this efficiently, cheerfully, and well. RiMG Series Editor Jodi Rosso patiently and expertly turned the edited chapters into final camera-ready copy. Finally, in the background but always present, was Jim Papike himself keeping us all on target and on schedule (and kicking some posteriors where posterior-kicking was needed). This volume is a testimony to Jim’s vision, and it would not have happened without him. Basaltic Volcanism Study Project (1981) Basaltic Volcanism on the Terrestrial Planets, Pergamon Press, New York Böhlke JK, de Laeter JR, De Bièvre P, Hidaka H, Peiser HS, Rosman KJR, Taylor PDP (2005) Isotopic compositions of the elements, 2001. J Phys Chem Ref Data 34:57-67 Clayton RN (2008) Oxygen isotopes in the early Solar System — A historical perspective. Rev Mineral Geochem 68:5-14 Lodders K (2003) Solar system abundances and condensation temperatures of the elements. Astrophys J 591:1220–1247

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 5-14, 2008 Copyright © Mineralogical Society of America

Oxygen Isotopes in the Early Solar System — A Historical Perspective Robert N. Clayton Enrico Fermi Institute, Department of Chemistry, Department of the Geophysical Sciences University of Chicago Chicago, Illinois 60637, U.S.A. [email protected]

ABSTRACT The first suggestion for the use of oxygen isotopes in cosmochemistry was that of H. C. Urey and colleagues in 1934, but appropriate instrumentation had not yet been developed. The modern era of oxygen isotope cosmochemistry began with the study of Apollo lunar samples in 1969 and of Allende refractory inclusions in 1973. The large (>5%) variations in 17O/16O and 18O/16O ratios, and small variations in 17O/18O were first interpreted as nucleosynthetic effects, but are now recognized to be the result of chemical processes early in Solar System history. Thus oxygen isotopes provide natural tracers for processes of formation of solid bodies in the inner Solar System. In particular, oxygen isotopes are very useful in recognizing genetic associations among meteorite groups. They also have been valuable in the study of parent body processes, such as metamorphism and aqueous alteration. There is conjecture that the ultimate cause of the oxygen isotope effects may be isotope-selective photodissociation of CO, which will be tested by isotopic measurement of solar oxygen and nitrogen collected in the NASA Genesis mission.

BEFORE ALLENDE Measurements of the oxygen isotopic compositions of extraterrestrial materials have provided unique insights into processes of formation of our Solar System (Clayton 2003). Since the three stable isotopes of oxygen (16O, 17O, 18O) are synthesized by different nuclear processes in different astrophysical sites, any incomplete homogenization in the interstellar medium could result in variations in the relative abundances of the isotopes in samples with diverse origins. In a remarkably prescient paper, Manian et al. (1934) attempted to observe such variations by measurement of 18O/16O ratios in meteorites that may have originated beyond the Solar System, as judged from their apparent hyperbolic orbits. Their search failed for two major reasons: (1) the absence of any “extra-solar rocks” in the size range of meteorites; and (2) the inadequacy of mass spectrometers at that time. In fact, the mass spectrum shown by Manian et al. has such low resolution that it does not even show the existence of 17O, quantitative determination of which is essential for recognition of extra-solar oxygen. Analogous measurements can be carried out today with modern mass spectrometers on micrometer-sized grains of oxides and silicates, found within meteorites, and do indeed yield variations in 18O/16O and 17O/16O by factors of ten or more (Nittler et al. 1994), which are interpreted as residual heterogeneities of nucleosynthetic origin (to be discussed below). It was anticipated by Brown (1947) that meteorites might contain daughter products of radionuclides with half-lives in the range 103–109 yr, yielding measurable quantities of the daughter isotopes in some phases. Reynolds (1960) found excess 129Xe from decay of the 1529-6466/08/0068-0002$05.00

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extinct radionuclide 129I (t1/2 = 17 Ma) in the Richardton chondrite, thus indicating a short time-interval between a nucleosynthetic process and formation of the Solar System. This observation was followed by the identification by Black (1972) of neon-E, a component within carbonaceous chondrites, consisting of almost pure 22Ne. Black stated: “… ‘E’ is not indigenous to the solar mix, but rather is to be assigned an extra-solar system origin”. Neon-E may be produced directly in a supernova explosion or indirectly as a decay product of 22Na (t1/2 = 2.6 yr) (Nichols et al. 1994). In light of the background information described above, it would be expected that a renewed search for “isotopic anomalies” in oxygen, measuring all three stable isotopes, would have been initiated. Such was not the case: oxygen isotope studies of meteorites in the 1960s and early 1970s continued to be based only on 18O/16O variations, interpreted as mass-dependent fractionation effects (Taylor et al. 1965; Onuma et al. 1972a). Mass-dependent isotopic fractionation in light elements, such as hydrogen, carbon, nitrogen, oxygen, and sulfur, was put on a firm theoretical foundation by the classic work of Urey (1947). In his tabulations of vibrational frequencies and isotopic partition functions for oxygen-bearing compounds, Urey considered only 16O and 18O, tacitly implying that information from 17O was redundant and unnecessary. Almost all lightelement isotope effects in terrestrial geochemistry follow the principles described by Urey, and, to this day, almost all terrestrial oxygen isotope studies are based only on 18O/16O variations. In some extraterrestrial applications, such as the study of oxygen isotope variations in lunar rocks, the approach based on terrestrial experience has worked well (Onuma et al. 1970; Taylor and Epstein 1970), although there are some clear exceptions, such as the virtual absence of deuterium in the hydrogen of implanted solar wind (Epstein and Taylor 1970). It was later observed that the 15N/14N ratio of solar wind nitrogen implanted in lunar soils varied by more than 30%, a range larger than can be explained by known mass-dependent effects (Thiemens and Clayton 1980), which could imply some specific solar process.

AFTER ALLENDE In 1969, laboratories around the world were gearing up for study of the first lunar samples, due in autumn of that year. On February 8, 1969, the remarkable CV3 chondrite, Allende, fell in Chihuahua State, Mexico, providing more than two tons of primitive meteorite for detailed chemical, mineralogical, and petrographic study. The fall was perfectly timed to contribute to L. Grossman’s doctoral research at Yale University (Grossman 1972). Of special interest to him were the refractory, calcium-aluminum-rich inclusions (CAIs; Fig. 1), which were interpreted as being primary condensates from a hot solar gas. This suggested a novel application of oxygenisotope thermometry: a determination of the condensation temperature and the composition of the nebular gas (Onuma et al. 1972b) by detailed isotopic measurements of the individual minerals that make up the CAIs, predominantly spinel, pyroxene, melilite, plagioclase and occasionally olivine. As was standard practice at the time, mass spectrometry was done with CO2 as the sample gas, with mass 46/mass 44 used for measurement of 18O/16O, and mass 45/ mass 44 for measurement of 13C/12C. In the chemical preparation of CO2, oxygen was liberated from the sample as O2, by reaction with BrF5, and was reacted with hot graphite to make CO2. Hence, the carbon isotope ratio should be constant for all samples. However, given the huge variations found for 18O/16O in CAI minerals, we also found a linear correlation between 46/44 variations and 45/44 variations, which we recognized were due to variations in 17O/16O, since 17 O also contributes a small amount to the ion beam at m/e = 45 (16O12C17O). Such a correlation is expected in mass-dependent fractionations (Craig 1957), but our observations gave a slope that was twice the Craig value, showing that the magnitude of the 17O/16O variations was equal to that of 18O/16O variations, rather than one-half of the 18O/16O variations, as is known for mass-dependent fractionation. These observations were the basis for the first paper on “oxygen

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Figure 1. A type B2 CAI from Allende, showing the roughly spherical shape and igneous texture indicative of crystallization from a melt. In this false-color BSE image, dark grey is melilite, light grey is clinopyroxene, medium grey is anorthite and black is spinel (MacPherson and Grossman 1981). In a very similar Allende inclusion, Al3S4, melilite has δ18O and δ17O near zero, whereas pyroxene and spinel have δ18O and δ17O near −40‰, even for spinels entirely enclosed within melilite (Clayton et al. 1977).

isotope anomalies” (Clayton et al. 1973). Subsequent oxygen isotope analyses have been done by measuring O2+, using O2 as the sample gas in gas-source spectrometers, or by measuring O– ions by Secondary Ion Mass Spectrometry (SIMS). In the 1970’s, it was shown that the dominant oxygen isotopic pattern in primitive extraterrestrial materials at the sub-centimeter scale (chondrules and CAIs) is a linear array with slope near 1 on a “three-isotope” graph of δ17O versus δ18O (e.g., Clayton et al. 1977). Such an array corresponds to a nearly constant 17O/18O ratio, with variable amounts of 16O (Fig. 2). This could be a mixing line between two end-members, one enriched in 16O, the other depleted in 16O. The simplest explanation for the meteoritic mixing line, or CCAM (carbonaceous chondrite anhydrous minerals), was the addition of a component enriched in 16O by stellar nucleosynthesis. In the limit of pure “16O-nuggets”, these would have to constitute about 5% of those mineral fractions that were the richest in 16O, and thus should be readily identified. This extreme case was clearly not supported by observation, and was replaced by the postulate of a component that was enriched by only a few percent in 16O, but still reflecting a nucleosynthetic origin (Clayton et al. 1977). One of the earliest observations of the oxygen isotope distribution in Allende CAIs was the recognition of a mineralogical control: the greatest 16O-enrichments occurred in spinel and clinopyroxene, and the smallest enrichments in melilite, plagioclase, and secondary garnet and feldspathoids (Clayton et al. 1977). This was attributed to a two-stage process, with 16O-rich minerals retaining their primary isotopic compositions, and 16O-poor phases reflecting some kind of secondary exchange, probably with a nebular gas.

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17

GO (‰ rel. SMOW)

0

CAI minerals Dark inclusions TF

-10

-20 AM CC

-30

-40

-50 -50

-40

-30

-20

-10

0

10

18

GO (‰ rel. SMOW) Figure 2. An oxygen three-isotope plot showing the CCAM (carbonaceous chondrite anhydrous mineral) line defined by minerals from Allende CAIs, and, for reference, the terrestrial fractionation line (TF) defined by terrestrial rocks, minerals and waters. In early interpretations of the CCAM line, the compositions at the upper end were considered “normal” and others were called “anomalous.” In the self-shielding interpretation, the solar composition lies at the lower end of the CCAM line, and other compositions reflect exchange with a photochemically processed gas.

One obvious test for a nucleosynthetic origin of 16O excesses is a search for correlated nuclear effects in other low-mass elements with which oxygen is chemically bound, such as magnesium and silicon. This test was performed on a set of typical Allende CAIs (Mittlefehldt et al. 2008), in which the fluorination reaction on each CAI yielded simultaneously O2 and SiF4, for oxygen and silicon isotope measurements. The oxygen data all followed the slope-1 CCAM trend, and the silicon data (δ29Si vs. δ30Si) all followed a slope-1/2 mass-dependent fractionation trend. The oxygen and silicon isotope data were uncorrelated, showing that two separate processes were involved, and in particular, casting serious doubt on a nucleosynthetic origin for the oxygen isotope effects.

FUN CAIs Early in the isotopic studies of Allende CAIs, two inclusions were found in which unusually large mass-dependent heavy-isotope enrichment was observed in magnesium (Wasserburg et al. 1977) and silicon (Clayton et al. 1978). These CAIs, unremarkable chemically and mineralogically, also have apparent nucleosynthetic isotope anomalies in all elements that have been studied (e.g. McCulloch and Wasserburg 1978; Lee et al. 1978). They were dubbed “FUN” inclusions, for “fractionated” and “unidentified nuclear” effects (Wasserburg et al. 1977). Their oxygen isotopic patterns differ from those of most CAIs, in that an intermediate stage of mass-dependent evaporation appears to have occurred between the time of primary crystallization and the secondary isotope exchange with a more 16O-poor reservoir (Clayton and Mayeda 1977; Lee et al. 1980; Clayton et al. 1984). The connection between an evaporation

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event and the presence of nuclear anomalies remains unknown. The FUN inclusions are also devoid of radiogenic 26Mg (MacPherson 2003). There are about six known FUN CAIs, all but one of which were known by 1984; only one, Vigarano 1623-5, has been discovered in the last 20 years (Davis et al. 1991), in spite of the isotopic analysis of hundreds of CAIs.

OXYGEN ISOTOPES IN PRESOLAR GRAINS The terms “presolar grains” and “stardust” refer to mineral grains that acquired their chemical and isotopic characteristics before the Sun and Solar System formed, and were subsequently incorporated into meteorite parent bodies (Anders and Zinner 1993). They are recognizable by their large differences in isotopic compositions in most elements with respect to Solar System materials. In most cases, the abundances of presolar grains in meteorites are at the level of tens of ppm, so their influence on the bulk isotopic compositions of meteorites is very small. Thus, although extremely interesting in their own right, they are not responsible for the isotopic variations discussed above. “Anomalous” isotopic variations in the noble gases due to presolar grains are indeed measurable in bulk meteorites, however; this is the property that led to the discovery of presolar grains (Lewis et al. 1987). Presolar grains condensed in cool atmospheres of earlier generations of stars, and their notable isotopic abundances result from nuclear processes in the parent stars or in previous generations of stars (galactic chemical evolution). The chemical separation techniques used by the Anders group favor isolation of carbon-rich phases, notably graphite, diamond, and silicon carbide (Anders and Zinner 1993), most of which were derived from carbon-rich stars of the Asymptotic Giant Branch (AGB). Subsequently, search techniques involving oxygen-isotope mapping have revealed presolar oxides (spinel, corundum) (Hutcheon et al. 1994) and silicates (Nguyen and Zinner 2004). The oxygen-rich grains have been assigned to four classes according to their oxygen isotopic compositions on a three-isotope plot (Nittler et al. 1997). Grains in all classes may have been derived from oxygen-rich red giant stars at various stages of stellar evolution. Grains enriched only in 16O, once believed to be present in meteoritic CAIs, are exceedingly rare (Nittler et al. 1998).

CHEMICAL ISOTOPE EFFECTS In 1983, a new process was discovered that could reproduce the slope-1 array for oxygen isotopes by a purely chemical means, generically known as a non-mass-dependent (NMD) isotopic fractionation (Thiemens and Heidenreich 1983). This phenomenon was first recognized in gas phase laboratory synthesis of ozone (O3) from oxygen (O2), with heavyisotope enrichment in the product ozone. The authors interpreted this result as being due to isotope-selective photodissociation of O2. Thiemens and Heidenreich (1983) included CO in a list of molecules that might exhibit this isotopic self-shielding effect (see below), and Navon and Wasserburg (1985) discussed the advantages of CO over O2 for the origin of oxygen isotope variations in meteorites. Since neither O2 nor O3 is expected to have been a major constituent of the hydrogen-rich solar nebula, it has been important to determine the underlying cause of the NMD chemical effect, to assess its generality and applicability to other molecular species. Although a number of systems have shown departures from classical mass-dependent isotope effects, only the O2-O3 system has produced an extended slope-1 relationship. A theoretical interpretation of the NMD effect, based on symmetry-dependent kinetics, has been presented (Gao and Marcus 2002), and Marcus (2004) has attempted to extend these ideas to include the formation of 16Oenriched minerals, as are seen in CAIs. Proposals to account for NMD effects in meteorites by such chemical processes have assumed that the oxygen isotopic composition of the nebular gas, and of the Sun, was similar to the terrestrial composition, and that CAIs are “anomalous”

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and require a special mechanism for their formation. In the self-shielding model, the solar composition is taken to be that of the CAIs, so that the terrestrial (and inner Solar System) composition is “anomalous” and requires a special mechanism. Determination of the isotopic composition of the solar wind, to be obtained by analysis of samples returned by the Genesis mission and assumed to be representative of the isotopic composition of the Sun, should allow us to distinguish between these proposed scenarios.

PHOTOCHEMICAL EFFECTS Isotope effects associated with the phenomenon of photochemical self-shielding have been discussed in the astrophysical literature since the early 1980s (Bally and Langer 1982; van Dishoeck and Black 1988). The isotope effects of self-shielding are based on the photodissociation of gaseous carbon monoxide, the most abundant oxygen-containing molecule in the Galaxy. This process occurs by predissociation, in which absorption of an ultraviolet photon (λ = 90-100 nm) excites a molecule to a short-lived state which then dissociates into ground-state C and O atoms. The excitation occurs in narrow lines that are wavelengthspecific for the various isotopologues of CO, due to the mass-dependence of their vibrational energies. Because of its much greater abundance, the 12C16O absorption line becomes optically thick (saturated absorption), while the lines of the rare isotopic species remain optically thin (unsaturated absorption). Thus the interior of a molecular cloud or solar nebula undergoes preferential dissociation of 13C16O, 12C17O and 12C18O relative to 12C16O, enhancing the local abundances of atomic 17O and 18O, which can then be incorporated into other molecules, such as H2O, and eventually into mineral grains. The photochemical process has been modelled quantitatively for molecular clouds (Warin et al. 1996) and the predicted deficits of molecular 12 17 C O and 12C18O have been observed by ultraviolet spectroscopy (Sheffer et al. 2002). The self-shielding effect predicts that all phases produced from the gas enriched in the heavier, rare isotopes will themselves be enriched in these isotopes. As a consequence, the isotopic composition of the initial bulk cloud or nebula must be at least as 16O-rich as the 16O-richest minerals. For the Solar System, this implies that the Sun should have an oxygen isotopic composition similar to that of the spinels in CAIs, i.e., at the lower end of the mixing line in Figure 2 (Clayton 2002). Analysis of samples returned by NASA’s Genesis mission will provide an oxygen isotopic analysis of the solar wind, thus providing a direct test of the selfshielding proposal. Of the three classes of proposals for the origin of the meteoritic oxygen isotope variation: nucleosynthetic; NMD chemical; and photochemical, none has been disproved, and each has adherents. The nucleosynthetic scenario led to predictions of correlated nucleosynthetic effects in magnesium and silicon, and to expectations of very 16O-rich presolar grains. Neither has been found. The NMD chemical process clearly does occur in the Earth’s atmosphere, but a viable extension to solar nebular conditions has not yet been found. The self-shielding mechanism has been demonstrated to occur in cold, low-density molecular clouds, but also has not yet led to a comprehensive model for production of isotopically “anomalous” solids in the Solar System.

INTERNAL ASTEROIDAL PROCESSES Both the chemical NMD processes and the photochemical processes inherently involve the gas phase in the solar nebula, with different degrees of interaction resulting in final solid products (chondrules, CAIs, planetesimals) with differing 16O excesses or deficits, expressed in terms of Δ17O = δ17O − 0.52 δ18O. Thus an asteroid or rocky planet has a characteristic value of Δ17O, inherited from the nebula, but remaining constant for subsequent internal processes,

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11

such as metamorphism or melting. For example, for Earth, Δ17O = 0 (by definition), but for Mars (SNC meteorites) Δ17O = +0.32 (Franchi et al. 1999); for the HED (howardite-eucritediogenite) parent body Δ17O = −0.22 (Wiechert et al. 2004); for equilibrated H-chondrites Δ17O = +0.73; and for equilibrated L-chondrites Δ17O = +1.07 (Clayton et al. 1991). Igneous and metamorphic processes within a parent body occur at constant Δ17O, following massdependent fractionation patterns, just as they do on Earth. Inter-mineral isotopic fractionations can, therefore, be used for isotopic thermometry, as was done for metamorphosed chondrites (Onuma et al. 1972a) and for lunar igneous rocks (Onuma et al. 1970). The use of oxygen isotope abundances to trace aqueous alteration processes in carbonaceous chondrites illustrates the additional power of the three-isotope system. It is likely that water accreted to the carbonaceous chondrites as ice, which almost certainly had Δ17O different from that of the anhydrous silicates and other minerals. For example, in Murchison (CM2), Δ17O for anhydrous minerals is −5.2‰, whereas Δ17O for phyllosilicates is −1.9‰, implying interaction with an aqueous phase with Δ17O more positive than −1.9‰. The data can be interpreted either in terms of simple closed-system mass balance (Clayton and Mayeda 1999) or in terms of a more complicated model including fluid flow (Young et al. 1999). These models assume the accretion of water as ice, which reacts with olivine and pyroxene upon melting. The aqueous alteration conditions are more akin to those in terrestrial sea-floor weathering than to terrestrial hydrothermal processes. An important advance in isotopic thermometry of carbonates has been made by John Eiler and colleagues (Ghosh et al. 2006), which allows estimation of precipitation temperatures without independent knowledge of the oxygen isotopic composition of the fluid phase. Application of this technique to carbonates in carbonaceous chondrites yields temperatures from 0 to 40 °C (Guo et al. 2007).

NITROGEN Another long-standing problem in Solar System isotope studies is the very large range of variations observed in the nitrogen isotopic composition. Because of the close similarity between the N2 and CO molecules (isoelectronic, isobaric), solution of one isotope problem (oxygen or nitrogen) may aid in solution of the other. Of course, nitrogen has only two stable isotopes, so tests for mass-dependence cannot be made, and the only observable variable is the 15 N/14N ratio, or δ15N. However, nitrogen from meteorites exhibits a range of more than a factor of two in 15N/14N, strongly suggesting that some process beyond ordinary mass-dependent fractionation has occurred. Furthermore, the observed isotopic variations, although systematic from one meteorite group to another, have not shown correlations with any other property. In many ways, the history of the development of nitrogen isotope studies of the Solar System parallels that for oxygen isotopes. “Mainstream” presolar SiC grains are enriched in 14N by a factor of 10 or more, due to operation of the CNO nuclear cycle in their source stars (Anders and Zinner 1993), but concentrations of these grains in meteorites are too low to account for the observed whole-rock isotopic variations. The large variation in 15N/14N in nitrogen implanted in lunar soils was first interpreted as evidence for a secular increase in the ratio due to nuclear reactions near the surface of the Sun (Kerridge 1975). Geiss and Bochsler (1982) showed that such nuclear processes were quantitatively inadequate to account for the isotopic variations in implanted lunar nitrogen. Models in the late 1980s assumed two implanted components, of either lunar or solar origin (Kerridge 1989). Hashizume et al. (2000) used ion microprobe techniques on lunar ilmenite grains to distinguish implanted nitrogen of solar origin (correlated with deuterium-free hydrogen) from implanted nitrogen of non-solar origin, which they labelled “planetary,” without identifying a specific origin. They concluded that lunar soil contains only one solar component, with δ15N in the range −250 to

12

Clayton

−300‰. A composition of about −300‰ has also been inferred for nitrogen in ammonia in Jupiter’s atmosphere (Owen et al. 2001). This interpretation leaves unanswered the question of the origin, presumably within the Solar System, of the huge range of nitrogen isotopic compositions among various reservoirs. Owen et al. (2001) and especially Terzieva and Herbst (2000) considered the possibility of large chemical isotopic fractionations among nitrogen compounds in cold clouds. The latter authors concluded that such processes could not account for 15N enhancements by factors greater than 1.5, as has been seen in one interplanetary dust particle (Messenger et al. 1996), whereas the observed meteoritic enhancement extends to a factor of 3 (Prombo and Clayton 1985).

CONCLUSIONS If we accept the generalization by Suess (1965) that the initial solar nebula was isotopically well-homogenized, then we must find late-stage processes, operating within the Solar System, to account for the observed large isotopic differences in oxygen and nitrogen between the Sun and the inner Solar System. The process of isotopic self-shielding in the photodissociation of CO and N2 may satisfy this requirement; the molecules have very similar dissociation energies and both have an abundant lighter isotope, and one or two rare heavier isotopes. The apparent similarity in nitrogen isotope ratios between the Sun and Jupiter would imply that the photochemical isotopic effects were limited to matter now found in the inner Solar System, effectively ruling out processes that affected the entire solar nebula (Yurimoto and Kuramoto 2004; Lyons and Young 2005). Measurements of the solar wind isotopic compositions of both elements in samples returned by NASA’s Genesis mission will go a long way toward testing these hypotheses.

ACKNOWLEDGMENT This research has been supported over many years by grants from the U.S. National Science Foundation and from NASA. The most recent support is from NASA grant NAG513165. G.J. MacPherson and M.H. Thiemens provided helpful reviews.

REFERENCES Anders E, Zinner E (1993) Interstellar grains in primitive meteorites: diamond, silicon carbide, and graphite. Meteoritics 28:490-514 Bally J, Langer WD (1982) Isotope-selective photodestruction of carbon monoxide. Astrophys J 255:143-148 Black DC (1972) Trapped helium, neon and argon isotopic variations in meteorites — II Carbonaceous meteorites. Geochim Cosmochim Acta 36:377-394 Brown H (1947) An experimental method for the estimation of the age of the elements. Phys Rev 72:348 Clayton RN (2002) Self-shielding in the solar nebula. Nature 415:860-861 Clayton RN (2003) Oxygen isotopes in the solar system. Space Sci Rev 106:19-32 Clayton RN, Grossman L, Mayeda TK (1973) A component of primitive nuclear composition in carbonaceous meteorites. Science 182:485-498 Clayton RN, MacPherson GJ, Hutcheon ID, Davis AM, Grossman L, Mayeda TK, Molini-Velsko C, Allen JM (1984) Two forsterite-bearing FUN inclusions in the Allende meteorite. Geochim Cosmochim Acta 48:535–548 Clayton RN, Mayeda TK (1977) Correlated oxygen and magnesium isotopic anomalies in Allende inclusions: I. Oxygen. Geophys Res Lett 4:295–298 Clayton RN, Mayeda TK (1999) Oxygen isotope studies of carbonaceous chondrites. Geochim Cosmochim Acta 63:2089–2104 Clayton RN, Mayeda TK, Epstein S (1978) Isotopic fractionation of silicon in Allende inclusions. Proc Lunar Planet Sci Conf 9th: 1267–1278 Clayton RN, Mayeda TK, Goswami JN, Olsen EJ (1991) Oxygen isotope studies of ordinary chondrites. Geochim Cosmochim Acta 55:2317–2337

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Clayton RN, Onuma N, Grossman L, Mayeda TK (1977) Distribution of the presolar component in Allende and other carbonaceous chondrites. Earth Planet Sci Lett 34:209-224 Craig H (1957) Isotopic standards for carbon and oxygen and correction factors for mass-spectrometric analysis of carbon dioxide. Geochim Cosmochim Acta 12:133-149 Davis AM, Clayton RN, Mayeda TK, Sylvester PJ, Grossman L, Hinton RW, Laughlin JR (1991) Melt solidification and late-stage evaporation of a FUN inclusion from the Vigarano C3V chondrite. Geochim Cosmochim Acta 55:621-637 Epstein S, Taylor HP Jr. (1970) The concentration and isotopic composition of hydrogen, carbon and silicon in Apollo 11 lunar rocks and minerals. Proc Apollo 11 Lunar Sci Conf, 1085-1096 Franchi IA, Wright IP, Sexton AS, Pillinger CT (1999) The oxygen isotopic composition of Earth and Mars. Meteorit Planet Sci 34:657–661 Gao YO, Marcus RA (2002) On the theory of the strange and unconventional isotope effects in ozone formation. J Chem Phys 116:137-154 Geiss J, Bochsler P (1982) Nitrogen isotopes in the solar system. Geochim Cosmochim Acta 46:529-548 Ghosh P, Adkins J, Affeck H, Balta B, Guo W, Schauble EA, Schrag D, Eiler JM (2006) 13C-18O bonds in carbonate minerals: a new kind of paleothermometer. Geochim Cosmochim Acta 70:1439–1450 Grossman L (1972) Condensation, chondrites and planets. Ph.D. Dissertation, Yale University, New Haven, Connecticut Guo W, Perronnet M, Zolensky ME, Eiler JM (2007) Temperatures of aqueous alteration on carbonaceous chondrite parent bodies. Meteorit Planet Sci 42:A61 Hashizume K, Chaussidon M, Marty B, Robert F (2000) Solar wind record on the Moon: deciphering presolar from planetary nitrogen. Science 290:1142-1145 Hutcheon ID, Huss GB, Fahey AJ, Wasserburg GJ (1994) Extreme 26Mg and 17O enrichments in an Orgueil corundum: identification of a presolar oxide grain. Astrophys J 425:L97–L100 Kerridge JF (1975) Solar nitrogen: evidence for a secular increase in the ratio of nitrogen-15 to nitrogen-14. Science 188:162-164 Kerridge JF (1989) What has caused the secular increase in solar nitrogen-15? Science 245:480-486 Lee T, Mayeda TK, Clayton RN (1980) Oxygen isotopic anomalies in Allende inclusion HAL. Geophys Res Lett 7:493-496 Lewis RS, Tang M, Wacker JF, Anders E, Steel E (1987) Interstellar diamonds in meteorites. Nature 326:160– 162 Lyons JR, Young ED (2005) CO self-shielding as the origin of oxygen isotope anomalies in the early solar nebula. Nature 435:317-320 MacPherson GJ (2003) Calcium-aluminum-rich inclusions in chondritic meteorites. In: Treatise on Geochemistry, vol. 1. Davis A (ed) Elsevier, Oxford UK, p 201–246 MacPherson GJ, Grossman L (1981) A once-molten, coarse-grained Ca-rich inclusion in Allende. Earth Planet Sci Lett 52:16-24 Manian SH, Urey HC, Bleakney W (1934) An investigation of the relative abundance of the oxygen isotopes O16:O18 in stone meteorites. J Am Chem Soc 56:2601-2609 Marcus RA (2004) Mass-independent isotope effect in the earliest processed solids in the solar system: A possible chemical mechanism. J Chem Phys 121:8201-8211 Messenger S, Keller LP, Thomas KL, Walker RM (1996) Nitrogen petrography in two 15N-rich IDPs. Meteorit Planet Sci 31:A88 Mittlefehldt DW, Clayton RN, Drake MJ, Righter K (2008) Oxygen isotopic composition and chemical correlations in meteorites and the terrestrial planets. Rev Mineral Geochem 68:399-428 Navon O, Wasserburg GJ (1985) Self-shielding in O2 — a possible explanation for oxygen isotopic anomalies in meteorites? Earth Planet Sci Lett 73:1–16 Nguyen A, Zinner E (2004) Discovery of ancient silicate stardust in a meteorite. Science 303:1496-1499 Nichols RH Jr., Kehm K, Brazzle R, Amari S, Hohenberg CM, Lewis RS (1994) Ne, C, N, O, Mg and Si isotopes in single interstellar graphite grains: Multiple stellar sources for Neon E (L). Meteoritics 29:510-511 Nittler LR, Alexander CMO’D, Gao X, Walker RM, Zinner EK (1994) Interstellar oxide grains from the Tieschitz ordinary chondrite. Nature 370:443-446 Nittler LR, Alexander CMO’D, Gao X, Walker RM, Zinner EK (1997) Stellar sapphires: the properties and origins of presolar Al2O3 in meteorites. Astrophys J 482:475–495 Nittler LR, Alexander CMO’D, Wang J (1998) Meteoritic oxide grain from supernova found. Nature 393:222 Onuma N, Clayton RN, Mayeda TK (1970) Oxygen isotope fractionation between minerals and an estimate of the temperatures of formation. Science 167:536–538 Onuma N, Clayton RN, Mayeda TK (1972a) Oxygen isotope temperatures of “equilibrated” ordinary chondrites. Geochim Cosmochim Acta 36:157-168 Onuma N, Clayton RN, Mayeda TK (1972b) Oxygen isotope cosmothermometer. Geochim Cosmochim Acta 36:169-188

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Owen T, Mahaffy PR, Niemann HB, Atreya S, Wong M (2001) Protosolar nitrogen. Astrophys J 553:L77–L79 Prombo CA, Clayton RN (1985) A striking isotope anomaly in the Bencubbin and Weatherford meteorites. Science 230:935-937 Reynolds JH (1960) Determination of the age of the elements. Phys Rev Lett 4:8-10 Sheffer Y, Lambert DL, Federman SR (2002) Ultraviolet detection of interstellar 12C17O and the CO isotopomeric ratios toward X-Persei. Astrophys J 574:L171-L174 Suess HE (1965) Chemical evidence bearing on the origin of the solar system. Ann Rev Astron Astrophys 3:217-234 Taylor HP Jr., Epstein S (1970) O18/O16 ratios of Apollo 11 lunar rocks and minerals. Proc Apollo 11 Lunar Sci Conf, 1613-1626 Taylor HP Jr., Duke MB, Silver LT, Epstein S (1965) Oxygen isotope studies of minerals in stony meteorites. Geochim Cosmochim Acta 29:489-512 Terzieva R, Herbst E (2000) The possibility of nitrogen isotopic fractionation in interstellar clouds. Mon Not Royal Astron Soc 317:563-568 Thiemens MH, Clayton RN (1980) Ancient solar wind in lunar microbreccias. Earth Planet Sci Lett 47:34-42 Thiemens MH, Heidenreich JE III (1983) The mass-independent fractionation of oxygen — a novel isotope effect and its possible cosmochemical implications. Science 219:1073-1075 Urey HC (1947) The thermodynamics of isotopic substances. J Chem Soc (London), 562-581 van Dishoeck EF, Black JH (1988) The photodissociation and chemistry of interstellar CO. Astrophys J 334:771802 Warin S, Benayoun JJ, Viala YP (1996) Photodissociation and rotational excitation of interstellar CO. Astron Astrophys 308:533-564 Wasserburg GJ, Lee T, Papanastassiou DA (1977) Correlated O and Mg isotopic anomalies in Allende inclusions: II Magnesium. Geophys Res Lett 4:299–302 Wiechert UH, Halliday AN, Palme H, Rumble D (2004) Oxygen isotope evidence for rapid mixing of the HED meteorite parent body. Earth Planet Sci Lett 221:373–382 Young ED, Ash RD, England P, Rumble D (1999) Fluid flow in chondritic parent bodies: Deciphering the compositions of planetesimals. Science 286:1331–1335 Yurimoto H, Kuramoto K (2004) Molecular cloud origin for the oxygen isotope heterogeneity in the solar system. Science 305:1763-1766

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 15-30, 2008 Copyright © Mineralogical Society of America

Abundance, Notation, and Fractionation of Light Stable Isotopes Robert E. Criss Washington University St. Louis, Missouri 63130, U.S.A. [email protected]

James Farquhar University of Maryland College Park, Maryland 20742, U.S.A. [email protected]

ABSTRACT Stable isotopes have become an essential tool to characterize and understand terrestrial and extraterrestrial matter. This chapter will briefly review the abundances of important light stable isotopes, demonstrate the link between abundance and atomic weight, introduce the notations and diagrams that are commonly used to report isotopic measurements, describe and partially explain the types of fractionation effects known to occur in nature, and direct the reader to more comprehensive sources of information on each subject. The special techniques needed to make accurate isotopic measurements gave rise to special notation for reporting stable isotope data, and these notations in turn gave rise to special diagrams that emphasize compositional differences and facilitate interpretation. Fundamental definitions are the isotope ratio R, representing the ratio of the abundance of a heavy isotope to that of a lighter, typically much more common isotope, and the isotopic fractionation factor α, representing the quotient RA/RB of the isotope ratios of two substances A and B. Under equilibrium conditions, lnα can theoretically vary linearly with 1/T at low temperatures or with 1/T 2 at high temperatures, forming the basis for a standard graph. For practical reasons the ratio R is difficult to measure and inconvenient to report, so stable isotope abundances are usually reported as delta values (δ-values) that describe their deviations from a defined “standard” material. Thus, the most important diagram for data interpretation is the “δ-δ plot” where the δ-values of two coexisting phases are simply plotted against each other. In systems where two different heavy isotopes exist, two different delta values may be defined, each normalizing the abundance of one of the heavy isotopes to the common light isotope. In such cases, a very important diagram called the “three isotope” plot involves simply plotting these two different δ-values against each other for a given material, and the slopes of data arrays on such graphs can be used to distinguish ordinary “mass-dependent” fractionation (MDF) effects from “non-mass-dependent” fractionations (NMF). Numerous algebraic convolutions of the above definitions have been made, providing special definitions that can elucidate different phenomena. The processes that govern isotope distribution have become progressively better understood, yet recent studies show that these processes are more diverse than anticipated only ten years ago.

INTRODUCTION Isotopic variations in light elements, particularly hydrogen, carbon, nitrogen, oxygen and sulfur (HCNOS), provide key information on planetary formation and on the evolution and interactions of their lithospheres, atmospheres and hydrospheres. The atomic abundances of 1529-6466/08/0068-0003$05.00

DOI: 10.2138/rmg.2008.68.3

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Criss & Farquhar

the isotopes of interest are only a few percent or less, and their absolute concentrations are difficult to accurately measure given their typically small natural variations. Accordingly, a “delta” (δ) notation was devised long ago to report isotopic abundances as per mil (‰), or parts per thousand, deviations from defined isotopic standards, as this notation exploits comparisons that can be precisely made with Nier-type mass spectrometers modified for this purpose (McKinney et al. 1950). In addition, a notation termed “Big-delta” (Δ) is commonly used both to report the simple difference in δ-values between two phases, and a similar notation, termed “Cap-delta,” is used to report the deviation of samples from a defined reference line on a “three isotope plot,” so attention to detail is required. In systems having multiple phases or components, the isotopes may “fractionate,” or differ in relative abundance, among those parts. The isotopic fractionation factor is analogous to a distribution coefficient, and it may reflect either equilibrium or disequilibrium processes or both. Normal fractionation processes depend critically on the vibrational frequencies, and hence on isotopic masses, so these effects are called “mass-dependent fractionations,” or MDF. Normal equilibrium fractionations among simple gases may be calculated from statistical thermodynamics (Urey 1947; Richet et al. 1977), but estimates for complex substances and for kinetic processes are more difficult to make. More recently, several “non-mass-dependent fractionation” (NMF) processes have been found to occur among rarefied gas molecules, which in special cases can be transferred to other molecules and even to solid phases. Processes that can produce NMF are discussed in detail in the chapter by Young et al. (2008), and include photo-dissociation in cases where the symmetry or atmospheric optical opacity differs for a given molecule and its various, isotopically-substituted forms (“isotopologues,” Fig. 1). Isotopomers, molecules with the same isotopic composition, but with a different arrangement (e.g., the linear molecules 14N15N16O and 15N14N16O) are also subject to small differences in their chemical behavior and have become a focus of recent research. This chapter will briefly review the abundances of important light stable isotopes, demonstrate the link between isotopic abundances and elemental atomic weight, introduce the notations and diagrams that are commonly used to report isotopic measurements, describe and partially explain the types of fractionation effects known to occur in nature, and direct the reader to more comprehensive sources of information on each subject. Subsequent chapters in the volume will apply this framework to several different isotope systems in diverse natural settings.

ISOTOPIC ABUNDANCES AND ATOMIC WEIGHTS Each chemical element is an assemblage of different types of atoms, or nuclides, that share the same number “Z” of protons but can differ in the number “N” of neutrons. The mass number “A,” representing the simple sum of Z plus N, is an integer that is close to the actual, non-integral atomic weight of the nuclide of interest. Isotopes of different elements are signified by a standard notation that is the chemical symbol for the element preceded by a superscript indicating the mass number. For example, 12C or “carbon-12” is the most common carbon atom, 13C is a stable isotope called “carbon-13,” and 14C, the most important radioisotope, is commonly called “carbon-14” or “radiocarbon.” All of these carbon atoms have six protons, as carbon is the sixth element in the periodic table, but they respectively have 6, 7, and 8 neutrons, a difference that affects the mass as well as the nuclear character and stability of each nuclide. Atomic weights are reported in atomic mass units, abbreviated “amu” or sometimes simply “u”. The amu is defined as exactly 1/12 of the mass of the carbon-12 nuclide, itself defined to have a rest mass of 12.000000 amu. Atoms are tiny, so the amu is a small unit, only 1.66054 × 10−27 kg. Nuclide masses are routinely reported to great precision, because isotopes

Abundance, Notation, Fractionation of Light Stable Isotopes

17

STABLE ISOTOPOLOGUES OF WATER 16

1

H

17

O

16 2

1

O

H

18

O

17 2

1

O

H

18.0106 amu Ab= 997,300 ppm 19.0148 amu Ab=400 ppm P°=23.756 torr P°=23.641torr

16

17

O

1 2 16

16

2

16 2

2

20.0231 amu Ab= 0.022 ppm P°=20.54 torr

17 2

O

H H O

Ab= .12 ppm

21.0211 amu Ab= 0.6 ppm

18

O

H

O

O

1 2 18

H H O

17

O

H

18

O

20.0211 amu

18 2

20.0148 amu Ab= 2,000 ppm P°=23.535 torr

1 2 17

H H O

19.0168 amu Ab=300 ppm P°=22.01 torr

O

2

O

H

O -5

21.0274 amu Ab=10 ppm

18 2

O -5

22.0274 amu Ab.=5*10 ppm

Figure 1. The water molecule has nine stable isotologues, all with different molecular weights and relative abundances (Ab), as indicated. Physical properties also vary among these molecules, as demonstrated by the vapor pressures at 25 °C (P°) that are provided or estimated for the most important isotopologues.

of different elements can share identical mass numbers (e.g., 40K, 40Ar, 40Ca), but in detail their masses differ slightly. In addition, tiny mass differences among nuclides can translate into tremendous energy differences according to Einstein’s formula, E = mc2. The masses and relative abundances of the stable HCNOS isotopes are given in Table 1. In addition, the approximate atomic weight of each element, representing the weighted average of the constituent isotopes in an average sample, is given in bold. The latter is the number reported in an ordinary periodic table, and is simply the sum of the atomic abundance (Abi) of each isotope multiplied by its atomic mass (Wti), i.e.: Element Atomic Weight = ∑ AbiWt i

(1)

i

For example, the atomic weight of 12.011 amu for the element carbon is calculated as follows: Carbon Atomic Weight = 0.989×12.0000 + 0.011×13.00335 (2)

18

Criss & Farquhar Table 1. Stable isotopes of the HCNOS elements. Element

Atomic Weight (amu)

Isotope

Hydrogen (Z=1)

Abundance (atom %)

1.0079 1

1.007825

99.985

2

2.014102

0.015

H (Protium) H (D, or Deuterium)

Carbon (Z=6)

12.011 12

C

12.00000

98.90

13

C

13.00335

1.10

14

N

14.003074

99.63

15

N

15.000109

0.37

16

O

15.994915

99.76

17

O

16.999131

0.04

18

O

17.999160

0.20

32

S

31.9720705

95.02

33

S

32.9714583

0.75

34

S

33.9678665

4.21

36

S

35.9670808

0.02

Nitrogen (Z=7)

14.0067

Oxygen (Z=8)

15.9994

Sulfur (Z=16)

32.07

(Source: Walker et al. 1989)

Note that the atomic weights of the elements are not intrinsic, as they vary slightly from sample to sample depending upon the exact atomic proportions of the constituent isotopes. In contrast, the atomic weights of the individual isotopes are intrinsic and invariant.

NOTATION Isotope ratios The isotope ratio R, here representing the simple quotient of an isotope of interest normalized to the most abundant nuclide of a given element (e.g., 2H/1H; 13C/12C; 18O/16O, etc.) is the fundamental variable of stable isotope geochemistry. Note the convention that R depicts the quotient of the heavy isotope abundance over the light isotope abundance, which is invariably used for the HCNOS isotopes, and is recommended for stable isotope studies of many other elements that are now being studied (Johnson et al. 2004). The ratio R is the variable of choice for understanding the processes that control isotope distribution; moreover, many needless approximations are avoided if R is used as the starting point in mathematical derivations. However, for numerous reasons R is not a convenient means for reporting natural variations of most light stable isotopes. First, Table 1 shows that the most common nuclide of each of the HCNOS elements constitutes more than 95% of the constituent atoms, so R is a small number, typically between 0.05 and 0.0001. Second, in most natural samples, the variations among the stable isotopes are rather small, and are difficult to measure as absolute ratios.

Abundance, Notation, Fractionation of Light Stable Isotopes

19

δ-values For all of the above reasons, a special method was developed long ago to report variations in the abundances of light stable isotopes. This method cleverly exploits the fact that a mass spectrometer can be used to determine the difference between the isotope ratios of two substances much more precisely than the absolute ratio of each individual substance. In particular, McKinney et al. (1950) invented “δ-values” to report the abundance ratio of a measured sample, Rx, as the normalized deviation from the abundance ratio, Rstd, of a defined isotopic standard, such that: ⎛ R − R std ⎞ δ x = 1000 ⎜ x ⎟ ⎝ R std ⎠

(3)

Ordinarily, the delta symbol is followed not by a subscript x but rather by notation indicating the isotope ratio of interest, i.e., δD or δ2H for the 2H/1H ratio, δ13C for the 13C/12C ratio, etc. The factor of 1000 magnifies these small deviations into convenient numbers, called per mil (‰) differences, where 1‰ is one tenth of one percent. For most natural samples, the variations in isotope abundances of the HCNOS elements are a few per mil to tens of per mil, but larger variations occur, particularly for hydrogen or for certain extraterrestrial materials. Mass spectrometers routinely measure isotopic differences to better than ± 1‰ for H and better than ± 0.1‰ for CNOS. The δ notation requires comparison to a defined isotopic standard for each element of interest. Samples of many different substances have been used as isotopic standards over the years, especially for oxygen. O’Neil (1986a) provides a concise summary of the most important HCNOS isotopic standards and the means to interrelate them.

Isotopic fractionation factor The isotopic fractionation factor “α” represents the partitioning of isotopes between two phases, and is analogous to a geochemical partition coefficient. This factor is most commonly used to represent the theoretical equilibrium condition between the phases, but fractionation factors may be used to quantify a non-equilibrium condition or process, or simply used to represent the measured isotopic difference between the phases. The fractionation factor is directly defined as the simple quotient of the isotope ratios RA and RB of two coexisting phases A and B: αA-B = RA / RB

(4)

Equation (3) can be used to translate this definition in terms of δ values: αA-B = (1000 + δΑ) / (1000 + δΒ)

(5)

The value calculated for α is independent of the isotopic standard chosen to report the δ values of samples A and B in Equation (5), but of course both samples must be reported as per mil deviations from the same standard. Because the δ-values of most natural samples differ by only a few per mil or so, most isotopic fractionation factors are close to unity, typically between 0.95 and 1.10.

Big delta and related approximations The symbol Δ, called “Big delta,” is commonly used to represent the simple isotopic difference between two phases A and B: ΔA-B = δΑ − δΒ

(6)

Big delta is commonly called the “isotopic fractionation” between A and B, because of the following approximations that can be directly derived from Equation (5) in the typical case where α is close to unity:

20 or

Criss & Farquhar 1000 (αA-B −1) ≈ δΑ − δΒ

(7a)

1000 ln αA-B ≈ δΑ − δΒ

(7b)

Thus, the simple difference in the δ-values of two coexisting substances can be conveniently related to the effective fractionation factor to reasonably good accuracy, and rapid estimates of expected isotopic differences can be made for a given fractionation factor. Of course, use of Equation (5) will avoid any inaccuracies.

Capital delta The symbol ΔA, called “Capital delta” is used to describe the difference between the isotopic composition of a substance (A) and a reference “mass-dependent” fractionation line, usually the terrestrial fractionation line (see below). As originally defined, “Capital delta” is similar to “Big delta” because it represented the simple difference between the “δ-values” for the rarest isotopes in a given sample and the associated reference value on the identified massdependent fractionation (MDF) line. The MDF line for oxygen can be approximated by (e.g., Clayton et al. 1973): 18 δ17 MFL ≈ 0.52 δ MFL

(8a)

and one of the corresponding definitions for “Capital delta” is given as: 17 18 Δ17 A = δ A − 0.5 δ A

(8b)

Another definition for “Capital delta” is formulated using an alternative definition of the MDF line for oxygen: ⎧⎛ 1 + δ18 ⎞ 0.52 ⎫ ⎪ ⎪ MFL δ17 = 1000 × ⎟⎟ − 1⎬ ⎨⎜⎜ MFL 1000 ⎠ ⎭⎪ ⎩⎪⎝

(9a)

and the corresponding definition for “Capital delta” is given as: ⎧⎛ 1 + δ18 ⎞ ⎪ 17 A Δ17 A = δ A − 1000 × ⎨⎜ ⎜ 1000 ⎟⎟ ⎠ ⎩⎪⎝

0.52

⎫ ⎪ − 1⎬ ⎭⎪

(9b)

Capital delta prime and delta prime Several recent studies have reintroduced a different definition for the delta-value, originally proposed by Hulston and Thode (1965a), that we will refer to here as “delta prime” (see Hulston and Thode 1965a; Angert et al. 2003; Ono et al. 2003; Young and Galy 2004). The “delta prime” is defined as: ⎛ R ⎞ δ′x = 1000 ln ⎜ x ⎟ ⎝ R std ⎠

(10a )

With this definition, Equation (7b) is expressed exactly as: 1000 ln α A-B = δ′A − δ′B

(10 b)

A corresponding definition for “Capital delta prime” has recently been advocated by Young and Galy (2004) as: Δ′A17= δ′ 17 O − 0.5247 δ′ 18 O (10c) For a discussion of the relative merits of “Capital Delta” and “Capital Delta prime” see discussions in Young and Galy (2004), Miller (2002), Angert et al. (2003), and Kaiser et al.

Abundance, Notation, Fractionation of Light Stable Isotopes

21

(2004). In practice, the values reported using each definition differ by only a few hundredths of a per mil.

Material balance Mass conservation relationships are simply and compactly expressed for the minor HCNOS isotopes. For a closed system constituted of many parts, the relationships may be expressed either in terms of the isotope ratios R or δ values: R system ≅ ∑ X i R i

(11a )

δ system ≅ ∑ X i δi

(11b)

i

and

i

where the Xi are mole fractions, defined in terms of the relative proportions of the element of interest in each phase i. For isotopes of the HCNOS elements, these simple approximations have great accuracy.

COMMONLY-USED DIAGRAMS δ−δ plot The δ−δ plot, a graph that directly represents the δ values of two coexisting phases A and B as the x and y axes, is the natural coordinate system for discussing measured variations in stable isotope systems (Fig. 2). As shown by Gregory and Criss (1986), both fractionation relationships and material balance effects may be directly understood on such graphs. For any given system state, equilibrium or otherwise, a condition of fixed isotopic fractionation between the phases is accurately represented by a straight line with a unit slope and a yintercept of ΔA-B, according to Equation (6). Different fractionation conditions are represented by parallel lines; of course, the condition of zero fractionation is the line δΑ = δΒ. The bulk δ-value of a multiphase system is represented by a single point along this zero-fractionation line (Fig. 2). A great strength of the δ−δ plot is the ease with which open-system and closed-system processes may be distinguished. Suppose that a system, originally sited on a specific fractionation line, is subjected to different conditions that induce a change to a different fractionation condition, represented by a point on a parallel line. Such a change requires the redistribution of isotopes among the various phases. In a closed, two-phase system, one phase must necessarily increase its heavy isotope content at the expense of the other, so either δΑ increases while δB decreases, or vice versa. Clearly, the dynamic trajectory followed during this change must move along a straight line with a negative slope, −XA/XB, as required by Equation (11b). In contrast, under open-system conditions, the isotopic composition of the entire system can change; under most geologically relevant conditions, rock systems subject to such processes react along trends having positive slopes (Fig. 2). Gregory et al. (1989) provide many examples of the latter type where the isotopic systematics clearly prove that the rocks have been pervasively infiltrated by fugitive, oxygen-bearing fluids. Extension of these relationships to systems with three or more phases has been discussed elsewhere (e.g., Criss 1999), but in most cases the qualitative consequences on δ−δ plots are similar.

Big Δ and Cap Δ plots Gregory and Criss (1986) discuss the characteristics of a number of other plots, for example the Δ−δ plot or the Δ−Δ plot, that are sometimes used to represent isotopic data. The advantages of these plots are offset by several factors, most importantly that the measured

22

Criss & Farquhar

Figure 2. Characteristics of a δ−δ plot, along with an illustrative example given by a graph of the δ18O values of coexisting plagioclase and pyroxene in some terrestrial and extraterrestrial rock suites (after Gregory and Criss 1986). Isotopic equilibrium of a rock suite at high temperatures is indicated by trend lines having unit slopes and small fractionation factors, as exemplified by the achondrite and lunar rock suites. Closed system re-equilibration processes would generally produce negative slopes on this diagram, but large effects are rare. Open-system processes such as fluid infiltration are common and generally produce elongated trends with positive slopes on this diagram. Examples of the latter are the Skaergaard and Skye suites that underwent alteration and exchange with heated, low-18O meteoric waters, which consequently lowered the δ18O values of all phases. Similarly, the Oman gabbros were altered by oceanic hydrothermal systems, but the δ18O values of all phases increased because ocean water is much higher in 18O than most meteoric waters. Unreasonable temperatures would be calculated for most of the plagioclase-pyroxene fractionations exhibited by these altered suites.

data are not directly represented, and such diagrams are subject to induced correlations that make data trends look much stronger than they actually are. Plots of “Capital delta” versus δ and Δ′ vs. δ′ are used to identify mass-dependent fractionations for similar reasons, and are discussed in Farquhar and Theimens (2000), McKeegan and Leshin (2001), and Young and Galy (2004).

Three-isotope plot For elements having two or more stable isotopes in addition to the reference nuclide, it is possible to generate a graph having axes that depict two different types of δ-values. Hulston and Thode (1965a) invented such a diagram to compare δ33S-values, based on the 33S/32S ratio, to the usual δ34S-values, based on the 34S/32S ratio (Fig. 3). Moreover, they exploited the fact that the sulfur system has yet another stable isotope, 36S, that presents an analogous opportunity to directly compare δ36S and δ34S values. Hulston and Thode (1965b) found excellent linear correlations on these diagrams, whose slopes of ~0.51 and ~1.9 conform to those predicted for sulfur isotope equilibrium fractionations.

Abundance, Notation, Fractionation of Light Stable Isotopes

23

40 30

Seawater Sulfate

Hulston and Thode 1965 a, b

34

20

NMF

36

D S

36

:D DF M

10

S

=

1.

87

D

S

+

0.

36

0 -10

10 8

33

D S

CHONDRITES CARBONACEOUS CHONDRITES IRONS ACHONDRITES EARTH Iron Phase

Seawater Sulfate

34

6 4

: DF M

NMF

2

33

D

S

=

0.

50

9

D

S

+

0.

03

0 -2 -4 -10

-5

0

5

10

15

20

25

34

D S Figure 3. Two versions of a “three isotope” plot, which together illustrate the relationships among the four stable isotopes in the sulfur system (32S, 33S, 34S, 36S), are exemplified by the data of Hulston and Thode (1965a,b). An analogous diagram, a plot of δ17O vs. δ18O, can be made for the oxygen isotope system, and it has provided valuable insights into the origin of meteorites and planetary materials. Many examples of the latter are given elsewhere in this volume. The regressions are appropriate for normal mass-dependent fractionation (MDF) processes and represent fits to all data, except for the iron meteorite “iron phase” samples (open crosses). Several of the latter samples exhibit non-mass dependent fractionations (NMF) produced by cosmic ray spallation.

In the oxygen system, we commonly directly compare δ17O values, based on the 17O/16O ratio, to δ18O values, that are based on the 18O/16O ratio. Several important discoveries were made using plots of δ17O vs. δ18O. Practically all samples from Earth lie along the “terrestrial fractionation line” that has a slope of ~0.52, closely conforming to the predicted MDF slopes for oxygen. Remarkably, compositions of lunar rocks also plot on this line, now called the “Earth-Moon line,” yet most meteorites lie along parallel but offset trends (Clayton et al. 1973;

24

Criss & Farquhar

Clayton and Mayeda 1975, 1983; Clayton et al. 1976). This result is in marked contrast to the observations of Hulston and Thode (1965b), who found that terrestrial rocks and numerous meteorite types all lie along a single trend line on each of the three-isotope plots in the sulfur system. Even more remarkable was Clayton et al.’s (1973) discovery that certain meteorite components lie along slope ~1 trends instead of slope ~0.52 trends on the δ17O vs. δ18O plot, demonstrating that some nuclear process or, as is believed today, some type of NMF process profoundly affected these samples.

ISOTOPIC FRACTIONATION PROCESSES Mass-dependent fractionation Qualitative explanation. Chemical and physical processes can fractionate isotopes because energy differences accompany isotopic substitution. Various contributions to total energy such as kinetic energy, gravitational potential energy, vibrational energy and rotational energy all include mass terms, and consequently, relative differences in energy arise among substances and their isotopologues. As a rule of thumb, the relative energy difference between a substance and its isotopologue tends to increase along with the magnitude of the dimensionless quantity Δm/m, where m is the atomic mass of the reference nuclide and Δm is the difference between this mass and that of the isotope of interest. Other factors are important, but an expected consequence of this “rule” is that MDF effects are largest for H and tend to decrease with increasing atomic number, as pointed out by many workers (e.g., Bigeleisen 1965; Johnson et al. 2004). Energy differences among substances and their isotopologues give rise to several different types of isotopic fractionation. Equilibrium isotope effects relate to the thermodynamic tendency for given compounds to concentrate the heavy isotope relative to the lighter one(s). Nonequilibrium isotope effects can involve such effects along with additional isotope effects that relate to translational kinetic energy, which affects diffusion or atmospheric escape, or to the breakage or formation of bonds, etc. A particular result is that MDF processes will produce δ18O differences that are approximately twice as large as the corresponding δ17O differences, because the mass difference attendant on substituting 18O for 16O is almost exactly twice as large as the difference for substituting 17O for 16O. This yields Equations (8a) and (9a), that describe mass fractionation arrays for oxygen. Equilibrium isotope fractionations. Under normal conditions, multiphase systems will tend to achieve a condition of isotopic equilibrium in which the various phases have differing thermodynamic tendencies to incorporate the heavy isotope. For equilibrium among sufficiently simple species like diatomic gases, the equilibrium fractionation factors can be calculated using the methods of statistical thermodynamics (Urey 1947). Elementary descriptions of Urey’s theory are given elsewhere (e.g., O’Neil 1986b; Criss 1999; Chacko et al. 2001; Schauble 2004); and detailed calculations for many gas species are tabulated by Richet et al. (1977). A qualitative explanation for equilibrium MDF effects is that, when a heavy isotope is substituted for the lighter isotope of the same element in a given molecular or lattice position, the vibrational and any rotational frequencies associated with that position are lowered. It follows that the quantized energy levels available to that subsystem are lowered. Because the thermodynamic character of a substance depends upon the average energy of its constituent molecules or unit cells, all phases favor incorporation of heavy isotopes to lower their total energy. The question is, among a collection of different substances, which will preferentially “win out” in the universal competition for the heavy isotope? As a rule of thumb, the heavy isotope is most concentrated in the substance having the strongest bonds to the element in question, because in such substances the greatest frequency reduction and energy lowering is realized upon the isotopic substitution. Entropy, of course, requires that all of the substances

Abundance, Notation, Fractionation of Light Stable Isotopes

25

contain at least some of the heavy isotope, so this is not an “all or nothing” process- in fact, the relative differences in isotopic concentrations are small. Temperature dependence. Isotopic fractionation factors approximate perfect geothermometers because the volume change upon isotopic substitution is small or insignificant, making pressure effects small under most normal conditions. However, the fractionation behavior of real systems can be complex and commonly cannot be accurately predicted. Nevertheless, the behaviors predicted for simple gases by Urey’s theory under “high temperature” (or low frequencies ) or “low temperature” (or high frequencies) conditions commonly approximate those exhibited by complex phases under these conditions (e.g., Fig. 4). The basic proportionalities predicted by statistical theory for isotopic equilibria under various temperature conditions are given below; derivations and discussion are provided elsewhere (Stern et al. 1968; Criss 1991). Low T

ln α = C1/T + C2

(12a)

High T

ln α = C3/T

(12b)

2

T→∞ ln α → 0 (12c) The numerical constants (C’s) in Equations (12a) and (12b) generally depend on vibrational frequencies. Laboratory determinations of fractionation factors are commonly reported in the forms of Equations (12a) or (12b), or as a combination of such terms, with empiricallydefined coefficients (Fig. 4). At extremely high temperatures, there is no preference for the heavy isotope to concentrate in any particular phase, and α approaches unity, as predicted by Equations (12b) and (12c). 1200°

1000°

800°

600°C

2

1000 ln D mineral-An50

Quartz

0 Diopside

-2 Forsterite Magnetite

-4

Perovskite CaTiO

-6

-8 0.4

3

0.6

0.8

1 6

10 /T

1.2

1.4

2

Figure 4. Plot of 1000 lnα, here representing the “isotopic fractionation” between intermediate plagioclase (Ab50-An50) and the indicated mineral, versus 1/T 2 for various minerals. At high temperatures, the isotopic fractionations among minerals vary as a linear function of 1/T 2. Fractionation data used in this diagram are derived from the tabulation of Chacko et al. (2001).

26

Criss & Farquhar

Multiple stable isotopes. A special fractionation relationship can be developed for the case where any of three or more isotopes having masses of m, m′ and m′′ can be substituted in a particular molecular or lattice site. Under the “high T” conditions of Equation (12b) the various isotopic fractionation factors for ordinary, MDF processes are interrelated by (Criss 1999): αm′/m = αθm′′/m

(13)

where the exponent θ depends only on the isotope masses and not on any characteristics of the site or bond, and its high temperature limit is given by: θ=

m′′ ⎛ m′ − m ⎞ m′ ⎜⎝ m′′ − m ⎟⎠

(14)

At lower temperatures, θ will vary from its high-temperature value, but usually only by a few percent. In such cases, θ must be determined from its fundamental definition. For example, for oxygen: θ AB =

( ln ( ln

17 18

fA

17

fB

fA

18

fB

) )

(15)

where 17fA refers to the reduced partition function for 17O-16O of phase A, and analogously for the other terms. Another term, “lambda” (λ), is used to describe observed mass-dependent isotopic fractionations: λ AB =

δ′17O A − δ′17O B δ′18O A − δ′18O B

(16)

Although it may seem that λAB and θAB should be equivalent, there are cases where they are not, such as during Rayleigh distillation and in reaction networks (c.f., Blunier et al. 2002; Angert et al. 2003; Farquhar et al. 2003).

Kinetic processes Stable isotope kinetics involve dynamic, unidirectional processes that redistribute isotopes among different phases or species. Such processes may or may not involve kinetic fractionation factors. Progress of exchange reactions. The most common and best understood kinetic processes involve temporal changes of the isotope ratios of unequilibrated phases that are undergoing mutual isotopic exchange as they strive to approach the equilibrium state. Most such phases participate in isotopic exchange reactions that are pseudo-first order in character (see Cole and Ohmoto 1986; Criss et al. 1987), although Johnson et al. (2004) discuss other possibilities. In numerous systems, the phenomenological rate of change of the isotope ratio RA of phase A is directly proportional to the deviation of that ratio from its equilibrium value with phase B at that instant (Criss et al. 1987): dR A = k ( R A − αR B ) dt

(17)

where k is a rate constant and t is time. Numerous solutions to this pseudo-first order equation are provided by Criss et al. (1987) and Criss (1999). Typical solutions exhibit an approach to equilibrium whose rate decreases exponentially over time, or decreases according to a simple sum of different exponential terms. Criss (1999) and Cole and Chakraborty (2001) provide experimental examples. Kinetic fractionation factors. For many processes, such as evaporation, isotope ratios progressively change yet equilibrium is not approached. Such processes involve kinetic

Abundance, Notation, Fractionation of Light Stable Isotopes

27

fractionation factors in addition to, or exclusive of, the normal equilibrium factors that govern exchange reactions. As an example, kinetic fractionation factors result when reaction rates, or rates of bond breaking, differ for chemical substances having different isotopic substitutions; see the review by Bigeleisen and Wolfsberg (1958). Chemical kinetic fractionation factors are described by partition function ratios involving partition functions (f) for both the reactants (A) and the transition state (‡): α A‡ =

k v ‡ ∏ fA = k * v*‡ f ‡

(18)

The vibrational frequencies (v‡) are imaginary frequencies that are related to the decomposition of the transition state. The mass-dependence of this fractionation factor has been described in limiting cases by Bigeleisen and Wolfsberg (1958) and the mass-dependent relationships of these cases have been described by Young et al. (2002) and Schauble (2004). The massdependent relationships at temperatures other than the limiting cases will follow: θ A‡ =

‡ ) θ A ln(∑ 17 f A ) − θ‡ ln( 17 f‡ ) + ln( v ‡ v17 ‡ ) ln(∑ 18 f A ) − ln( 18 f‡ ) + ln( v ‡ v18

(19)

where θA and θ‡ refer to the mass-dependent relationships between the reduced partition functions for the 17O- and 18O-substituted species of the reactants (A) and the transition state (‡). As a cautionary note, we point out that values for θA‡ calculated on the basis of statistical thermodynamics for specific reactions and for conditions other than the limiting case may be different from those for the limiting cases presented in the literature.

Non-mass-dependent fractionations Qualitative explanation. A class of isotopic fractionation processes distinguished as “non-mass-dependent” deviate significantly from the mass-dependent relationships. We use the descriptive term “non-mass-dependent” rather than “mass-independent” because some of the theoretical treatments proposed to explain these effects include mass terms and indicate that the processes need not be fully independent of mass, but that other factors contribute a nonmass-dependent component to the isotopic fractionations. Non-mass-dependent fractionations have been attributed to a variety of chemical effects, including electronic-nuclear spin coupling involving odd-mass isotopes, differences in the densities of bound states during transition state chemistry that result from isotopically-induced changes in molecular symmetry, differences in the probability of intersystem crossings or internal conversions between bound electronic states induced by isotopic substitutions, and differences in photochemical reaction rates caused by self and mutual shielding. Spin coupling effects. Clear descriptions of magnetic isotope fractionation effects can be found in Turro (1983). These effects occur as a result of a spin-dependent reaction rate enhancement for odd-mass isotopes because they have an unpaired nuclear spin that can be coupled to the electron spin, allowing for the electronic spin to change sign while conserving total spin. Spin-selective reactions have been studied for radical pairs that are prevented from reacting because of incompatible electronic spin. When the radical pair is held in proximity by a “cage” for timescales comparable to those for spin reorganization, preferential recombination reactions can occur for reactions involving odd mass isotopes. If the timescales are too short for the spin reorganization to occur, or too long to allow the escape of the non-magnetic radical pairs, the rate advantage is diminished or lost. Effects associated with metastable and bound states. Kinetic fractionation factors will be mass-dependent when the reaction rates depend on the relative mass difference between the isotopes, but when additional considerations such as symmetry come into play, the relative reaction rates of some isotopically substituted species can be significantly enhanced,

28

Criss & Farquhar

producing anomalous NMF effects. R. Marcus and coworkers (Hathorn and Marcus 1999, 2000; Gao et al. 2001) have developed a modified Rice-Rammelsberger-Kassel-Marcus transition state explanation that uses an empirical factor (η) to describe the unusual and anomalous rate enhancements associated with ozone chemistry. In its most basic form η reflects the effects of a higher density of states for the potential energy surface of ozone formation for asymmetric transition state species. More recently, Babikov et al. (2003) have examined ozone recombination reactions using ab initio quantum dynamical treatments of energy transfer mechanisms and metastable states. They have been able to attribute the effect to specific metastable states and also to provide a more detailed assessment of the reasons why the isotope partitioning reactions occur. Other NMF effects that have been investigated experimentally have been attributed to enhancement of reaction rates associated with rovibronic overlap involving intersections of potential energy surfaces for bound states and between predissociative states and dissociative states (Coleman et al. 1997; Bhattacharya et al. 1999; Zmolek et al. 1999). Bhattacharya et al. (1999) document changes in the non-mass-dependent character for oxygen produced during photodissociation of CO2 by 185.4 nm radiation. They observed significant changes in the δ17O and δ18O fractionations when 13CO2 was photodissociated compared to experiments in which 12 CO2 was photodissociated. Bhatacharya et al (1999) argue that the carbon isotope substitution shifts the populations of oxygen-substituted species among the rotational-vibrational energy levels within a predissociative bound state. This changes the distribution of oxygen-substituted species relative to the intersections between the predissociative potential energy surface and the dissociative PES, and shifts the relative rate advantages for 16O-, 17O-, and 18O-substituted species. Related experimental observations and hypotheses were presented by Zmolek et al. (1999) for CS2. Shielding effects. The availability of electromagnetic radiation for chemistry depends both on the nature of the source of the radiation and the absorption characteristics of material through which the radiation passes. Absorption of electromagnetic radiation occurs when a photon couples in some way with the molecular vibrations, electronic transitions, or other properties of the material through which it passes. The absorption spectra that result from coupling of electromagnetic radiation with molecular vibrations and electronic transitions have a dependence on isotopic substitutions, and small shifts in absorption spectra are observed for molecules that differ only in their isotopic compositions. The absorption depth is a property that describes the distance that electromagnetic radiation will be transmitted through a material, and the absorption depths of gases with different isotopic substitutions vary depending on both the relative abundances and the absorption cross sections of the isotopically-substituted species. Shielding refers to situations where radiation that is required to drive photochemical reactions has been fully absorbed by matter in the path of the radiation. Self-shielding refers to the case where the absorbing material is the same material as the material that is shielded (e.g., 12C16O shielding 12C16O further along the radiation path). Mutual shielding occurs when the absorbing material is distinct from the material that is shielded (e.g., 32S16O2 shielding an absorption band for another SO2 isotopologue further along the radiation path). Shielding becomes important when the radiation that is attenuated is necessary for photochemistry, and non-mass-dependent isotopic fractionations can arise if differences in the absorption depths for molecules with different isotopic abundances provide a rate advantage for reactions involving some isotopically-substituted species over others.

CONCLUSIONS Stable isotopes have become an essential tool for characterizing and understanding terrestrial and extraterrestrial matter. The special techniques needed to make accurate isotopic measurements gave rise to special notations for reporting stable isotope data, and these notations

Abundance, Notation, Fractionation of Light Stable Isotopes

29

in turn gave rise to special diagrams that emphasize compositional differences and facilitate interpretation. The processes that govern isotope distribution have become progressively better understood, yet recent studies show that these processes are more diverse than anticipated only ten years ago.

REFERENCES Angert A, Rachmilevitch S, Barkan E, Luz B (2003) Effects of photorespiration, the cytochrome pathway, and the alternative pathway on the triple isotopic composition of atmospheric O2. Global Biogeochem Cycles 17:doi 10.1029/2002GB001933 Babikov D, Kendrick BK, Walker RB, Pack RT, Fleurat-Lesard P, Schinke R (2003) Formation of ozone: metastable states and anomalous isotope effect. J Chem Phys 119:2577-2589 Bhattacharya SK, Savarino J, Thiemens MH (1999) A new class of oxygen isotopic fractionation in photodissociation of carbon dioxide: Potential implications for atmospheres of Mars and Earth. Geophys Res Lett 27:1459-1462 Bigeleisen J (1965) Chemistry of isotopes. Science 147:463-471 Bigeleisen J, Wolfsberg M (1958) Theoretical and experimental aspects of isotope effects in chemical kinetics. Adv Chem Phys 1:15-76 Blunier T, Barnett B, Bender ML, Hendricks MB (2002) Biological oxygen productivity during the last 60,000 years from triple oxygen isotope measurements. Global Biogeochem Cycles 16:doi 10.1029/2001GB001460 Chacko T, Cole DR, Horita J (2001) Equilibrium oxygen, hydrogen and carbon isotope fractionation factors applicable to geologic systems. Rev Mineral Geochem 43:1-81 Clayton RN, Grossman L, Mayeda TK (1973) A component of primitive nuclear composition in carbonaceous chondrites. Science 182:485-488 Clayton RN, Mayeda TK (1983) Oxygen isotopes in eucrites, shergottites, nakhlites, and chassignites. Earth Planet Sci Lett 62:1-6 Clayton RN, Mayeda TK (1975) Genetic relations between the moon and meteorites. Proc Lunar Sci Conf 6:1761-1769 Clayton RN, Onuma N, Mayeda TK (1976) A classification of meteorites based on oxygen isotopes. Earth Planet Sci Lett 30:10-18 Cole DR, Chakraborty S (2001) Rates and mechanisms of isotopic exchange. Rev Mineral Geochem 43:83223 Cole DR, Ohmoto H (1986) Kinetics of isotopic exchange at elevated temperatures and pressures. Rev Mineral 16:41-90 Coleman JJ, Xu XP, Thiemens MH, Trogler WC (1997) Photopolymerization and mass-independent sulfur isotope fractionations in carbon disulfide. Science 273:774-776 Criss RE (1991) Temperature dependence of isotopic fractionation factors. Geochem Soc Special Pub 3:1116 Criss RE (1999) Principles of Stable Isotope Distribution. Oxford University Press, New York Criss RE, Gregory RT, Taylor HP Jr. (1987) Kinetic theory of oxygen isotopic exchange between minerals and water. Geochim Cosmochim Acta 51:1099-1108 Farquhar J, Thiemens MH (2000) Oxygen cycle of the Martian atmosphere-regolith system: Δ17O of secondary phases in Nakhla and Lafayette. J Geophys Res-Planets 105(E5):11991-11997 Farquhar J, Johnston DT, Wing B.A, Habicht KS, Canfield DE, Airieau A, Thiemens MH (2003) Multiple sulphur isotopic interpretations of biosynthetic pathways: implications for biological signatures in the sulphur isotope record. Geobiology 1:27-36 Gao YQ, Marcus RA (2001) Strange and unconventional isotope effects in ozone formation. Science 293:259263 Gregory RT, Criss RE (1986) Isotopic exchange in open and closed systems. Rev Mineral 16:91-127 Gregory RT, Criss RE, Taylor HP Jr. (1989) Oxygen isotope exchange kinetics of mineral pairs in closed and open systems: Applications to problems of hydrothermal alteration of igneous rocks and Precambrian iron formations. Chem Geol 75:1-42 Hathorn BC, Marcus RA (1999) An intramolecular theory of the mass-independent isotope effect for ozone. I. J Cheml Phys 111:4087-4100 Hathorn BC, Marcus RA (2000) An intramolecular theory of the mass-independent isotope effect for ozone. II. Numerical implementation at low pressures using a loose transition state. J Chem Phys 113:9497-9509 Hulston JR, Thode HG (1965a) Cosmic-ray-produced in S36 and S33 in the metallic phase of iron meteorites. J Geophys Res 70:4435-4442

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Hulston JR, Thode HG (1965b) Variations in the S33, S34, and S36 contents of meteorites and their relation to chemical and nuclear effects. J Geophys Res 70:3475-3484 Johnson CM, Beard BL, Albarede F (2004) Overview and general concepts. Rev Mineral Geochem 55:1-24 Kaiser J, Rockmann T, Brenninkmeijer C AM (2004) Contribution of mass-dependent fractionation to the oxygen isotope anomaly of atmospheric nitrous oxide. J Geophys Res-Atmospheres 109(D3). Art no D03305 McKeegan KD, Leshin LA (2001) Stable isotope variations in extraterrestrial materials. Rev Mineral Geochem 43:279-318 McKinney CR, McCrea JM, Epstein S, Allen HA, Urey HC (1950) Improvements in mass spectrometers for the measurement of small differences in isotope abundance ratios. Rev Sci Instrum 21:724-730 Miller MF (2002) Isotopic fractionation and the quantification of 17O anomalies in the oxygen three-isotope system – an appraisal and geochemical significance. Geochim Cosmochim Acta 66:1881-1889 O’Neil JR (1986a) Terminology and standards. Rev Mineral 16:561-570 O’Neil JR (1986b) Theoretical and experimental aspects of isotopic fractionation. Rev Mineral 16:1-40 Ono S, Eigenbrode JL, Pavlov AA, Kharecha P, Rumble D, Kasting JF, Freeman KH (2003) New insights into Archean sulfur cycle from mass-independent sulfur isotope records from the Hammersley Basin, Australia. Earth Planet Sci Lett 213:15-30 Richet P, Bottinga Y, Javoy M (1977) A review of hydrogen, carbon, nitrogen, oxygen, sulphur, and chlorine stable isotope fractionation among gaseous molecules. Ann Rev Earth Planet Sci 5:65-110 Schauble EA (2004) Applying stable isotope fractionation theory to new systems. Rev Mineral Geochem 55:65-111 Stern MJ, Spindel W, Monse EU (1968) Temperature dependence of isotope effects. J Chem Phys 48:29082919 Turro NJ (1983) Influence of nuclear spin on chemical reactions: Magnetic isotope and magnetic field effects (A review). Proc Nat Acad Sci USA-Phys Sci 80:609-621 Urey HC (1947) The thermodynamic properties of isotopic substances. J Chem Soc (London) 562-581 Walker FW, Parrington JR, Feiner F (1989) Nuclides and Isotopes, 14th ed., General Electric Co., San Jose, California Young ED, Galy A (2004) The isotope geochemistry and cosmochemistry of magnesium. Rev Mineral Geochem 55:197-230 Young ED, Galy A, Nagahara H (2002) Kinetic and equilibrium mass-dependent isotope fractionation laws in nature and their geochemical and cosmochemical significance. Geochim Cosmochim Acta 66:10951104 Young ED, Kuramoto K, Marcus RA, Yurimoto H, Jacobsen SB (2008) Mass-independent oxygen isotope variation in the solar nebula. Rev Mineral Geochem 68:187-218 Zmolek P, Xu XP, Jackson T, Thiemens MH, Trogler WC (1999) Large mass independent sulfur isotope fractionations during the photopolymerization of (CS2)-12C and (CS2)-13C. J Phys Chem A 103(15):24772480

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 31-53, 2008 Copyright © Mineralogical Society of America

Nucleosynthesis and Chemical Evolution of Oxygen Bradley S. Meyer Department of Physics and Astronomy Clemson University Clemson, South Carolina 29634, U.S.A. [email protected]

Larry R. Nittler and Ann N. Nguyen Department of Terrestrial Magnetism Carnegie Institute of Washington Washington, D.C. 20015, U.S.A.

Scott Messenger Robert M. Walker Laboratory for Space Science NASA Johnson Space Center Houston, Texas 77058, U.S.A. ABSTRACT Of the elements strictly synthesized in stars, oxygen is by far the most abundantly produced. We review the nucleosynthesis and galactic chemical evolution of this important element. We then review its isotopic composition in presolar grains recovered from primitive meteorites and from interplanetary dust particles. As we describe, knowledge of these isotopic compositions provide important constraints on theories of nucleosynthesis, stellar evolution, and galactic chemical evolution.

INTRODUCTION The isotopes of oxygen are crucial diagnostics of nucleosynthesis and galactic chemical evolution. This is due in large part to oxygen’s high abundance. It is the third most abundant element in the Galaxy, though considerably less abundant than hydrogen and helium, which are principally produced in primordial nucleosynthesis. Of the elements strictly synthesized in stars, oxygen is by far the most abundantly produced. Oxygen’s abundance is nearly equal to that of all the other heavy elements (elements with atomic number greater than that of helium) combined (e.g., Lodders 2003). This high abundance means that it is readily observable in the atmospheres of stars. It also means that oxygen is a dominant player in chemical reactions taking place in stellar outflows, in the interstellar medium, and in the early Solar System. The abundance of oxygen is dominated by that of 16O, which is the third most abundant isotope in the Galaxy. The other stable isotopes of oxygen are considerably less abundant. 18O is the 25th most abundant isotope in the Solar System, with an abundance about 0.2% that of 16O. 17 O is the 34th most abundant isotope in the Solar System, and its abundance is 0.03% that of 16 1 O . Despite the low relative abundances of 17O and 18O, they are still sufficiently abundant that 1

The isotopic abundances and rankings presented here are those according to the compilation of Lodders (2003). For further details on oxygen abundances in the Solar System, please see Davis et al. (2008). The interested reader may also wish to explore the Solar System abundances and their rankings with the Webnucleo online solar abundances tool available at http://www.webnucleo.org/home/online_tools/solar_abundances/

1529-6466/08/0068-0004$05.00

DOI: 10.2138/rmg.2008.68.4

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they can be measured in stellar atmospheres. As we shall also see, oxygen isotopic abundances are also readily measured in presolar grains in meteorites and interplanetary dust particles. These measurements provide key constraints on nucleosynthesis and stellar evolution. Also contributing to oxygen’s importance as a diagnostic is the fact that its three stable isotopes are predominantly made in different burning epochs during the life of a star. As we shall see below, 17O is predominantly made in hydrogen burning, 18O is predominantly made in the early stages of helium burning, and 16O is made in the later part of helium burning. Thus, the isotopic composition of a stellar atmosphere or a presolar grain provides clues to the stellar region in which the major part of the material was synthesized. Finally, a third significant aspect of oxygen’s role as a diagnostic of nucleosynthesis and galactic chemical evolution is the primary versus secondary nature of the different isotopes. 16 O is a primary isotope, that is, an isotope that can be synthesized in a star initially composed only of hydrogen and helium, while 17O and 18O are secondary isotopes, which means that their formation requires pre-existing seed nuclei from previous stellar generations. As a consequence, the abundance of 16O relative to those of 17O and 18O changed with time in the Galaxy’s history. Measurements of oxygen abundances in stars or in presolar grains of different ages can thus provide important clues about the chemical evolution of the Galaxy (e.g., Clayton 1988). The goal of this paper is to review the nucleosynthesis and chemical evolution of the isotopes of oxygen and the manifestations of these processes in astronomical and cosmochemical samples. We begin in the section “Nucleosynthesis of the Isotopes of Oxygen” with a review of the nucleosynthesis of oxygen isotopes in the different burning stages in stars and follow up with an analysis of the oxygen isotopic abundances from high-mass and low-mass stars. We then turn in the section “Chemical Evolution of the Isotopes of Oxygen” to a review of the key issues of the galactic chemical evolution of the oxygen isotopes. In the section “Oxygen in Presolar Grains,” we review some of the implications of oxygen data on presolar grains for nucleosynthesis and chemical evolution. In the last section, we present some concluding remarks.

NUCLEOSYNTHESIS OF THE ISOTOPES OF OXYGEN As a whole, our Universe has only about 10−10 nucleons for every photon. By contrast, the corresponding number in stars is close to unity. The low nucleon-to-photon ratio in the early Universe strongly inhibited build-up of heavy nuclei, and, as a consequence, primordial nucleosynthesis did not produce any oxygen. The oxygen present in our Solar System was thus made in stars. In this section, we first review oxygen nucleosynthesis in the main stellar burning stages. We then analyze the oxygen yields predicted for high-mass and low-mass stars, as well as for novae and Type Ia supernovae.

Production of oxygen in mainline stellar burning stages Hydrogen burning. Stars shine. As a consequence, they radiate away energy. To compensate this energy loss and maintain their pressure support against gravitational contraction, stars convert nuclear energy into thermal energy by fusing lighter nuclei into heavier ones. This nuclear burning proceeds through a series of stages in which the ashes of one stage serve as the fuel for the next. This sequence of burning stages progresses until either electron degeneracy replaces thermal pressure to support the star or the star burns the nuclei all the way to iron, from which no more nuclear energy can be extracted by fusion. The first nuclear fuel available to stars is hydrogen, which stars burn into 4He. Stars with the mass of the Sun or less burn hydrogen at relatively low temperatures (10-20 million K) via the so-called PP chains. Because the PP chains only involve isotopes of hydrogen, helium, lithium, and beryllium, these reaction chains do not affect the stellar abundance of oxygen. In contrast,

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stars more than about 1.2 times the mass of the Sun attain somewhat higher temperatures and can thus burn hydrogen by the more efficient process known as CNO burning. In this process, the production of 4He from 1H is catalyzed by two coupled cycles involving isotopes of carbon, nitrogen, and oxygen. In the CN cycle, the reaction sequence is: 12C + 1H → 13N + γ (here γ represents energy released during the reaction as a gamma ray which subsequently deposits its energy locally); 13N → 13C + e+ + νe (here e+ represents a positron and νe represents an electrontype neutrino); 13C + 1H → 14N + γ; 14N + 1H → 15O + γ; 15O → 15N + e+ + νe; and 15N + 1H → 12 C + 4He. This last reaction closes the cycle. Roughly 0.1% of the time, however, 15N does not return to 12C, but rather captures a proton to become 16O: 15N + 1H → 16O + γ. This initiates the NO cycle: 15N + 1H → 16O + γ; 16O + 1H → 17F + γ; 17F → 17O + e+ + νe; and 17O + 1H → 14 N + 4He. It is the NO cycle that predominantly involves oxygen isotopes in hydrogen burning, although the minor pathway 17O + 1H → 18F + γ; 18F → 18O + e+ + νe; and 18O + 1H → 15N + 4 He brings 18O into play as well. During CNO burning, the slow reaction is 14N + p → 15O + γ (here p represents a proton; this reaction can also be written as 14N(p,γ)15O). This means that CNO cycling tends to convert the initial abundances of carbon and oxygen into 14N. This is evident in Figure 1, which shows the mass fractions of various species yielded by a hydrogen burning calculation with the Clemson

Figure 1. Evolution of the mass fractions of the indicated species during hydrogen burning at a temperature of 3×107 K and 10 g/cm3.

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nucleosynthesis network code (Jordan and Meyer 2004). The calculation began with a solar abundance distribution (Anders and Grevesse 1989) and was run at a constant temperature of 3×107 K and a density of 10 g/cm3, conditions fairly typical of hydrogen burning in the cores of massive stars. As is evident, CN burning converts 12C into 14N early in the burning while the reaction 18O + p → 15N + 4He quickly depletes the initial 18O and involves O in the main CNO cycling. As time progresses in the calculation, the 17O abundance rises due to proton capture onto 16O and achieves a maximum enrichment of about a factor of ten relative to its initial abundance. After about 104 years, however, the 17O abundance attains a steady state as destruction via 17O(p,4He)14N balances the production from 16O. Finally, after 106 years, the full CNO bi-cycle achieves a steady state. The 1H converts into 4He and the oxygen isotopes convert into 14N. After complete CNO burning, 4He and 14N are enriched while 12C and 16,17,18O are all depleted (Fig. 1). Nevertheless, if the CNO cycling is not complete or the burning happens at a lower temperature, the matter may be enriched in 17O and depleted in 16,18O, as is the case in Figure 1B at a time of 106 yr. This is typically the case in the envelopes of stars that have experienced dredge up (mixing) of matter from a hydrogen burning shell. Helium burning. After a star has burned its hydrogen into helium, the next available fuel is the 4He. Due to the lack of stable isotopes with mass number five and eight, the helium burning proceeds via the triple-alpha process, which may be viewed as the reaction 4He + 4He + 4He → 12 C + γ. Helium burning typically occurs at temperatures of ~1-3×108 K and densities near 1000 g/cm3. Figure 2 shows the evolution of mass fractions relative to solar values in a helium burning calculation at a temperature of 2.5×108 K and a density of 1000 g/cm3 and using as the initial abundances the final yields from the previous hydrogen burning calculation. In the initial stages of helium burning, the abundant 14N captures 4He to produce 18F, which decays to 18O, thereby strongly enriching 18O relative to 16O and 17O. The 18O itself then captures another 4He, which depletes the 18O and creates 22Ne. At the high temperature of the calculation in Figure 2, this conversion of 14N into 22Ne occurs within the first 0.1 yr. As this is occurring, the triple-alpha reaction is also converting 4He into 12C. Beginning at about one year, the 12C becomes sufficiently abundant that it can capture 4He to become 16O. This strongly enriches 16O relative to the other oxygen isotopes. The 12C/16O ratio resulting from helium burning determines the nature of the subsequent carbon burning in the star, which, in turn, determines the whole subsequent shell structure of the star (e.g., El Eid et al. 2004). For this reason, the reaction 12C(4He,γ)16O, which is not yet fully characterized in the laboratory, is the subject of intense experimental study. As a final point, it is worth noting that the reaction 16O + 4He → 20Ne + γ does not occur at helium-burning temperatures because of the lack of an appropriate resonance in 20Ne. A dominant mechanism for two nuclei to react is to proceed through formation of a compound nucleus, which is, in effect, a resonance in the scattering state of the two interacting nuclei. The compound nucleus subsequently breaks apart into different nuclei or de-excites to the product nucleus ground state. For example, when 12C and 4He combine, they can form an excited state (that is, a resonance) of 16O if the energy and the spin and parity changes in the interaction are appropriate. Most of the time, the excited 16O nucleus will break apart into 4He and 12C again. In some cases, however, the excited 16O decays by photon emission to the ground state. In the case of 20Ne, the typical interaction energies of 16O and 4He in helium burning are appropriate to form the 4.969 MeV state in 20Ne, which would make an ideal compound nucleus, but the spin and parity changes are not right, so the reaction does not occur (e.g., Clayton 1968). Alpha capture therefore ceases at 16O, and, as a consequence 16O is, in fact, the dominant product of helium burning. From Figure 2, it is clear that there is a phase in helium burning in which 18O is enhanced while 16O and 17O are depleted. As the burning progresses, however, the 18O converts to 22Ne

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Figure 2. Evolution of the mass fractions of the indicated species during helium burning at a temperature of 2.5×108 K and a density of 1.0×103 g/cm3.

and 16O becomes the dominant oxygen isotope. Because the 16O so greatly dominates the other oxygen isotopes after helium burning, we will henceforth neglect these minor isotopes in the subsequent discussion. Carbon burning. The next stellar burning stage is carbon burning. Carbon burning typically occurs at temperatures near 9×108 K and densities near 105 g/cm3. It predominantly converts 12 C into 20Ne and 24Mg via the reactions 12C + 12C → 20Ne + 4He and 12C + 12C → 24Mg + γ. At carbon-burning temperatures, there are resonances in the 16O(4He,γ)20Ne reaction available that allow the reaction to proceed efficiently. This means that the 16O left after helium burning is depleted somewhat during carbon burning due to capture of 4He produced by the main carbonburning reactions. Neon burning. After carbon burning, a star burns neon. This effectively occurs in a twostep process. The first reaction is 20Ne + γ → 16O + 4He. This reaction is endothermic. The 4 He produced, however, then captures on another 20Ne to produce 24Mg. This latter reaction is sufficiently exothermic that the net reaction, 20Ne + 20Ne → 16O + 24Mg, is exothermic. Clearly, neon burning replenishes some of the 16O depleted during carbon burning.

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Oxygen and silicon burning. Once neon burning is complete, the next burning stage is that of oxygen. The dominant reaction is 16O + 16O → 28Si + 4He. This phase clearly depletes the oxygen. After oxygen burning, the star will burn 28Si into 56Fe. The effective reaction is 28 Si + 28Si → 56Ni + γ followed by decay of 56Ni to 56Fe; however, the burning actually occurs by burning through quasi-equilibrium clusters (e.g., Woosley et al. 1973). Little oxygen is produced in this 28Si burning (e.g., Bodansky et al. 1968). Stellar explosion. Stars that are able to evolve all the way to silicon burning develop iron cores. At that point, the nuclei are in the state with the highest nuclear binding; hence, any rearrangement of their abundances cannot release energy. This means that the star can no longer burn fuel to maintain its pressure against gravity and the stellar core collapses homologously. The inner core collapses subsonically. Once it attains nuclear matter densities, the collapse halts and the core bounces. The outer core, however, collapses supersonically and therefore does not receive the signal to stop collapsing. It crashes onto the already collapsed inner core. This highly non-adiabatic effect generates a shock wave that propagates out through the outer layers of the star. As the shock passes through the stellar layers, it heats and compresses them. This causes further burning (known as explosive burning). The energy imparted to the outer layers by the shock eventually expels them and injects them into interstellar medium. This dramatic stellar death is known as a supernova. In particular, it is a “core-collapse” supernova since the energy powering the explosion arose from the release of gravitational binding energy in the collapse of the stellar core. For a review, see, for example, Woosley et al. (2002). The modification of the oxygen isotopes due to explosive burning is largely confined to the inner layers of the star. In particular, explosive oxygen burning will deplete the 16O in the 16 O-rich layers and explosive neon burning will enhance the 16O in the 20Ne-rich layers. By the time the shock reaches the 12C-rich and 4He-rich regions, however, the post-shock temperatures and densities do not lead to significant changes in the pre-supernova oxygen abundances.

Analysis of the oxygen yields from massive stars Armed with an understanding of the nucleosynthesis in the mainline burning stages of stellar evolution, we may now analyze the yields from massive stars—those with masses greater than about ten solar masses. Such stars are able to burn their nuclei all the way to iron. Stars with mass less than about ten solar masses develop cores that are supported by electron degeneracy, which halts their evolution before the advanced nuclear burning stages can occur. For a discussion of degeneracy, see the section on low-mass stars below. To analyze the yields from massive stars, we consider a single model. While this is only one particular model, its yields are fairly representative of the ejecta from any star more than ten times the mass of the Sun. Chemical evolution models using detailed stellar model yields have long shown that the Galaxy’s supply of 16O is dominantly from massive stars (e.g., Tinsley 1980). Because 16O so dominates the oxygen abundance, massive stars are thus nearly the sole contributors to the Galaxy’s inventory of this important element even though they compose only a few percent of the stars born in any stellar generation. Massive stars also dominate the galactic production of 18O (e.g., Timmes et al. 1995). Low-mass stars only dominate the galactic synthesis of the low-abundance isotope 17O. The model presented here is that of Meyer (2005), using the stellar evolution code of The et al. (2000) and El Eid et al. (2004) but with coupled nuclear burning and convection using routines written by Jordan et al. (2005), for a star that began with a mass 25 times greater than that of the Sun. By the end of its life, it had lost 3.4 solar masses of material due to stellar winds. The model ran through silicon burning. The collapse of the core and the explosion of the star were then simulated by numerically increasing the energy in the core of the stellar model. This led to the generation of an outwardly propagating shock wave in the model. Shock propagation and the concomitant explosive burning were tracked with the appropriate computer codes.

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The final yields (after stellar explosion) of the isotopes of oxygen relative to their initial abundances are shown in Figure 3. The abscissa is the so-called Lagrangian mass coordinate, Mr. The value of Mr indicates the mass contained within a given spherical shell in the star. For example, Mr = 0 is the center of the star. Mr = M, where M is the total mass of the star, is the star’s surface. Mr = 10 solar masses locates a spherical shell that contains 10 solar masses inside. The Lagrangian mass coordinate is convenient because, although the radius of a shell expands or contracts as the star evolves, the mass it contains does not. As Figure 3 shows, the abundances of the oxygen isotopes vary dramatically inside the stellar ejecta. Particularly noteworthy is the fact that abundances are uniform over certain mass ranges. This indicates mixing within the stellar regions. For example, the abundances of 16O, 17 O, and 18O are uniform from Mr ~7.8 to 21.6 solar masses. This represents the outer envelope of the star, which stretches from what was the hydrogen-burning shell to the stellar surface. The strong convection in the envelope prior to the explosion homogenizes the abundances. Similarly, the helium-burning shell in the star extends from about Mr ≈ 5.5 to 7.3 solar masses. It is also worth noting that stellar burning proceeds through a sequence of core and shell burning. In particular, a given burning stage commences in the center of the star since the temperature is generally highest there. If the energy release is strong enough, the burning drives convection. This draws fuel in from throughout the convective core. Once the burning is complete, the fuel for that burning stage is exhausted throughout the convective core. The star will then contract until the next burning stage begins. Shell burning occurs when the region outside the formerly convective core contracts and heats to the point that the appropriate nuclear fuel can ignite. With these preliminaries in mind, we may trace a star’s evolution from the oxygen isotopes in the star. It is convenient to consider major zones in the stellar ejecta and use the terminology from Meyer et al. (1995) to label these zones by the most dominant elemental abundances, working our way inward from the surface of the star. The H envelope stretches from Mr ≈ 7.8 to 21.6 solar masses. As the star evolves past the main sequence phase of its life, convection begins in the envelope and reaches down to matter that had experienced partial burning of hydrogen

Figure 3. Final mass fractions (after stellar explosion) of the isotopes of oxygen relative to solar in a onedimensional stellar model of an initially 25 solar mass star as a function of Lagrangian mass coordinate Mr. The final stellar mass was 21.6 solar masses, and the abundances are uniform from Mr = 7.8 to 21.6 solar masses (the surface).

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by the CNO cycle. As a consequence, the abundances throughout the envelope look similar to the early stages of CNO burning in Figure 1, with modest enrichments in 14N and 17O. It is also noteworthy that mass loss begins as envelope convection commences and carries away some products of CNO burning prior to the explosion. The He/N zone stretches from Mr ≈ 7.3 to 7.8 solar masses. It is a thin, radiative (nonconvective) shell that was originally part of the convective hydrogen-burning core. Since it largely completed hydrogen burning at the time of stellar explosion, it converted most of its oxygen into 14N, in analogy with the late stages of the single-zone calculation shown in Figure 1, in which all three oxygen isotopes are depleted relative to their initial abundances. The He/C zone stretches from Mr ≈ 5.5 to 7.3 solar masses. This is the part of the star that had experienced convective core hydrogen burning and was burning helium in a convective shell at the time of stellar explosion. In this zone, the 18O is enriched, and the abundances resemble those in the single-zone calculation in the early stages of the burning. It is worth emphasizing that, because of the convective burning, a 14N atom that captures a 4He atom in the burning region is likely to mix out into the non-burning parts of the shell before capturing another 4He. This means that the 18O has to cycle back into the burning region before capturing another 4He to become 22Ne, which lengthens the lifetime of the 18O compared to that in the single-zone calculation. This, in turn, explains why the convective helium shell can be enriched in both 12C and 18O, even though Figure 2 might suggest that 12C builds up only after the 18O is declining. The O/C zone contains material from Mr ~4.5 to 5.5 solar masses. This is matter that was part of the convective helium-burning core but did not partake in convective carbon shell burning. Since this matter completed helium burning, the oxygen here is dominated by 16O. Nevertheless, in this particular stellar model, some 17O and 18O are present due to non-convective burning in this region just prior to the stellar explosion. The O/Ne zone, ranging from Mr ~3 to 4.5 solar masses, is a region of the star that experienced convective carbon shell burning. As discussed above, at the temperatures of carbon burning, 16O can capture 4He nuclei released by the carbon burning reactions. This thereby depletes the 16O left over from the previous convective core helium burning. As is evident from Figure 3, however, this depletion is fairly slight, and the 16O still strongly dominates the oxygen abundances. In the present stellar model, the O/Si zone ranges from Mr ≈ 1.8 to 3 solar masses. This is the region of the star that experienced neon shell burning. As discussed above, the effective neonburning reaction is 20Ne + 20Ne → 24Mg + 16O, which increases the 16O abundance slightly. Finally, inside Mr ≈ 1.8 solar masses in the present model, the star burns its 16O into 28Si and heavier isotopes both in pre-supernova and supernova nucleosynthesis. These regions of the stellar ejecta, the Si/S and Ni zones, are devoid of any oxygen, except for trace amounts produced during the stellar explosion. In broad summary, the ejecta from a massive star is characterized by an 17O-rich envelope, which contains most of the stellar ejecta, a narrow helium shell enriched in 18O, and the inner regions, which are strongly enriched in 16O. It is important to emphasize that, although it is true that massive stars can eject several solar masses of 16O, the massive star ejects all three O isotopes, and the bulk ejecta are not necessarily 16O-rich.

Low-mass stars Apart from mass loss during pre-supernova evolution, which can be significant, massive stars tend to lose all their mass in a single event, namely, the supernova explosion. In contrast, low-mass stars, those with masses less than roughly eight solar masses, lose their mass in periodic episodes during their post-main sequence phases. It is thus necessary to follow the evolution of the star in detail to understand the ejected oxygen yields. We are brief in our

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discussion; for more detailed reviews of the evolution of, and mass loss from, low-mass stars, please see Busso et al. (1999), Herwig (2005), or Straniero et al. (2006). Before discussing the evolution of low-mass stars, it is important to consider the concept of degeneracy in gases. The principal subatomic components of stars are electrons, protons, and neutrons, which all have intrinsic angular momentum quantum number (“spin”) ½, which means they are fermions. By the Pauli Exclusion Principle, no two fermions can occupy the same quantum state. This is not a problem at low densities when there are a vastly greater number of states energetically available than electrons present. When the density increases, however, the number of energetically available states does not greatly exceed the number of electrons. This is the condition of degeneracy, and the result is that electrons may find themselves trapped in highenergy states because lower-energy states are already occupied by other electrons. Because these high-energy electrons carry momentum, they contribute significantly to the pressure of the gas. The remarkable consequence is that even a zero-temperature degenerate gas can have a high pressure. To put this in perspective, consider the pressure at the center of the Sun. The temperature (T) = 1.5×107 Kelvins (K) and the mass density is about 150 g/cm3, which corresponds to a particle density of about n = 2×1025 particles/cm3 in the fully ionized plasma. This gives rise to a gas pressure P = nkBT = about 4×1016 dynes/cm2, where kB is Boltzmann’s constant. There are essentially no degenerate electrons in the center of the Sun, so the degeneracy pressure is zero. In the evolved core of a low-mass star, by contrast, the temperature is roughly 108 K and the mass density is roughly 106 g/cm3. The gas pressure is thus about P = 8×1021 dynes/cm2. At the same conditions, however, the degeneracy pressure is about 2×1022 dynes/cm2. The degeneracy pressure dominates that due to the thermal motions of the gas and can be great enough to support the star against contraction due to gravity. A final introductory point regards burning under degenerate conditions. When nuclear burning occurs, nuclear binding energy is released, which raises the temperature of the local environment. Under non-degenerate conditions, the pressure rises. This causes gas in the local environment to expand, and thus cool. This feedback mechanism can thus lead to steady, stable burning. Under degenerate conditions, however, the pressure is dominated by the degenerate electrons and is thus largely insensitive to the temperature. When burning occurs, the temperature rises but the pressure does not. The rising temperature causes the nuclear burning to proceed at a more rapid rate, which leads to a further increase in temperature and even faster burning. This is a runaway condition that stops only when the temperature is so high that the degeneracy is lifted and the normal gas pressure again dominates. Such violent burning can lead to nova outbursts or to Type Ia supernova explosions (see below). The typical evolution of a low-mass star begins with core hydrogen burning during main sequence evolution. Once the hydrogen is exhausted, the center of the star contracts, and hydrogen shell burning commences on top of the hydrogen-exhausted core. The star begins to ascend the giant branch in the Hertzsprung-Russell diagram. As the star ascends the giant branch, it grows in radius, and the outer envelope expands and cools. This increases the opacity, and the envelope becomes convective. The convective envelope grows in extent and eventually “dredges up” material that had experienced CNO processing (the “first dredge up”). This enriches the envelope material in 17O and depletes it slightly in 16O and 18O. The stellar core continues to contract until helium burning ignites. This occurs either under electron-degenerate conditions as a “flash” for stars less than roughly 2 solar masses or under non-degenerate conditions and, hence, more quietly for stars of mass greater than ~2 solar masses. After the core has exhausted its helium, it contracts again and ascends the “asymptotic” giant branch (AGB). At this time, stars more massive than ~4 solar masses experience “second dredge up,” which brings products of hydrogen shell and helium core burning to the surface. The CNO burning products tend to dominate so that, again, the 17O is enriched.

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By this point, the star’s structure consists of an inert, degenerate C/O core, surrounded by a helium-burning shell, surrounded by a hydrogen-burning shell, surrounded by the stellar envelope. The hydrogen- and helium-burning shells burn alternately in a complicated choreography. Most of the time, hydrogen shell burning occurs quiescently. Eventually, however, the temperature and density rise in the helium-rich region between the C/O core and the hydrogen shell. The helium ignites in a thermal pulse, which drives convection within the helium-rich zone and extinguishes the hydrogen-burning shell. The convective envelope reaches down into matter that has experienced helium burning and dredges it up (“third dredge up”) to the surface, thereby enriching it in the helium-burning products, including 16O and 18O, as well as 17O from the hydrogen shell. Once the helium burning has ceased, the newly produced carbon and oxygen settle onto the core, and the hydrogen-burning shell reignites. The thermal pulses help drive the periodic mass loss from the AGB stars. As a typical AGB star progresses through multiple thermal pulses, the C/O ratio in the envelope increases as the helium shell nucleosynthesis and third dredge up tend to preferentially add 12C over 16O. Because the oxygen abundance initially dominated that of carbon, the star began with a C/O abundance ratio less than one. The preferential addition of 12C increases the C/O ratio until it becomes a “carbon” star (C/O > 1). Eventually the pulsations drive off most the star’s envelope, and only a degenerate C/O white dwarf star with a thin hydrogen or helium atmosphere is left behind. The thin atmosphere is due to the remnants of the star’s original envelope, or possibly to accretion from the interstellar medium. So-called “naked” white dwarf stars that show no surface hydrogen or helium have also been observed (Werner et al. 2004). Naked white dwarf stars have clearly lost their entire envelopes. In white dwarf stars, degeneracy pressure is sufficient to hold the star up against gravity, even at zero temperature, so these burnt-out cinders slowly radiate and cool with time. It is interesting that, given good estimates of bare white dwarf cooling rates, one may estimate the age of the galactic disk from the populations of white dwarf stars as a function of their luminosity (e.g., Oswalt et al. 1996). While this picture of stellar evolution is largely successful in explaining surface abundances of low-mass stars, certain observational puzzles present themselves. First, the surface abundance ratio of 12C/13C observed for many red-giant branch stars is lower than predicted by the models. This suggests extra mixing below the conventional convective envelope. Such mixing would bring envelope material into the vicinity of the H-burning shell and allow for some nuclear processing, and then transport it back to the surface. This has been termed “cool bottom processing” (Boothroyd et al. 1995). In addition to helping to explain the carbon abundance puzzle in low-mass red-giant branch stars, cool bottom processing can also explain the low 18 O/16O ratio in certain AGB star atmospheres and presolar oxide grains (Wasserburg et al. 1995; Boothroyd and Sackmann 1999; Nollett et al. 2003) and the carbon and nitrogen isotopic compositions of many presolar SiC grains (Huss et al. 1997; Nollett et al. 2003). Another puzzle is the large surface 17O enrichments in the handful of observations of J type carbon stars which cannot be explained by standard low-mass-star evolution. The answer may lie in so-called “hot bottom burning”. In this scenario, the convective envelope of a star with mass greater than ~5 solar masses extends down to the H-burning shell so that the envelope material itself experiences H burning, heavily depleting 18O and enriching 17O at the stellar surface (e.g., Boothroyd et al. 1995; Lattanzio et al. 1997). As we shall see in the section “Oxygen in Presolar Grains,” one generally invokes hot bottom burning and cool bottom processing to explain the oxygen isotopic ratios in certain presolar oxide and silicate grains as well as in stellar atmospheres. For completeness, it is worth noting that stars with masses in the range of ~8-10 solar masses do not develop degenerate C/O cores; they proceed to core carbon burning. They then eject their envelopes in a manner similar to lower mass stars and leave behind O/Ne/Mg white dwarf stars. As we have seen, the pressure in high-mass stars (stars with mass greater than about

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ten times that of the Sun) is never dominated by electron degeneracy. These stars therefore proceed all the way through silicon burning.

Novae and Type Ia supernovae Novae are thermonuclear explosions that occur when material from a binary stellar companion (main sequence or red giant star) is gradually accreted onto a white dwarf. As the density increases in the outer layers of the white dwarf due to ongoing accretion, nuclear reactions begin to occur. Because these reactions take place under degenerate conditions, a relatively brief thermonuclear runaway occurs, resulting in the optical outburst we know as a “nova.” A small amount of processed accreted matter and even some underlying white dwarf material is ejected. Then the system settles down to accrete more material until the next nova outburst. Although there are many uncertainties in nova modeling, there is reasonably good agreement between nova nucleosynthesis calculations and observed elemental abundances in novae. Novae are not believed to be significant contributors to the galactic budget of most elements; however, they are probably an important source of some rare isotopes, including 13C, 15 N, and of interest here, 17O. Calculations predict ejecta with 17O/16O ratios 25-2000 times the solar ratio (José et al. 2004). The ejecta are also typically enriched in 18O. Nevertheless, He burning in massive stars is the dominant source of the galactic 18O, and novae contribute only a small fraction to the galactic inventory. In some white dwarf-stellar companion binary systems, the central density in the accreting white dwarf can build up to such high temperatures that carbon-burning reactions can ignite. Like the nova outburst, such burning occurs under degenerate conditions but, given the higher temperature sensitivity of carbon-burning reactions compared to those of the CNO bi-cycle relevant for novae, this burning is much more violent. The result is the complete thermonuclear disruption of the white dwarf in what is most likely a Type Ia supernova. Because of the high temperatures attained in a Type Ia event, much of the matter reaches nuclear statistical equilibrium, which means that such supernovae are major contributors to the Galaxy’s supply of irongroup isotopes. Nuclear statistical equilibrium does not favor oxygen, however, so the inner parts of Type Ia’s do not synthesize much oxygen. In the outer layers of Type Ia supernovae for which the burning front is subsonic, considerable 16O can be produced, principally by explosive carbon and neon burning. Nevertheless, given the relative infrequency of Ia events, they probably only contribute at best a few percent to the Galaxy’s inventory of oxygen (e.g., Thielemann et al. 1986). Confirmation of this basic picture comes from the fact that very old, low-metallicity stars show O/Fe ratios greater than the solar value (e.g., McWilliam 1997). Such stars formed early from gas that inherited the ejecta from massive stars that lived and died in our Galaxy’s first few million years. Massive stars produce a higher ratio of oxygen (16O) to iron than that present in the Solar System, which explains the high O/Fe ratio in low-metallicity stars. Type Ia supernovae, on the other hand, did not start occurring until the first white dwarf stars formed, many tens to hundreds of millions of years after the Galaxy formed. Since Type Ia supernovae produce much iron with little oxygen, their ejecta lowered the O/Fe ratio in the interstellar medium from its high value in the early Galaxy to the solar value by the time of the Sun’s birth.

CHEMICAL EVOLUTION OF THE ISOTOPES OF OXYGEN A fundamental distinction in nucleosynthesis theory is between primary and secondary isotopes, and this distinction has important consequences for the evolution of the oxygen abundances in the Galaxy. A primary isotope is one that can be produced from a star initially composed only of hydrogen and helium. A secondary isotope is one that can only be made from pre-existing seed nuclei. These pre-existing seed nuclei come from previous generations of stars, hence the name secondary.

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The oxygen isotopes provide clear examples of this distinction. 16O is a classic example of a primary isotope. A star composed initially only of the principal products of Big-Bang nucleosynthesis (1H and 4He) could produce 16O. It would do this by converting its initial 1 H into 4He via the PP burning chains. When the 4He subsequently burns by the triple-alpha process, the dominant products would be 12C (another primary isotope) and 16O. By contrast, 17 O and 18O are both secondary isotopes, because a star initially composed only of 1H and 4 He could not form them. 17O is made by proton-capture on 16O. This occurs during hydrogen burning, and a star initially composed only of 1H and 4He would have no 16O during its hydrogen-burning stage to capture a proton. Similarly, 18O traces its origin back to 4He capture by 14N in the early stages of helium burning. As evident from previous discussions, CNO burning converts most of the star’s initial C, N, and O isotopes into 14N because of the slowness of the 14N(p,γ)15O reaction. A star composed initially only of 1H and 4He would not have the C, N, or O isotopes necessary to convert into 14N and, hence, would produce no 18O. Figure 4 demonstrates the primary vs. secondary nature of the oxygen isotopes. Shown are the results from a suite of stellar evolution and nucleosynthesis calculations by Woosley and Weaver (1995) for a range of initial stellar mass and metallicity. In chemical evolution studies, metallicity is an astronomical term that refers to the mass fraction of isotopes heavier than helium. Early astronomical studies of galactic abundances and chemical evolution only distinguished between hydrogen, helium, and everything else, and the respective mass fractions were X (hydrogen), Y (helium), and Z (the “metals”). In the Solar System, the mass fraction of “metals” is Z~~0.02, with 16O accounting for half of that component. As Figure 4A shows, massive stars with a metallicity 1/10,000 as large as that in the Solar System eject as much 16O as do stars with solar or even twice solar metallicity. This confirms 16 O’s status as a primary isotope because it shows that the 16O yield from a star is essentially independent of the star’s initial abundance of elements other than hydrogen or helium. The general trend for all metallicities is that the larger the mass of the star, the more 16O it ejects. This reflects the fact that larger stars have larger helium-burning cores; thus, they produce more 16O. The exception to this trend is that some stars in the suite of models with mass greater than 30 times that of the Sun produce less than their lower-mass peers. This is not a nucleosynthetic effect; rather, it is a consequence of the fact that these models experienced significant fallback during their explosion and formed black holes. The black holes swallowed up much of the 16O, which reduced the ejected yield. Figures 4B and 4C show the corresponding ejected yields from the suite of stellar models for 17O and 18O. From these plots it is clear that the larger the initial metallicity of the star, the greater the production of 17O and 18O. It is also interesting that, for a given initial metallicity, the yield is only weakly sensitive to initial stellar mass. As we previously saw, the ejected 17O and 18O are produced in shell burning. The size of convective hydrogen and helium burning cores is strongly dependent on the stellar mass. By contrast, the sizes of convective hydrogen and helium shells are less dependent on the stellar mass. This explains the weaker dependence of 17,18O on initial stellar mass than that of 16O. An important caveat is that the rates for the key proton-capture reactions 17O(p,α)14N and O(p,γ)18F have been revised since the calculations shown in Figures 4B and 4C were run. In particular, the preferred values for these reaction rates have increased (Blackmon et al. 1995; Angulo et al. 1999). These increases cause greater destruction of 17O during CNO burning, thereby leading to a decrease in the expected yield of 17O from massive stars. In particular, massive star models using the new 17O proton-capture rates typically show nearly ten-fold reductions in the yield of 17O (e.g., Rauscher et al. 2002). This, in turn, suggests that other sites, particularly AGB stars with hot bottom burning and novae play a significant, if not dominant, role in the synthesis of this isotope. Nevertheless, as in the massive star models, the production of 17O is secondary in these other sites (e.g., Romano and Matteucci 2003). 17

Figure 4. Yields of O isotopes from stellar models of type II supernovae of different initial metallicity and mass, from Woosley and Weaver (1995). This figure was produced by the online Clemson University Galactic Chemical Evolution (CUGCE) Tool available at http://webnucleo. org. With this tool, the interested user may explore the primary vs. secondary nature of many other isotopes from massive stars.

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With these basics in mind, the galactic evolution of the abundances of the oxygen isotopes is now easy to understand. The high entropy of the Universe allowed primordial nucleosynthesis to produce only 1H, 2H, 3He, 4He, and trace amounts of lithium, beryllium, and boron from the initial soup of neutrons, protons, and electrons. The initial composition of the first generation of stars, then, was devoid of C, N, and O, and as a consequence, produced little 17O and 18O but normal amounts of 16O. As the Galaxy evolved, the succeeding generations of stars formed from interstellar media that had initial compositions enriched in “metals” from previous generations. These stars were then able to produce increasingly large amounts of 17O and 18O. This behavior is evident in Figure 5, which shows a one-zone model of the evolution of the oxygen abundances in the Galaxy. Like Figure 4, Figure 5 was produced from the online CUGCE Tool at http://webnucleo.org (for a description, see Meyer et al. 2001). The model followed a single zone in the Galaxy, used the instantaneous recycling approximation, and employed Clayton’s family of analytic galactic infall models (Clayton 1984). Figure 5 shows the mass fraction of oxygen isotopes in the interstellar gas, normalized to their mass fraction at the time of Solar System formation in the model. The behavior of 16O is distinctly different from that of 17O and 18O in the figure. Because 16O is a primary isotope, each generation of stars ejects about the same number of grams of 16O per gram of mass going into stars; thus, the gas becomes enriched in 16O in a linear fashion. In contrast, 17O and 18O are secondary isotopes; therefore, the number of grams of these isotopes ejected per gram going into stars increases with each generation. This gives rise to the quadratic evolution seen in Figure 5. At the time of Solar System formation, the oxygen isotopes had all reached their solar values in the Galaxy. The galactic chemical evolution presented in Figure 5 is a simplification, given the fact that the model contains only a single zone and uses the instantaneous recycling approximation—the real Galaxy is an inhomogeneous mix of different interstellar phases and stars that have finite lifetimes. Nevertheless, the general trend of increasing 17O and 18O relative to 16O in the evolution

Figure 5. Mass fraction of the isotopes in the interstellar gas in a chemical evolution model as a function of time. The mass fractions are normalized to their values at the time of Solar System formation (roughly 9.5 Ga after Galaxy formation in the present model). On the scale of the plot, t=0 corresponds to the point in time when the Galaxy starts forming. Build up of the mass of the Galaxy occurs over ~2-3 Ga after t=0 by infall of intergalactic gas.

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of the Galaxy is quite evident from astronomical data. In particular, stars that formed recently in the local stellar neighborhood show a lower 16O/17O ratio than the solar value, which indicates the ratio has declined over the last 4.5 Ga, as expected from galactic chemical evolution (e.g, Wilson and Rood 1994). As for 18O, the 16O/18O ratio decreases inward from the Sun’s position in the Galaxy. Since the inner part of the Galaxy is thought to be more chemically evolved than the outer regions, this supports a secondary nature of 18O. Despite this agreement between observations and galactic chemical evolution expectations for oxygen, there are two significant puzzles. The first is that the solar 16O/18O ratio is actually less than that in today’s interstellar medium in the local solar neighborhood (Wilson and Rood 1994). This is not expected, because the 16O/18O ratio in the Galaxy should decline with time. The second puzzle is that the 18O/17O ratio in today’s interstellar medium is about 3.5, which is significantly less than the solar value of 5.2. From Figure 5, one would expect this ratio to be little changed since the time of the Sun’s formation. A proposed solution to the “18O puzzle” is that the Sun formed in an association of highmass stars, a so-called “OB association”. In such a stellar association, the high-mass stars evolve, explode, and self-enrich the cluster. Since the high-mass stars preferentially eject 16O and 18O, the 18O/16O and 18O/17O ratios can grow over the course of a ~10-20 Myr period of the association’s evolution (e.g., Prantzos et al. 1996). This means that the Sun would have formed with higher values for these ratios than the ambient interstellar medium from which the entire association itself formed. D. D. Clayton has proposed an alternative scenario. It is based on his idea (Clayton 2003) for origin of the silicon isotopes in presolar silicon carbide grains (see section “Oxygen in Presolar Grains”) in which the absorption of a metal-poor satellite galaxy (like the Magellanic Clouds) by the Milky Way initiated a burst of star formation. The number of AGB stars formed during this starburst exceeded that formed by normal galactic star formation, so the former contributed the bulk of the SiC in this period. Since the parent composition of these AGB stars was a mix of the evolved Milky Way Galaxy and the metal-poor satellite, the silicon isotopic compositions of the SiC grains from these stars essentially lie on a mixing line between these two initial compositions. In this scenario, the Sun itself formed with a composition relatively close to the metal-poor satellite; thus, most SiC grains have a silicon isotopic composition that looks more evolved than solar. This scenario has interesting implications for the 18O puzzle (Clayton 2004). In particular, immediately after the merger, high-mass stars quickly evolved and enriched the interstellar medium with 16O and 18O but not 17O, which had to wait for the low-mass stars to evolve. The Sun formed during this period of enriched 16O and 18O, with an 18O/17O ratio of 5.2. Over the course of the last 4.5 Ga, however, low-mass stars formed during the starburst returned their 17O-enriched matter and lowered the interstellar 18O/17O ratio to its present value of 3.5. Clayton (2004) points out that a consequence of this scenario would be a correlation between 18 O/16O and 17O/16O and the 30Si/28Si in presolar SiC grains. Oxygen and silicon isotopes in presolar silicates from low-mass stars would also be expected to show these correlations. The current presolar silicate data set (see section “Oxygen in Presolar Grains”) does not show an unambiguous correlation between Si and O isotopes, but there are some technical difficulties with such measurements (Nguyen et al. 2007). While details of this scenario need to be worked out, it is a nice illustration of oxygen at the nexus of nuclear physics, stellar evolution, galactic astronomy, and presolar grains.

OXYGEN IN PRESOLAR GRAINS Presolar grains are rare and small (nm to 10 μm) mineral grains recovered from primitive meteorites and interplanetary dust particles (IDPs; 90% of which are believed to have formed in C-rich AGB stars and only 1% from supernovae. Some recent studies of cold dust in supernova remnants suggest that supernovae are much more prodigious dust producers than previously believed (e.g., Dunne et al. 2003). However, other studies have questioned these results (Dwek 2004; Krause et al. 2004) and the presolar grain evidence indicates that supernovae were relatively minor contributors to dust in the Galaxy at the time that the Solar System formed. Finally, we note that presolar grains are identified by virtue of having highly anomalous isotopic compositions relative to the range of materials known to have formed in the Solar System. However, this is essentially an operational definition, and analytical uncertainty plays a significant role in deciding whether any given grain is demonstrably a stellar condensate. That is, searches for presolar oxides and/or silicates generally involve the measurement of large numbers of grains, and grains whose isotopic compositions are significantly different (e.g., 3σ) from the rest are defined to be presolar. Figure 7 shows the O-isotopic compositions (expressed as δ-values; see the caption or Criss and Farquhar (2008) for the definition) of many of the presolar grains. Also shown as small ellipses near the origin is the range of O isotopes measured with high precision in meteoritic materials. The grey shaded ellipse illustrates schematically the region of the plot for which analytical uncertainty on small grain measurements has excluded identification of presolar grains. Almost certainly, as analytical methods improve, this region will get smaller and grains with less extreme composition will be defined as presolar. However, it is possible that some presolar grains have isotopic compositions similar to the solar values. One source of such grains may have been other young stars in the molecular cloud from which the Solar System formed. Young stars are observed to eject prodigious amounts of materials in bipolar outflows, and the Sun’s stellar neighbors should have had isotopic compositions close to solar values. In any case, we wish to emphasize that establishing the true limits (if any) of presolar grain isotopic compositions is an important topic for future research.

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Figure 7. O isotopic ratios of presolar oxides and silicates, expressed as δ-values: δiO = 103×[(iO/16O)grain/ (iO/16O)Terrestrial − 1 ]. Dotted lines indicate solar isotopic ratios (assumed to be terrestrial). Solid ellipses near origin indicate range of high-precision isotopic measurements of Solar System-derived meteoritic materials. Grey shaded ellipse indicates region in which analytical errors preclude identification of presolar grains. See Fig. 6 for data sources.

CONCLUDING REMARKS We have reviewed in detail the nucleosynthesis and chemical evolution of the O isotopes in the Galaxy. Although the relative abundances of these isotopes are affected by many nuclear processes occurring in many types of stars, it is clear that most 16O and 18O atoms in the Universe were synthesized in massive stars, whereas a major fraction of the 17O probably formed in lower-mass stars and nova ejecta. This basic picture is confirmed by astronomical observations of stars and molecular clouds and by the ability of nucleosynthesis and galactic chemical evolution models to roughly reproduce the composition of the Solar System. Nonetheless, there are significant puzzles, especially the unusual 18O/16O ratio of the Solar System, and many important remaining uncertainties in the details of the stellar models. In this regard, presolar grains provide particular promise for improving our understanding of the origin and evolution of O in the Galaxy. Advances in the sensitivity and spatial resolution of isotopic imaging capabilities of SIMS instruments have made it easier to find and study presolar oxides and silicates in situ, especially those of submicron size. As we have seen above, such advances have already significantly enhanced our knowledge of stellar evolution by constraining mixing in supernova ejecta and cool bottom processing and hot bottom burning in red giant branch stars. And we anticipate study of oxygen-rich presolar matter will continue to reward us with insights into issues such as chemistry in supernova debris (for example, low-density graphite), chemical evolution (for example, the possible origin of the Group 4 oxide grains), and transit times of dust through the interstellar medium and conditions in the Sun’s parent molecular cloud (for example, presolar silicates in IDPs). Such rich rewards are only fitting for oxygen, the king of the strictly stellar synthesized elements.

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ACKNOWLEDGMENTS The authors would like to thank the organizers, particularly Glenn MacPherson, for the stimulating 2005 Oxygen in the Early Solar System workshop in Gatlinburg. BM, LRN and SRM all acknowledge NASA’s Cosmochemistry program for financial support of this work.

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 55-72, 2008 Copyright © Mineralogical Society of America

Oxygen in the Interstellar Medium Adam G. Jensen Center for Astrophysics and Space Astronomy, University of Colorado Boulder, Colorado 80309-0389, U.S.A. (Present address: Goddard Space Flight Center, Code 665,Greenbelt, Maryland 20771, U.S.A) [email protected]

F. Markwick-Kemper Jodrell Bank Centre for Astrophysics University of Manchester M13 9PL, Manchester, United Kingdom [email protected]

University of Virginia, Department of Astronomy, P.O. Box 400325, Charlottesville Virginia, 22904-4325, U.S.A.

Theodore P. Snow Center for Astrophysics and Space Astronomy University of Colorado Boulder, Colorado 80309-0389, U.S.A. [email protected]

ABSTRACT The oxygen that is observed in the Solar System today is a remnant of the interstellar oxygen that was in the dense molecular cloud that collapsed to form the Solar System. While the chemical evolution of the Galaxy has progressed since then, processes in the interstellar medium (ISM) that involve oxygen are relevant to the origins of oxygen in the Solar System. Oxygen in the ISM can be found as neutral or ionized atomic gas and as a constituent of molecular gas, volatile ices, and refractory minerals in dust, with the dominant state depending on the specific environment. The gas-phase abundance of atomic oxygen is well-known in the diffuse ISM that fills most of the Galaxy’s volume, but the state of oxygen in denser environments is poorly understood. The ISM abundances of isotopes of oxygen other than 16O cannot be easily determined due to observational constraints. Oxygen in interstellar dust is primarily found in the form of silicates that are created in evolved stars and then ejected into the ISM before being incorporated into the formation of new solar systems. Some of the important unknowns concerning oxygen in the ISM include the “cosmic” (i.e., total) abundance of oxygen, the abundance of oxygen in dust, and the details of dust grain processing in the ISM.

INTRODUCTION The oxygen found in today’s Solar System originated in the interstellar medium; it was once part of the dense molecular cloud that collapsed to form the Sun and the Solar System. The Solar System has preserved a sample of interstellar medium material that was available ~5 billion years ago, but the chemical evolution of the Galaxy has progressed since then. It is, however, 1529-6466/08/0068-0005$05.00

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worthwhile to study the processes that take place in the interstellar medium today to obtain clues to the origins of oxygen in the Solar System. Oxygen is an important component in both gas-phase interstellar chemistry and models of solid-state interstellar material (dust and ices). The dominant form of oxygen in the interstellar medium (ISM) varies greatly between different environments. We discuss the abundance and evolution of oxygen in the ISM, both in gas-phase atomic and molecular forms and as a component of oxygen-rich solid-state components.

Phases in the interstellar medium Based on gas density, astronomers distinguish several phases in the interstellar dust and gas reservoir (Snow and McCall 2006). The diffuse atomic regions take up approximately ~90% of the available volume, but only ~1% of the mass, in the interstellar medium, while the dense molecular clouds contain the bulk of the mass (~90%), mostly in molecular form, in only ~1% of the volume of the ISM. The remaining intermediate-density material, which accounts for about 9% of the total mass in 9% of the total volume in the ISM, can be divided into diffuse molecular clouds and translucent clouds, based on their density and extinction1, and whether or not carbon is ionized, neutral, or in the form of CO (Snow and McCall 2006). Phase transitions between these phases occur, and dust and gas can cycle between the diffuse, translucent and dense phases. Material can only leave this eternal cycle in star formation processes when it becomes part of a star or a planetary system. Replenishment of the interstellar medium occurs through stellar ejecta; nucleosynthesis in stellar interiors may have altered the original compositions of stars, and the stellar ejecta often no longer resemble the original ISM material. In addition, the conditions in stellar winds are often favorable for additional chemical and physical processing of the stellar ejecta. Oxygen is found in all phases in the interstellar medium, and can be present in gas phase species (atomic or molecular), as well as in solid-state components (dust or ices). Overall, the interstellar medium is oxygen-rich and enough oxygen is available to drive an oxygen-rich chemistry.

Forms of oxygen in the interstellar medium The dominant form of oxygen in the ISM varies with the type of gas cloud. In the densest molecular clouds, oxygen can be found in molecules such as CO, OH, and H2O. Over a very wide range of densities, including both diffuse molecular clouds and diffuse atomic clouds, oxygen is found in its neutral form (O I). In the hot intercloud medium, oxygen is found in various ionized forms, with the most easily detectable and diagnostically important (though not the most abundant) form being O VI. In dust (found to varying degrees in all phases of the ISM) oxygen can be found in silicates such as Mg2(1−x)Fe2xSiO4, oxides such as FeO, and ices such as H2O. Measuring the amount of oxygen in each form presents its own unique challenges. In the next section we summarize the methods of observing oxygen in the gas phase, significant results, and implications for the amount of oxygen in solid-state forms.

OXYGEN IN THE GAS PHASE Measurements of gas-phase oxygen Neutral atomic oxygen. Neutral atomic oxygen (O I) has most of its resonance transitions in the far-UV (below 1100 Å), though two of the most useful lines are found at 1302 Å and 1356 Å. A wide range of instruments have been used to undertake studies of interstellar O I, 1 The degree to which radiation from a background source is extinguished. This is usually expressed as I =I0 e−τ, with I0 the intensity of the background source, I the measured intensity and the optical depth τ as a measure of extinction.

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from Copernicus (e.g., York et al. 1983; Keenan et al. 1985) to the Goddard High Resolution Spectrograph (GHRS) and Space Telescope Imaging Spectrograph (STIS) instruments onboard the Hubble Space Telescope (HST) (e.g., Meyer et al. 1998; Cartledge et al. 2001, 2004; André et al. 2003) to the Far Ultraviolet Spectroscopic Explorer (FUSE) (e.g., Moos et al. 2002; Jensen et al. 2005). The difficulty in measuring any atomic species in the ISM comes from the necessity to measure absorption lines that are either very weak or very strong. If lines are weak enough, then they can be assumed to be free of saturation, and the equivalent width (a measure of total absorption relative to the continuum) of the absorption line is linearly proportional to the column density of absorbing material. Unfortunately, it is difficult to find lines that are both weak enough to meet this requirement yet strong enough to be detected and accurately measured, especially in lines of sight with heavy reddening (due to dust) where the signal-tonoise ratio in the UV range is low. Additionally, some species are expected to be “depleted” (that is, in forms other than the gas phase), and subsequently the gas-phase column density that creates these absorption lines is even smaller. Weak lines are still useful if they have slight saturation, if the velocity structure2 of another atomic species can be assumed and an appropriate saturation correction is applied. Conversely, when very strong absorption lines exhibit “damping wings” as a result of the Heisenberg Uncertainty Principle applied to the transition lifetimes, the shape of the wings is fairly sensitive to the total column density, and the total equivalent width of the absorption feature is proportional to the square root of the column density. In these cases, as long as the spectral resolving power is sufficient to clearly delineate the damping wings, column densities can be accurately measured either through fitting the shape of the wings or by measuring the equivalent width of the absorption feature. However, most elements of astrophysical interest do not have transitions strong enough to exhibit such wings at typical column densities, especially when potential depletions are factored in. Instead, in many cases astronomers are left with lines of intermediate strength, which are relatively insensitive to column density, making it difficult to derive accurate abundances from them, even in some cases where the velocity structure is accurately known. If the only available lines are of moderate strength, the alternative is to construct a curve of growth. A curve-ofgrowth method fits the equivalent widths of multiple absorption lines to a curve that represents the expected values of the equivalent width of a line for a range of combinations of oscillator strength and column density. The free parameter in this fitting process is the column density. The shape of the curve also depends on the line-of-sight velocity structure. Either a certain velocity structure must be specified, or a single-component velocity dispersion can be assumed. In the latter case, the b-value of the dispersion is a second free parameter in the fitting process, with the various b-values creating a family of curves. The accuracy of the curve-of-growth method relies on having a wide range of absorption line strengths that are measured and the validity of the assumed velocity structure (see Spitzer 1978 for a review). Once oxygen column densities have been determined, abundances relative to hydrogen can be determined in diffuse (or denser) clouds by measuring hydrogen through profile fitting of the Lyman-α line (for atomic hydrogen) and the many low-J (rotational) states of H2. The major observational barriers are the spectral type of the background star (where cooler stars may introduce significant contamination of the Lyman-α line for the atomic hydrogen measurement) and the difficulty of obtaining good data quality in the far-UV, where most H2 transitions lie. 2

“Velocity structure” refers to the distribution in velocity space of the absorbing material that is reflected in the profile of an absorption line. This structure includes the number of cloud components and their relative column densities; the velocity offsets of the clouds relative to Earth; and the velocity dispersion of each cloud (a “b-value” quantifies a dispersion that is assumed to be Gaussian).

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The early Copernicus studies of York et al. (1983) and Keenan et al. (1985) relied primarily on the weak 1356 Å line of oxygen, using a curve-of-growth of neutral nitrogen to infer any possible saturation corrections. Neither study found any evidence of correlations in the oxygen abundance with the total hydrogen column density. Meyer et al. (1998), André et al. (2003), and Cartledge et al. (2001, 2004) studied O I abundances with the GHRS and STIS instruments onboard HST. The Meyer et al. (1998) study examined 13 lines of sight and did not find any statistically significant variation of O/H with respect to N(Htot), the total line-of-sight column density of atomic and molecular hydrogen, measured in particles per unit area; nH, the average volume density of hydrogen in a line of sight, equal to N(Htot)/r, where r is the line-of-sight pathlength; or the fraction of hydrogen atoms in molecular form, f(H2)=2N(H2)/[2N(H2) + N(H I)], where N(H2) and N(H I) are the molecular and atomic hydrogen column densities, respectively. The more extensive study by Cartledge et al. (2004), however, found that O I depletion from the gas phase did increase with the average volume density of hydrogen, nH. Cartledge et al. found that O/H is approximately 100 parts per million (~25% of the total observed gas-phase O/H ratio) smaller in lines of sight with larger nH. Cartledge et al. interpreted this as representative of two phases in the ISM (warm and cold), with a transition at nH ≈ 1 cm−3. Jensen et al. (2005) used the FUSE satellite to probe reddened lines of sight with large fractions of molecular hydrogen. FUSE does not provide information on the weak 1356 Å line or the potentially damped 1302 Å line, but does provide information on a wide range of absorption lines of intermediate strength that are beyond the spectral range of HST, and at greater sensitivity and/or resolution than past instruments such as Copernicus. The Jensen et al. (2005) study did not find statistically significant correlations of O/H with any important line of sight parameters, but could not rule out possible correlations within the large errors. However, a small subset of lines of sight with both HST and FUSE data available (and subsequently smaller systematic and statistical errors) did show the hint of a correlation of decreasing O/H with an increasing ratio of total to selective extinction, RV ≡ AV/EB-V. Typical interstellar grains will preferentially cause extinction at shorter wavelengths, and EB-V is often correlated with the total column density of typical interstellar dust. Larger grains, however, will cause extinction in visible light in a non-preferential manner (i.e., grey extinction). Therefore, increasing values of RV, a measure of significant extinction in the visual band (AV), but relatively little preferential extinction at shorter wavelengths (EB-V), are thought to correlate with increasing grain size. Other ionization states of atomic oxygen. The ionization potential of O I is nearly identical to that of H I (13.618 eV and 13.598 eV, respectively). This is convenient, because combined with the fact that the ionization of O I relative to H I is also governed by a charge-exchange reaction (Field and Stiegman 1971), we can draw the simple conclusion that in regions dominated by H I and H2 (from diffuse atomic clouds to dense molecular, star-forming clouds) the vast majority of oxygen will be either neutral or found in a molecular or solid-state form. More specifically, a gas cloud in collisional equilibrium at temperatures up to 104 K should have over 99% of its oxygen in the form of O I (Sutherland and Dopita 1993). Observing other ionization states of oxygen poses significant observational difficulties in addition to their relatively small fractions. O II, O VII, and O VIII do not have ground-state transitions with wavelengths in the UV, and the few such transitions of O III, O IV, and O V have such small f-values that they are unlikely to be observed in interstellar absorption (Morton 2003). These elements are, however, observed in emission. For example, O III is commonly observed in planetary nebulae. Abundances relative to certain other elements can be accurately determined in these cases, but direct measurements of the O/H ratio are not possible in such regions, where the hydrogen is largely ionized. Again, however, O III does not play a significant role in diffuse clouds.

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While it does not compose a significant fraction of the total oxygen in the Galactic ISM, and is never the dominant form of oxygen (Sutherland and Dopita 1993), the lithium-like ion O VI is an important diagnostic of the hot intercloud medium, as well as gas in the Galactic halo and the intergalactic medium. This is due in part to its detectability, and also because its large ionization potential, which is above the ionization edge of He II, means that its creation is dominated by collisional excitation rather than photoionization. O VI is observed through a pair of relatively strong transitions at 1032 Å and 1038 Å, in a spectral range detectable with FUSE (and past instruments such as Copernicus) but not HST. The broad thermal distributions of O VI that are typically seen imply that O VI absorption is not significantly saturated and O VI column densities can be determined directly and accurately. The study of O VI is important for studying gas at the interface between hot and cold gas; at collisional equilibrium temperatures of ~3×105 K, O VI reaches its peak ionization fraction of ~20% (Sutherland and Dopita 1993). However, most O VI is observed in non-equilibrium regions, such as shock fronts, and traces cooling and heating, itself being an important coolant. Oxygen in gas-phase molecules. The dominant form of oxygen in molecules is carbon monoxide. All isotopologues3 of CO have been detected in the interstellar medium through radio emission (from rotational transitions) except for 13C17O (Morton and Noreau 1994), though detection through ultraviolet absorption (from electronic transitions) is also common. Sheffer et al. (2002) and Federman et al. (2003) have presented some of the most recent observations of many different forms of CO in UV absorption with the STIS and GHRS instruments onboard HST. These results show just how small the abundance of molecules is in the diffuse interstellar medium. For example, the 12C16O column density for the line of sight toward X Persei found by Sheffer et al. (2002) is a factor of ~50 smaller than the column density of gas-phase atomic oxygen found by Jensen et al. (2005). Other isotopologues of CO are even less abundant. The study of CO is particularly important for the cooling of dense molecular clouds; CO is collisionally excited and then radiates at wavelengths where the cloud is optically thin. In these cases, the CO abundance jumps up dramatically; as much as one-sixth of the total carbon abundance may be in CO. Other molecules of potential importance include H2O, OH, O2, and H2CO (formaldehyde). Goldsmith et al. (2002) presented a tentative detection of O2 in a high-velocity outflow seen in the line of sight toward ρ Ophiuchi, later confirmed by Liseau et al. (2006). The abundance relative to H2 is approximately 10−5. In general, though, limits have been placed on the O2 abundance to be less than 10−7 relative to hydrogen. (See Goldsmith et al. (2002) for further discussion of a scenario explaining the large abundance in this line of sight.) Aggregate measurements of oxygen in gas and dust. Measuring oxygen in all of its forms is possible in some cases. The inner K-shell absorption edge of oxygen can be detected through X-ray spectroscopy, measuring oxygen along the line of sight in all of its forms— atomic, molecular, and in dust. Fine structure can further reveal whether the oxygen is in the gas form or tied up in solids. The results in the case of X Persei (Cunningham et al. 2004) are consistent with the results of gas-phase oxygen abundance studies and the requirements of dust models based on UV extinction. To summarize, the gas-phase atomic oxygen abundance is relatively constant over a wide range of conditions. However, the evidence from Cartledge et al. (2004) suggests that oxygen depletion from the gas phase increases somewhat with increased average gas density, and the 3 Isotopologues are chemicals with different isotopes of the same atomic constituents. An isotopomer refers to chemicals with the same isotopes of atomic constituents but different atomic arrangements (isotopomer is a contraction of isotopic isomer). Though isotopomer is commonly applied to the various forms of CO in astrophysical literature, isotopologue is the correct term.

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evidence from Jensen et al. (2005) hints that grain size is related to an increase in oxygen depletion. However, astronomers still do not have a perfectly formed idea of how oxygen varies in the most extremely reddened environments, or even in translucent clouds (a phase of the ISM that has long been postulated but for which direct evidence is sparse; see further discussion in Rachford et al. 2002). The Cosmic Origins Spectrograph (COS), which will be installed on HST as a part of Servicing Mission 4 (currently projected for August 2008), will have several times the sensitivity of STIS, but at somewhat lower resolution. COS will be able to probe oxygen column densities through detection of the weak 1356 Å line of oxygen even at AV > 5 mag, much more reddened than has ever been observed at high S/N in the ultraviolet. This will allow probing of the gas-phase oxygen abundance in the much denser environments that are not yet fully understood.

Isotope measurements from gas-phase oxygen and carbon monoxide Atomic isotope shifts. Attempting to measure relative isotopic atomic abundances of elements in the interstellar medium is very difficult. In the case of single-electron (hydrogenlike) atoms, the difference in the reduced mass of the electron and the nucleus, μ ≡ memnucleus/ (me+mnucleus), is what provides the change in the energy and subsequent shift in wavelength of a given transition between two different isotopes. While this can be significant for the shift between hydrogen and deuterium (~3×10−4 of the transition energy, corresponding to the absolute value of the fractional shift in the wavelength), it is increasingly small with increasing atomic mass. For example, the difference in the transition energy due to the different reduced masses of 16O and 18O is ~4×10−6 (assuming hydrogen-like O VIII). For species with more than one electron, the shifts depend on more than just the reduced mass, as the total momentum of all electrons in the system must be considered. The resulting energy shifts are difficult to calculate (see Clark 1984, who calculates shifts for many atomic species but not O I); however, they are unlikely to be substantially larger than the above approximation (which is equivalent to a shift of ~1 km/s for far-UV lines), which puts it at the threshold of the resolution of even a precision instrument such as STIS. No study of interstellar atomic oxygen has turned up clear evidence of other isotopes of atomic oxygen, due to a combination of the following limitations: (1) The inherently small shift between isotopes in the energy for a given transition, especially compared to the resolution of available instruments; (2) The broadening of interstellar profiles (~3-20 km/s for oxygen; see Jensen et al. 2005), which is a combination of thermal broadening, multiple clouds (with a range of velocity offsets) in the line of sight, and in the case of strong, damped lines, “natural” broadening; and (3) The relatively small abundances of isotopes other than 16O. Isotopologues of carbon monoxide and other oxygen-bearing molecules. The situation with molecules is more promising. The difference in isotope mass of the atoms in a molecule influences the energy of its rotational and vibrational states, in some cases splitting the energies of electronic transitions by resolvable amounts. Further helping matters is the fact that molecules tend to form in cold environments with small thermal widths, resulting in absorption features that are more easily resolved. Radio observations of rotational transitions can be used to detect various isotopologues; in fact, as mentioned above, almost all variations of carbon monoxide have been measured in the ISM through radio observations. The ratio of 12CO to 13CO does seem to show a spatial variation that can be correlated with stellar evolution that turns 12C into 13C (Langer and Penzias 1990). However, the relative abundances of the various isotopologues of CO are unlikely to be good proxies for the relative isotopic abundances of oxygen. Both Sheffer et al. (2002) and Federman et al. (2003) discuss

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the fractionation of CO isotopologues. In all combinations, the fractionation is more severe than the relevant isotopic ratios; e.g., 12C16O/12C18O is found to be ~3×103, compared to the Solar System 16O/18O ratio of ~500 (Lodders 2003). This is because the less abundant isotopologues remain optically thin and therefore susceptible to photodissociation, while the more abundant isotopologues may begin to become optically thick and self-shielded from photodissociating radiation. The impact of slightly different RMS velocities—due to different isotope masses—on molecular formation rates may also play a small role in the difference between the isotopologue ratios of CO and the isotopic ratios of C and O. The isotope 17O is less abundant than both 16O and 18O. Penzias (1981) attempted to constrain the C18O/C17O ratio, and found a value of ~3.5, consistent within the errors at all measured galactocentric radii. Wouterloot et al. (2005) found a ratio of ~4.1, while Ladd (2004) found ratios of ~4.0 or ~2.8, depending on the model used to interpret the line intensities. All of these measurements, however, are somewhat smaller than the Solar System 18O/17O ratio, ~5.5 (Anders and Grevesse 1989; Lodders 2003). These low ratios cannot be due to differences in self-shielding effects, because C18O is more likely to be self-shielded than C17O, which would increase the observed C18O/C17O ratio (Ladd 2004); deep within molecular clouds, both are likely to be well-shielded. Vastly different 18O/17O ratios are also observed in the Large Magellanic Cloud (LMC) and two nuclear starburst galaxies, implying that metallicity4 might play a role in determining this ratio (Wouterloot et al. 2005 and references therein). Studies of isotopologues of OH have also been used to explore the isotopic abundances of oxygen. Bujarrabal et al. (1983) found ratios of 16OH/18OH to be ~370 and ~440 in two different cloud components toward Sgr B2, with a lower limit of ~200 in a third cloud component. Within the errors, these derived ratios are in reasonable agreement with the Solar System ratio, and Bujarrabal et al. suggest that previous interstellar results with lower ratios are likely the result of excitation of the more abundant 16OH. The 18OH/17OH ratio is found to be 3.6±0.5, however, again lower than the Solar System value of ~5.5 and in agreement with the results from CO isotopologues. Polehampton et al. (2003, 2005) also used isotopologues of OH to explore the interstellar isotopic abundances of oxygen, finding 16OH/18OH ratios also roughly consistent with the Solar System 16O/18O ratio (17OH was also detected, but 18OH/ 17 OH was not independently derived). They also found no evidence of a gradient in 16O/18O with galactocentric distance, but their results also do not conclusively rule one out. The lack of a gradient in the isotope ratio would be in contrast with several other studies that show a low value in the galactic center (e.g., the compilation of 16O/18O ratios determined from H2CO isotopologues in Wilson and Rood 1994). For further discussion, including information about the uncertainties of the various studies, see Polehampton et al. (2005) and references therein.

Inferring gas-phase depletions of oxygen While there is some consensus on the gas-phase abundances of oxygen in environments that are not too extreme, what is less clear is the interpretation of the gas-phase oxygen measurements as it relates to forms other than the gas phase. Typically, derived abundances of gas-phase oxygen are compared to a certain standard (usually the solar oxygen abundance) and the “missing” oxygen is inferred to be the total amount of oxygen in molecules and dust. The problem with this method is two-fold. First, the appropriate standards are not well known. For example, measurements of the solar oxygen abundance vary by nearly a factor of two, even within the last decade. Secondly, and perhaps more importantly, there is reason to doubt that the solar oxygen abundance is the correct standard, or that any such standard exists at all. What is the solar oxygen abundance? Large ranges are found when considering photospheric models. Anders and Grevesse (1989) derived a value for the solar O/H of 853 4 In astrophysics, the term metallicity is often used to mean the mass fraction contained in “metals,” i.e. elements heavier than helium, in an astrophysical environment.

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parts per million (ppm), while the more recent 3-D photospheric models of Holweger (2001) and Asplund et al. (2004) put solar O/H at 545 and 460 ppm, respectively. More recently, Ayres et al. (2006) have argued for a solar O/H ratio of 700 ppm based on chromospheric CO as a tracer of the total oxygen abundance. For a time, it was thought that many elemental abundances (e.g., carbon, oxygen, and krypton) in the ISM were very nearly 2/3 of the solar value, but this is no longer a widely held opinion, based largely upon the revisions to, or at least the uncertainty in, the solar values of the carbon and oxygen abundances (the disparity with krypton may still exist, however; see Sofia and Meyer 2001). Instead of the solar photospheric elemental abundances, perhaps protosolar abundances— i.e., “Solar System abundances”—are better estimates of the cosmic abundances. Estimates for abundances in the protosolar disk are most commonly derived from meteorites, but this is problematic for oxygen because it readily forms gaseous compounds, and it is therefore almost certain that there was much oxygen in the protosolar disk that did not condense into rocks. Protosolar abundances, such as that for oxygen, are derived by applying a systematic factor to the photospheric abundances; this systematic factor is based on understanding of condensation in the protosolar disk, including the meteoritic abundances of some of the heavier elements. Two examples of this method, Cameron (1973) and Lodders (2003), derive a range of protosolar O/H ≈ 575-700 ppm, due in part to the different photospheric abundances used. For further discussion on the protosolar oxygen abundance and related issues, see the chapter in this volume (Davis et al. 2008) Perhaps our Solar System and Sun are unique and a better oxygen abundance standard is the oxygen abundance in other stars. Studies such as Snow and Witt (1996) and Sofia and Meyer (2001) have examined this issue. Specifically, Snow and Witt found O/H = 380 ppm for field B stars, and O/H = 490 ppm for cluster B stars and field F and G stars. For B stars, Sofia and Meyer found a weighted average of O/H = 350 ppm, and an average O/H of 445 ppm for F and G stars (not weighted, as it was based on a very small sample). The observed gas-phase ISM abundances of oxygen rule out the field B star abundances of Snow and Witt (1996) and the B star abundances of Sofia and Meyer (2001) as potential standards. Various theoretical models of interstellar dust, such as those cited in Snow and Witt 1996, attempt to explain the nature of interstellar extinction, including features such as the 10- and 18-μm bands believed to be from silicates; these models generally require >120 ppm of oxygen relative to hydrogen. Combined with the ISM averages of André et al. (2003) and Cartledge et al. (2004) of gas-phase O/H≈400 ppm (the “warm ISM” in the Cartledge et al. case), even the F and G star abundances fall a little short of explaining the total gas- and dust-phase oxygen in the ISM. What might cause the oxygen abundance in the ISM to differ from the oxygen abundance of the Sun and other stars? It is important to note that the spread in stellar abundances is greater than the spread in ISM abundances between various lines of sight, leaving the question of how such varied stars could form from a well-mixed ISM. A likely answer is that processes occurring during stellar formation heavily influence stellar abundances. Snow (2000) discusses two possibilities that may leave stars metal-poor: ambipolar diffusion and sedimentation. Ambipolar diffusion would magnetically exclude refractory elements, possibly resulting in stellar abundances that are smaller by a factor of 2-3. Sedimentation is the process by which large grains form in a protostellar disk but do not accrete onto the star, decoupling refractory elements from the gas, and resulting in metal-poor stars. Another process that could potentially factor into producing stars that are metal-poor relative to the ISM is the photoablation of protostellar disks. If protostellar disks are destroyed soon after stellar formation, then the late infall of potentially metal-rich material is disrupted. Additionally, if there is some late infall of gas (either metal-rich or metal-poor) after stellar convection has begun, this gas may remain in the photosphere, distorting observed versus true stellar abundances.

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OXYGEN IN INTERSTELLAR DUST On first sight, with an overall gas-to-dust ratio in the interstellar medium of 100-200, dust may not appear to be an important contributor to the total mass budget of oxygen. However, almost all of the gas in the ISM is actually in the form of hydrogen or helium, while these two species rarely compose a significant fraction of the dust mass. Thus, for the remaining elements, including oxygen, the solid phase appears to be important. In fact, at a metallicity of Z=0.02, and with a gas-to-dust ratio of 100, about half of the metals—including oxygen—are contained in the solid phase. Most of the oxygen that is found in the solid state in the diffuse interstellar medium is found in silicates, while some may also be present in metal oxides (Whittet 1984). In the dense interstellar medium, the remaining gas phase oxygen might condense out in the form of ices, in particular H2O and CO2, as evidenced by infrared spectroscopy of protostars (Gibb et al. 2004). Ices are much more volatile than silicates and are therefore less likely to contain a record of their formation and processing history. In addition, the phase in which ices can be observed astronomically spans only a short period of time in the total evolution cycle of dust. We will therefore limit the discussion on interstellar oxygen in the solid state to the silicates and will discuss such properties as composition and lattice structure, in the context of the evolution of silicate dust.

Solar System silicates The most recent phase in the evolution of silicate dust, Solar System silicates, is discussed extensively in other chapters in this book. The most primitive silicates found in the Solar System today are believed to be the GEMS (Glasses with Embedded Metals and Sulfides; Bradley 1994), which are present in some interplanetary dust particles (IDPs), and the silicates that have been locked up in comets. Evidence for crystallinity is found in the 8-13 μm spectroscopy of both short and long period comets (Hanner et al. 1994; Honda et al. 2004). Ground-based observations of the Deep Impact encounter with comet 9P/Tempel 1 show that the silicates observed before and after the impact were mostly amorphous, and predominantly in the form of large grains, while the dust plume caused by the impact contains an abundance of smaller and crystalline silicates (Harker et al. 2005). This is interpreted as evidence that cometary silicates initially are significantly crystalline, and that these silicates gradually become more amorphous upon passing the Sun several times. A full 2-45 μm spectrum, obtained for Oort Cloud comet Hale-Bopp (Crovisier et al. 1997) supports this view, by showing strong forsterite features in the 20-40 μm range. Estimates for the degree of crystallinity of the silicates in Hale-Bopp typically fall in the 20-30% range (Brucato et al. 1999; Galdemard et al. 1999; Wooden et al. 1999; Hayward et al. 2000; Harker et al. 2002; Bouwman et al. 2003), but more recently Min et al. (2005) have convincingly shown that the degree of crystallinity is probably closer to 7.5%. Because very little processing occurs inside comets in the Oort cloud, the crystalline fraction in long-period comets observed today essentially reflects the composition of the silicates in the comet-formation zone in the planet-forming disk around the young Sun. The fact that cometary silicates are more crystalline than GEMS and than interstellar silicates (Kemper et al. 2004), indicates that crystallization occurs in the planet-forming disk itself, or perhaps in the dense molecular cloud that pre-dates the Solar System. The process of crystallization of the amorphous interstellar silicates, as they get incorporated in the Solar System, is probably best studied in other planetary systems in formation.

Silicates in circumstellar environments of young stars One can distinguish two types of circumstellar disks around young stars. Pre-main sequence stars often have relatively stable and massive primordial planet-forming disks around them, while young main-sequence stars can exhibit debris disks in which the sources of dust are planetesimal collisions. When the era of planetesimal collisions ends, the disk rapidly clears

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out due to Poynting-Robertson drag and radiation pressure. Debris disks are seen around stars up to an age of about 400 million years (Habing et al. 1999). The spectral properties in such disks are consistent with the planetesimal origin of the grains. Recent Spitzer observations have shown that typical debris disks are populated by large (Jura et al. 2004) and crystalline (Beichman et al. 2005) silicate grains. The planetesimals from which these grains originated may bear resemblance to either asteroidal or cometary bodies (Beichman et al. 2005), both of which, in our Solar System, are known to contain crystalline silicates. While crystallization in dense planet-like bodies occurs due to the heat of formation, the presence of crystalline silicates in loosely packed cometary bodies is not entirely understood. Formation heat never reaches sufficiently high temperatures to anneal the silicates, so the crystalline fraction observed in comets reflects the composition of the circumstellar environment where the comets are formed. Indeed, crystalline silicates are observed in the planet-forming disks around pre-main-sequence stars (e.g., Waelkens et al. 1996). A range of crystalline fractions is observed—always the Mgrich end-members forsterite and enstatite—however, it rarely exceeds a few percent of the total silicate mass. The crystallinity appears to be more enhanced for larger overall grain sizes in the circumstellar disk (Bouwman et al. 2001). Although the evolutionary status of the disk may correlate with the particle size of the silicate grains, a clear correlation between stellar age and crystallinity remains to be confirmed (Bouwman et al. 2001). It is clear that energetic processing is required to build up the observed degree of crystallinity (a few percent), but it remains unknown how this level of crystallinity can be achieved in the comet-forming region, where the radiative heating from the young star is not sufficient to heat the grains to crystallization temperatures. At least two different mechanisms have been proposed. First, radial mixing of grains from the warm inner radius of the disk toward the outer regions has been suggested as a mechanism to transport crystalline material to the comet-forming region (Bockelée-Morvan et al. 2002). This theory is supported by the recent results from the Stardust mission, where the presence of refractory Ca-, Al-rich grains among the cometary particles, as well as the low abundance of grains with presolar isotopic compositions, in particular the 18 O/16O ratio, suggests large-scale radial mixing (Brownlee et al. 2006; McKeegan et al. 2006; Zolensky et al. 2006). Alternatively, it might be possible that shocks propagating through the pre-solar nebula sufficiently heated the grains in situ to cause crystallization, even at distances of 10 AU (Harker and Desch 2002). Recent interferometric observations of three pre-main sequence stars clearly show a gradient in crystallinity in the circumstellar disk, with the highest crystalline fractions near the inner radius of the disk (van Boekel et al. 2004), thus providing supporting evidence for the radial mixing theory.

Dust properties in the interstellar medium In all three phases of the interstellar medium (ISM), silicates are the most important depletion sink for oxygen. Studies of silicate properties have concentrated mostly on the diffuse interstellar medium however, which contains only 1% of the dust mass in the ISM, although it occupies about 90% of its volume. The properties of the dust composition in the denser phases of the interstellar medium have received less attention. In particular, silicate properties in young stellar objects (YSOs)—the solar nebula analogs—are relatively unexplored. Kessler-Silacci et al. (2005) present a sample containing nine approximately solar mass (M~) and three high mass (M > 8 M~), deeply embedded YSOs, showing silicate absorption features in the 8-13 micron wavelength range. They find that the silicates toward these protostars reveal an amorphous lattice structure, and only in more evolved objects, where the silicate feature appears in emission, can crystalline forsterite be observed. Demyk et al. (1999) have set an upper limit of 1-2% on the degree of crystallinity in the silicates towards two massive protostars, and find that the shape of the silicate absorption feature suggests an amorphous pyroxene composition. Early observations of the silicate profile towards the stars in the Trapezium cluster suggested an amorphous lattice structure for the silicates in the diffuse interstellar medium (Forrest et al.

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1975). The shape of this silicate feature has been used to construct the widely used optical constants for interstellar silicates (Draine and Lee 1984), under the assumption that silicates in the ISM are completely amorphous. In 2001, Li and Draine investigated the composition of interstellar dust from infrared emission, and were able to set an upper limit of 5% to the degree of crystallinity. More recently, the silicates in the 8.5 kpc-long line of sight towards the galactic center were shown to be 1 μm depth), where other surface-correlated components such as the oxide layer (likely acquired after the recovery of the sample) were mostly removed.

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for solar wind implantation was a key indicator of the solar wind, and the corresponding carbon and nitrogen compositions at those depths were inferred to be solar. IHNC06 chose a recent soil because it offered the best prospect for recent exposure to the solar wind, and would therefore be most analogous to the exposure of synthetic materials on the Genesis mission (see below). The high proportion of grains with high track densities and the 36Ar concentration attest to this exposure. HC05 argue that the bulk soil contains many components of diverse origins and therefore the young soils are compromised. However, the issue is one of ascertaining contributions on a grain-by-grain basis, such as was carried out by HC05 for carbon and nitrogen. In the case of carbon and nitrogen, the isotopic compositions are indeed at the appropriate depth and associated with low D/H. For oxygen isotope measurements, the oxide rims potentially compromise the analyses, with that oxygen coming from an unknown source. For this reason, IHNC06 chose not to analyze this material. HC05 and IHNC06 attempted to perform D/H isotopic measurements on the metal grains, but the samples have D/H at terrestrial levels and this is interpreted as a result of diffusion from the terrestrial atmosphere after the samples were returned to Earth. The depths of the solar signatures in the two studies are quite different and require quite different sources. SRIM3 calculations show that for solar wind energies, the depth of penetration is less than 50 nm for oxygen implantation into iron metal. Even for extreme solar wind energies (1000 km s−1), oxygen ions are stopped by 150 nm. These results are broadly consistent with the depth of oxygen measured by IHNC06, given some degree of uncertainty in the sample geometry and sputtering calibration, although the abundance of the oxygen at deeper depths is higher than expected based on the relative proportions of modern normal solar wind velocities (400 km s−1, corresponding to an energy of a few keV/nucleon) to extreme solar wind velocities (1000 km s−1). Implantation of the oxygen into the 79035 iron metal grains is attributed to solar energetic particles (HC05). SRIM calculations show that for energies on the order of 10 MeV/nucleon, penetration depths for oxygen into Fe metal are indeed on the order of 3 μm. There appears to be an issue, however, regarding the relative contributions of SW and SEP in lunar samples. The solar wind contribution should be much larger in lunar materials compared to the rare events responsible for SEP. It appears that SW is preferentially lost relative to the SEP contribution, however; for example, the SEP contribution of noble gases composes 10-20% of neon, argon, and krypton in lunar soil 68501 (Becker and Pepin 1994). Possible mechanisms for this loss include ablation and diffusion. The SW contribution in the 79035 metal grains is masked because of the presence of the oxide layers. Only when the oxide contribution declines with depth does the SEP component become apparent. Its presence in only very few of the grains suggests that not all grains experience the same exposure conditions. The oxygen isotopic compositions of two metal grains from 10084 reveal high 17O and O at depths consistent with solar wind implantation. Two other grains had oxide layers and these analyses were discontinued after isotopically normal Δ17O values were obtained. Ireland et al. (2007) report additional analyses of metal grains from the Apollo 16 and 17 sites. In lunar soil 61141, several grains with negligible oxide layers and several with thicker oxide layers were analyzed. None of these grains has nonterrestrial Δ17O. In lunar soil 78481, oxide layers were apparent and little deviation from terrestrial Δ17O was noted. These data suggest that the presence of implanted oxygen may be specific to individual grains within individual soils. 18

To elucidate the discrepancy currently observed in solar oxygen studies of lunar samples, it is imperative to make clear the inventory and time variations of extraselenial oxygen fluxes supplied to the lunar surface. 3

SRIM, an acronym for Stopping and Range of Ions in Matter, is a widely used computer program and is available at http://www.srim.org.

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The implications of the two different reports of the solar oxygen isotopic composition, which are at present difficult to reconcile, are treated separately. Implications of an 16O-rich solar composition. In this section, the value of HC05 is adopted for discussion, whereas the discussion based on the result of IHNC06 is done in the following section. In Table 1, the solar isotopic compositions are compiled, and are compared with the isotopic compositions of the terrestrial upper mantle, as an example of a large and well-studied planetary reservoir possibly representing, or at least not far from, the average of solid materials in the inner Solar System. The δsolar values, the deviation of the average planetary isotopic composition relative to the solar composition, are calculated from these values. The true δsolar values may be even larger, since the solar isotopic ratios obtained from the lunar samples are upper limits. For example, if we adopt the Jovian atmospheric isotopic composition of nitrogen (δ15Nair = −380 ± 120‰), as the solar value instead of the upper limit value obtained from lunar samples (δ15Nair < −240 ± 25‰), the δ15Nsolar value will be 2 times larger than the one given in Table 1. Here, to enable comparison of the δsolar values among carbon, nitrogen and oxygen, we adopt the lunar-based values. The δDsolar value, however, is calculated based on the Jovian atmospheric value (D/H = (2.0 ± 0.35) × 10−5; Geiss and Gloeckler 2003), representing the initial solar value before the Sun destroyed its deuterium. This is because the protosolar D/H ratio is not preserved in the present (main sequence) Sun or in lunar soil samples, whereas the solar photosphere is representative of the whole Sun for carbon, nitrogen and oxygen isotopic composition (Kallenbach et al. 2003). In Table 1, there are large differences between the planetary and solar isotopic compositions. It was argued for a long while that the observed range of oxygen isotopic composition among meteorites is primarily due to contributions to the planetary solid materials of presolar components with strikingly different isotopic compositions (e.g., Clayton 1993). However, the available data on oxide and silicate presolar grains do not favor this hypothesis. The majority of presolar oxide and silicate grains are enriched in 17O but depleted in 18O compared to solar composition (reviewed by Zinner 2007, and Meyer et al. 2008). The major population of presolar SiC grains is enriched in 14N (Zinner 1998). Nanodiamonds separated from carbonaceous chondrites may be of presolar or of partial Solar System origin (Dai et al. 2002), and also are 14N-rich (Zinner 1998). The observed differences in the oxygen and nitrogen bulk

Table 1. Comparisons of elemental and isotopic abundance ratios between planetary and solar compositions. Isotopic ratios (Conventional δ values, in ‰)2

Elemental abundances1 (log10 values) Meteorite – Photosphere3 H N C O

−3.75 −1.53 −0.99 −0.27

(‰)

R

Solar

Earth mantle

δsolar4

D/H N/14N 12 13 C/ C 17 O/16O 18 O/16O

−870 ± 205 ≤−240 ± 256 ≤−105 ± 206 ≤−40 ± 86 (−39)6,7

−80 −5 −6 +3 +6

+6100 +3105 +1105 +455 +475

15

Compiled by Asplund et al. (2005); 2δ = (Rsample/Rstandard − 1) × 1000, where Rstandard is the ratio for the terrestrial reference standard; 3Si-normalized abundance ratios are compared; 4δsolar = (Rbulk planetary/Rprotosolar − 1) × 1000, where bulk planetary composition is represented by the Earth’s upper mantle values; 5Jovian atmospheric value; 6Lunar sample-based solar wind values; 7Solar wind δ17O and δ18O values are obtained by combining the observed Δ17O (=δ17O − 0.52 δ18O) value of ≤−20±4‰, and the prediction of Young and Russell (1998) that the solar component should exhibit a relationship of δ17O = δ18O − 1.

1

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isotopic compositions between planetary materials and solar composition cannot be explained by contributions of these presolar components to the normal components produced in the Solar System. Hereafter, we attempt instead to interpret the differences in the isotopic compositions of hydrogen, carbon, nitrogen and oxygen between the solar and bulk planetary compositions by isotopic fractionation processes that operated on a common source reservoir during the formation of the Solar System.

δsolar (‰)

The δsolar values of the four ma10000 jor light elements are systematically D/H positive, and appear to be in the order of their volatilities. In Figure 3, 1000 the δsolar values are plotted against the Si-normalized abundance ratio 15N/14N of primitive chondrites normalized by the solar photospheric abundance 13C/12C 100 ratio. We explain the trend in a very qualitative sense that elements with 17O/16O & 18O/16O higher volatility, that is, those that require larger energy to condense, 10 show larger fractionation degrees in -4 -3 -2 -1 0 their isotopic ratios. This reasoning Si-normalized Abundances is partly supported, for the case of Log10(Meteorites / Photosphere) hydrogen and nitrogen, by Aléon and Robert (2004). They explain the difFigure 3. The δsolar value (δsolar = (Rplanetary / Rsolar − 1) × 1000), ference in the isotopic fractionation where R is the respective isotopic ratio which is noted at the side of each point) plotted against the elemental abundance degree between hydrogen and nitroratios in primitive meteorites normalized by the abundance gen by assuming formation of organratios in the solar photosphere, which are taken from Asplund ic molecules now present in primiet al. (2005). The δsolar values appear to increase from right to tive meteorites, IDPs and comets at a left, i.e., with increasing volatility of respective elements. temperature range between 50 K and 80 K. The difference in degree of isotopic fractionation results from differences in exothermicities between hydrogen and nitrogen for the ion-molecule reactions that fractionate their isotopic compositions. This theory could be extended to carbon and oxygen. There is, however, another factor that affects the δsolar values for the elements like carbon and oxygen, which are not as highly volatile as hydrogen and nitrogen. The trend observed between carbon and oxygen can be explained by a simple equation derived from a mass balance between the gas and the solid phase in the protosolar nebula. The equation is derived as follows. First, the mass balance equation of an isotope system is given by: Rsolar = Rnebular gas (1 − ƒsolid) + Rsolid ƒsolid

(1)

where R represents the mean isotopic ratio of respective reservoirs, and ƒsolid is the fraction of the element in the solid phase. The parameter ƒsolid can be approximated by the abundance ratio in primitive chondrites normalized by the solar abundance ratio. The mean isotopic fractionation degree Δ between the nebular gas and the solid can be defined as follows: Rsolid = (1 + Δ) Rnebular gas

(2)

Combining the two formulae, we obtain the relationship between ƒsolid and δsolar (x- and y-axes in Fig. 3), as follows: Rsolid/Rsolar ≡ δsolar /1000 + 1 = (1 + Δ)/(1 + ƒsolid Δ)

(3)

The curve expressed by Equation (3) connects two end-points, one at (ƒsolid = 1, δsolar = 0), representing the case of refractory elements, and the other at (ƒsolid = 0, δsolar = 1000 Δ), the

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δsolar (‰)

case of highly volatile elements. An 350 example of the curve for the case 15N/14N of Δ = 0.12 is plotted in Figure 4. 300 13 Interestingly, points for δ Csolar, 250 δ17Osolar and δ18Osolar plot close to this single curve. This suggests that 200 the mean degree of the gas-to-solid fractionation degree (Δ) among the 150 13 12 13C/12C C/ C, 17O/16O and 18O/16O isotopic 100 systems were similarly about 120‰. 17O/16O & 18O/16O In principle, the Δ values should be 50 different for different elements and isotopes. The coincidence of the 0 0 0.2 0.4 0.6 0.8 1 Δ values among the three isotopic systems may not be chance. The Si-normalized Abundances parallel degree of enrichments of the (Meteorites / Photosphere) minor isotopes (e.g., 13C, 17O and 18 Figure 4. The δsolar value plotted against the elemental O) in the solid phase is predicted abundance ratios in primitive meteorites normalized by to occur by the self-shielding effect the abundance ratios in the solar photosphere. Results upon photodissociation of CO of Equation (3) in the main text (where ƒsolid and δsolar molecules by UV light (van Dishoeck correspond to the x and y axes, respectively), for the case of and Black 1988; Warin et al. 1996), Δ = 0.12, are plotted as a line. which has been proposed as a mechanism that played an important role in production of the planetary solid materials in the inner Solar System (Thiemens and Heidenreich 1983; Thiemens 1999; Clayton 2002; Yurimoto and Kuramoto 2004; Lyons and Young 2005). For better understanding of the processes and places where the oxygen isotope fractionation among the planetary solids took place, further studies on isotopic fractionation processes of carbon, nitrogen or hydrogen might provide clues and means of tests, since these elements seem to partially share gas-to-solid formation pathways. The relatively narrow distribution of carbon isotopic composition among primitive meteorites (e.g., Alexander et al. 1998; Messenger 2000), compared to other organics-forming elements discussed here, such as hydrogen or nitrogen, is often explained by the existence of several chemical pathways with opposite and partially canceling carbon isotope fractionation effects, such as CO photodissociation and the ion molecule reactions involving CO that might have prevailed in the interstellar molecular cloud (e.g., Tielens 1997). The picture shown in Figures 3 and 4, that the planetary carbon is systematically enriched in 13C, might suggest that the formation of carbon-bearing planetary (organic) solids effectively took place in a location where the photodissociation of CO, among others, was the most effective triggering process to produce active carbon that led to the formation of carbon-bearing solids, such as in a relatively warm (>>10 K) gas medium around the proto-Sun (Hashizume et al. 2004), consistent with the proposed formation temperature (50-80 K) of meteoritic/cometary organics inferred by Aléon and Robert (2004). Nitrogen in the planetary solid phase is highly enriched in 15N, especially compared to the degrees of fractionation of carbon and oxygen. This might indicate that nitrogen was more effectively fractionated by processes, such as ion-molecule reactions, different from those for carbon and oxygen. However, the isotopic fractionation of nitrogen by a photochemical process similar to the cases of carbon and oxygen might also have played an important role in determining its bulk isotopic composition (Clayton 2002), although lack of precise spectroscopic data necessary to estimate the isotopic fractionation degree in the photodissociation process of N2 molecules hinders further discussion on this issue.

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Another marked difference between the nitrogen and carbon isotopes in the planetary solids is the much wider range in nitrogen isotopic composition observed among meteorites. The δ15Nsolar values among bulk meteorites range between −100‰ and +1000‰ (e.g., compiled by Hashizume et al. 2000, 2004). The high enrichment of 15N remains a puzzle. Ion-molecule reactions (Tielens 1997; Terzieva and Herbst 2000; Charnley and Rodgers 2002) can fractionate nitrogen isotopes significantly, but cannot explain the full range in bulk meteorites. Terzieva and Herbst (2000) demonstrated that the 15N/14N enrichment due to the ion-molecule reactions in the cold interstellar molecular cloud might be as high as +25%, insufficient to explain the entire range of nitrogen isotopic composition observed among meteorites. Charnley and Rodgers (2002), after Terzieva and Herbst (2000), proposed that 15N enrichment might have taken place in the cold midplane in the outer region of the accretion disk, where CO molecules in the gas, which plays a role in decreasing the nitrogen isotopic fractionation degree by the ion-molecule reactions, had totally condensed. Other possibilities are that some of the nitrogen isotopic variations are due to atmospheric loss (as in the case of the Martian atmosphere) or are nuclear in origin, either by spallation or by presolar grain carriers. The reason for the wide distribution of nitrogen isotopic composition among meteorites remains a totally unanswered question, which might be an important clue in understanding the nitrogen isotope fractionation processes in the Solar System, and could also help understanding the oxygen and carbon fractionation processes. Implications of an 16O-poor solar composition. IHNC06 raise the possibility that the oxygen isotopic composition of the solar photosphere and solar wind is not representative of bulk solar composition. If UV self-shielding raised the δ17O and δ18O of nebular water, late infall of this water into the solar photosphere could have led to an isotopically heavy photosphere. However, current understanding of solar structure and evolution suggests that the composition of the photosphere is representative of the initial composition of the Sun, apart from a limited amount of gravitational settling of all heavy elements (Lodders 2003). If the photosphere has a significantly positive Δ17O, as the IHNC06 measurements suggest, the UV self-shielding models cannot be correct in their current form, as they are only capable of increasing Δ17O relative to initial Solar System composition. If this is the case, other explanations, such as nucleosynthetic components or photochemistry of the sort advocated by Thiemens (1988, 1999, 2006) and Marcus (2004), must be sought for the distribution of oxygen isotopes in the Solar System.

Oxygen isotopic composition of the solar wind: direct measurements There is a direct measurement of the 16O/18O ratio in the solar wind, giving a value of 446 ± 90, which corresponds to δ18O = 118 ± 226‰ (Bochsler 2007). This value is not precise enough to distinguish among any of the other current values discussed above. The Genesis spacecraft was designed to collect samples of the solar wind by exposing high purity materials in space for a period of about three years. The spacecraft did this successfully and returned to Earth in September, 2004. It suffered a hard landing because of the failure of the parachute to deploy. Almost every one of the over 300 individual collectors was shattered, but enough large pieces remain that nearly all of the planned experiments can be done. During planning of the Genesis mission, a list of mission goals was prepared. The highest priority was determination of the oxygen isotopic composition of the Sun. In order to enhance collection of nitrogen and oxygen, an electrostatic concentrator, which was intended to concentrate elements in the mass range 4 to 28 amu by a factor of ~20 (Wiens et al. 2003; Heber et al. 2006), was flown on the spacecraft. Remarkably, three of the four quadrants of the small concentrator target survived the hard landing without breaking. Despite the minimal damage to the concentrator, two difficulties remain in using it for oxygen isotopic analysis: first, there is surface contamination by soil from the landing site in Utah, grains of which will have to be

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completely removed or avoided; and second, most of the Genesis collector surfaces are coated with a “brown stain,” a refractory hydrocarbon-fluorosilicone film that contains significant amounts of oxygen. Three approaches will be taken for analysis of the oxygen isotopic composition of Genesis concentrator samples (Wiens et al. 2004). (1) The MegaSIMS, an ion microprobe in which a compact tandem accelerator mass spectrometer replaces the conventional magnetic sector mass spectrometer, is nearing completion at UCLA (McKeegan et al. 2004; Mao et al. 2006). One of the difficulties in conventional ion probe mass spectrometry of Genesis samples is the interference of 16 OH− on 17O−. In naturally-occurring minerals, even hydrated ones, this interference can be dealt with by using high mass resolution. In the Genesis samples, the H/O ratio is so high that even a small tail on the 16OH− peak leads to overwhelming interference on 17O−. In the MegaSIMS, secondary O− ions are accelerated in the first half of the tandem accelerator and passed through an argon stripper gas. This strips several electrons off the O− ions and removes all hydrides, leaving positively charged oxygen ions (+1 to +3). Analyses will be done on silicon carbide from the concentrator. (2) Techniques are under development at the Open University to isotopically analyze CO liberated from a diamond target by laser ablation. Since use of diamond of normal carbon isotopic composition would lead to an overwhelming interference of 13C16O on 12C17O, one of the four quadrants of the concentrator is coated with 13C diamond. (3) Techniques are being developed at the University of California, San Diego, to laser fluorinate silicon carbide from the Genesis concentrator target. O2 so liberated from this target would be purified and mass-analyzed. Techniques have been developed over the last three years to minimize contamination from Utah dust and the “brown stain” in Genesis sample collector materials, but oxygen isotopic analysis of Genesis concentrator targets remains difficult. Once results are obtained, it will be necessary to correct for isotopic fractionation of oxygen in the concentrator itself. This will be done through electrostatic modeling (e.g., Wiens et al. 2003) constrained by precise neon isotopic measurements as a function of distance from the center of the concentrator target. The gold-coated cross that separates the four quadrants has proven to be ideal for this purpose (Heber et al. 2007). At the time Genesis was launched, the CO self-shielding models had not yet been developed and the most important question for oxygen isotopes was where the Sun fit within the fairly narrow range of ±5‰ in Δ17O. Thus, the stated goal was to measure oxygen isotopes in the Sun to a precision of better than 1‰, preferably to 0.1‰. Now, the first-order question is whether the Sun is 5% richer, 5% poorer, or about the same in 16O as the terrestrial planets.

Summary of solar oxygen isotopic composition The solar oxygen isotopic record from lunar metal grains is currently puzzling: is the Sun O-rich or 16O-poor compared to the Earth and all other planetary bodies whose compositions are known? The two studies based on lunar metal grains sampled different components (solar energetic particles and solar wind) that were implanted at different times (more than one billion years ago and relatively recently). An 16O-rich Sun is predicted by all three varieties of models that appeal to self-shielding of CO as an explanation of non-mass-dependent oxygen isotopic variations in the Solar System. An 16O-poor Sun is difficult to reconcile with any current model for oxygen isotope variations in the Solar System, but is perhaps in better agreement with the spectroscopic value of δ18O = 41+−63 61 ‰ (Scott et al. 2006). On the other hand, perhaps both lunar metal grain studies are correct, but one or both represent isotopic components other than the Sun. The next few years are bound to be interesting as these 16

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various possibilities are sorted out with high-quality measurements of the Genesis collectors and further studies of the lunar record.

ACKNOWLEDGMENTS This work was supported in part by the National Aeronautics and Space Administration and a number of other national and international funding agencies. We thank Jérôme Aléon and Edward D. Young for their constructive reviews.

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 93-140, 2008 Copyright © Mineralogical Society of America

Redox Conditions in the Solar Nebula: Observational, Experimental, and Theoretical Constraints Lawrence Grossman Department of the Geophysical Sciences and Enrico Fermi Institute The University of Chicago Chicago, Illinois 60637, U.S.A. [email protected]

John R. Beckett Division of Geological and Planetary Sciences California Institute of Technology Pasadena, California 91125, U.S.A.

Alexei V. Fedkin, Steven B. Simon Department of the Geophysical Sciences The University of Chicago Chicago, Illinois 60637, U.S.A.

Fred J. Ciesla Department of Terrestrial Magnetism Carnegie Institution of Washington Washington, D.C. 20015-1305, U.S.A. ABSTRACT Crystallization experiments on liquids with compositions similar to those of compact Type A, Type B1 and Type B2 refractory inclusions were conducted under controlled temperature and fO2 conditions. Application of the results to the compositions of coexisting Ti3+ -bearing fassaitic clinopyroxene + melilite pairs in natural inclusions shows that, if they crystallized at ~1509 K, they did so at log fO2 = −19.8 ± 0.9, only slightly below the equilibrium log fO2 of a partially condensed system of solar composition at the same temperature, −18.1+−00..23, or IW-6.8. Fassaite is the only fO2 indicator that shows that anything in chondrites formed in a system that was close to solar in composition. Solar composition is so reducing that equilibrium calculations predict vanishingly small FeO/(FeO + MgO) ratios in the condensate until temperatures fall below 800 K, where significant oxidation of metallic iron and formation of fayalite in solid solution with previously condensed forsterite begin. The mechanism for the latter process is diffusion of Fe2+ through forsterite, but the diffusion rate is nearly zero at these temperatures. By comparison to what is achievable in a system of solar composition, the mean FeO/(FeO + MgO) ratio of the olivine in chondrules in unequilibrated ordinary chondrites (UOCs) is very high, ~0.15. Making such ratios in chondrule precursors by solar nebular processes requires sufficiently high fO2 for iron to become oxidized above temperatures where diffusion of Fe2+ becomes very slow. Two dynamic models for enrichment of oxygen relative to carbon and hydrogen were investigated quantitatively: radial transport of water ice-rich migrators across the snow line into the inner part of the solar nebula where the ice evaporates; and coagulation, vertical settling and evaporation of anhydrous dust in the median plane of the inner nebula. In both cases, the maximum achievable fO2, ~IW-4.5, produces a maximum XFa before diffusion 1529-6466/08/0068-0007$05.00

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Grossman et al. ceases that is a factor of >7 less than would be required for UOC chondrule precursors, even for grains only 0.1 μm in radius and nebular cooling times as high as 106 yr. The same dynamic models are also incapable of creating environments sufficiently oxidizing to produce olivine with XFa = 0.15 during formation of chondrules by melting of FeO-poor precursors. If, instead, chondrule precursors were made of very FeO-rich, non-equilibrium condensates, reduction of chondrule melts by nebular gas may have been arrested before the mean XFa of chondrule olivine could fall below 0.15 because chondrules were hot for such a short time. A nebular origin for the mineral assemblage of unequilibrated enstatite chondrites (UECs) requires fO2 significantly below that of a system of solar composition. In particular, after fractionation of specific amounts of predicted high-temperature condensates, equilibrium condensation in a system whose Ptot = 10−4 atm and whose initial composition is solar except for a C/O ratio of 0.83 yields an assemblage characterized by a very large enstatite/forsterite ratio, the presence of oldhamite and niningerite, metallic nickel-iron containing several wt% Si, and small amounts of pure silica and albitic plagioclase, very similar to the mineral assemblage of EH3 chondrites. Log fO2 in this system varies from IW-8.9 at 1500 K to IW-13 at 900 K. The mechanisms proposed to date for fractionation of C, O and H from one another are quantitatively insufficient to produce the magnitude of nebular fO2 variations needed to account for primitive features of UOCs and UECs.

INTRODUCTION Chondrites formed in the solar nebula. Although most were modified by metamorphic processes in the parent bodies into which they accreted, a common theme of studies of the least metamorphosed chondrites is that they possess mineralogical, chemical and isotopic characteristics inherited from the solar nebula at the time of their formation. If such characteristics are interpreted correctly, they can be unique and valuable clues to the composition of and physico-chemical conditions in the solar nebular region where they formed. One such characteristic is the relative amount of a particular element in each of two or more oxidation states in a given mineral assemblage. This is a direct measure of the equilibrium partial pressure of oxygen, or oxygen fugacity, fO2, of the gas with which the assemblage may have equilibrated, in this case, the gas in the region of the solar nebula where the assemblage formed. The fO2 is largely determined by the relative abundances of carbon, oxygen and hydrogen in cosmic gases. In this work, an experimental calibration of the relationship between fO2 and the Ti3+/Ti4+ ratio of pyroxene is determined and used to infer the fO2 of the gas with which coexisting melilite and pyroxene equilibrated when they crystallized from melt droplets to form refractory inclusions. This fO2 is well below that required to produce the mean fayalite content of olivine in chondrules in unequilibrated ordinary chondrites (UOCs). Fractionations of oxygen from carbon and hydrogen resulting from two dynamic nebular processes, radial transport of water ice-bearing migrators, followed by evaporation of the ice; and coagulation, vertical settling and evaporation of anhydrous dust, are investigated quantitatively to see if they can produce the inferred fO2 variation. Finally, the fO2 of the gas is determined that can account for the distribution of silicon between metal and silicates, and of calcium and magnesium between sulfides and silicates, in EH3 chondrites by condensation.

OXYGEN FUGACITY DURING CRYSTALLIZATION OF REFRACTORY INCLUSION MELTS Experimental technique Synthesis experiments were conducted in S. E. Haggerty’s laboratory at the University of Massachusetts. Starting compositions for melting experiments were obtained by weighing and mixing CaCO3 (Baker’s ACS), MgO (Muscle Shoal’s Electrode Co. or ALFA Puratonic), Al2O3·3H2O (Fisher Certified Reagent) or Al2O3 (ALFA Puratonic), SiO2 (optical grade quality

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quartz crushed to -200 mesh or ALFA Puratonic), and TiO2 (Baker’s Analyzed Reagent). These mixtures were ground under ethanol by hand in an agate mortar for one hour, in an automatic agate mortar for 2-3 hours and allowed to dry in a drying oven or in air overnight. The oxide mixes were placed in a Pt crucible, brought to 1000 °C in stages over a period of one day and held there for 2-3 days. They were then heated in air, either in a Pt-wound resistance furnace or in a Del Tech VT-31 furnace at ~1300 °C for several hours. The samples were placed in a Pt crucible that was suspended at the end of an alumina rod into the hot spot of the furnace. Samples were quenched by removing a plug from the bottom port of the furnace, detaching the sample holder from a brass fitting at the top of the furnace and dropping the sample holder and crucible through the furnace onto a block of ice beneath the bottom port. The glasses so produced were analyzed by electron microprobe, and the three starting compositions used in this work are shown in Table 1. Spinel crystals were not observed. The bulk compositions of B1-1 and B2-1 represent attempts at producing charges having the same bulk compositions as TS33 and TS20, specific examples of Type B1 and Type B2 inclusions, respectively. Both are clinopyroxene-rich types of refractory inclusions found in CV3 chondrites, the Table 1. Bulk compositions studied former having more normative melilite due in this work (wt%). to their lower SiO2 contents than the latter. The bulk composition of ETEG was chosen B1-1 B2-1 ETEG to be undersaturated with respect to spinel MgO 8.5 12.0 6.4 and to yield large amounts of clinopyroxene Al2O3 31.8 28.3 19.2 having high Ti2O3 contents. This bulk 27.6 32.9 37.4 SiO2 composition does not correspond to that of CaO 28.1 24.7 30.8 any particular inclusion, and is much lower 4.0 2.1 6.5 TiO2 in MgO and Al2O3 and much higher in SiO2 Total 100.0 100.0 100.3 and TiO2 than all Type B inclusions (Simon and Grossman 2004). For quenching experiments, a vertical Del Tech VT-31 furnace was used, together with a Eurotherm model 90 temperature controller. Temperatures were measured with a Pt-Pt10Rh type S thermocouple calibrated against an NBS calibrated type S thermocouple and against the melting points of ITS-90 Au and Ni (1064.18 and 1455 °C, respectively). The gas mixing system is a slight modification of that designed by Williams and Mullins (1976). A mullite reaction tube was inserted into the furnace, and the gas mix was introduced through a gas inlet at the top. Chips of glass starting material were placed into cages made of 0.12 mm Ir wire, which were suspended from a long alumina rod. While the sample holder port at the top of the furnace was temporarily plugged, the furnace temperature was adjusted to within a few degrees of the desired temperature. The gas mix was then introduced at a flow rate that gave a positive pressure within the reaction tube. After minor adjustments of flow rate and temperature, the plug was removed, and the sample holder assembly was inserted through the port such that the sample was in the hot spot of the furnace adjacent to the thermocouple. Air is introduced at this time, but it was found that the fO2 at the sample site stabilized after 10-15 minutes. Samples were first placed in the experimental gas mix at a temperature 10-40 degrees above that of the first appearance of clinopyroxene for 1-2 hours, allowing the Ti3+/Ti4+ ratio of the melt to equilibrate with the gas mix. The furnace temperature was then reduced to the run condition temperature. Runs were terminated by quenching into a beaker of water by removing a plug at the base of the furnace, and passing an electric current through the Ir wire holding the sample cage to the alumina rod. A horizontal, Kanthal wound, model 1201 Marshall furnace was used for some calibration experiments. Temperature was controlled to ± 2 °C by a Marshall model 4045 temperature controller, resulting in a hot spot that varied by no more than ± 3 °C over a three inch length. Temperatures were measured as in the Del Tech furnace.

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Low fO2 was achieved by bubbling a gas with 15 vol% H2 and 85 vol% N2 through two polyethylene bottles containing a concentrated sulfuric acid solution. A plug of glass wool was inserted into the tubing beyond the point where the gas had passed through the acid baths, in order to prevent sulfuric acid droplets in the gas stream from entering the furnace. The gas mix was nonexplosive due to dilution of H2 in the N2 carrier gas. Since a slightly positive total pressure, Ptot, (~1 atm) was maintained in the gas system, the partial pressure of H2, PH2, in the N2-H2 gas mix was 0.15 atm. The partial pressure of H2O, PH2O, over the sulfuric acid was fixed by the temperature of the acid bath (24 ± 1 °C) and the concentration of H2O in the acid. Oxygen fugacities were found to be independent of flow rate between 125 and 1000 ml/min, as measured at ambient temperature. A rate of 500-1000 ml/min was used in the experiments. This generated a specific H2/H2O volume ratio in the gas and buffered the oxygen fugacity of the gas mix used in the experiments. All experiments were conducted using aliquots from the same master batch of sulfuric acid solution. No significant change in either the H2O content of the acid bath or in measured H2/H2O ratios in the resulting gas was observed. The acid bath was changed every two weeks when continuous experimentation was conducted. The densities of the sulfuric acid solutions were measured by weighing known volumes of acid before and after each batch of acid was changed. The density obtained, 1.805 ± 0.008 g/cm3, corresponds to 10-12 wt% H2O (Perry and Chilton 1973). Taking the range of PH2O in equilibrium with sulfuric acid solutions with this range of water contents at 24 °C (Perry and Chilton 1973) and assuming PH2 = 0.15 atm gives PH2/PH2O of 11100 ± 3500. The oxygen fugacity of an H2-H2O gas mix is governed by the equilibrium H2(g) + 1/2 O2(g) = H2O(g). Using the data in Chase (1998) for the free energy of this reaction, the experimental gas mix has log fO2 = −19.5 ± 0.3 at 1500 K. In another calibration study, the reaction 2Cr(s) + 3/2 O2(g) = Cr2O3(s) was reversed in the experimental gas mix in the Marshall furnace between 908 and 951 °C. Powdered Cr or Cr2O3 was placed in a mullite boat attached to a thermocouple well by Pt wires. The thermocouple well and sample were in the cooler part of the furnace, 100-200 °C, while the furnace was brought up to the desired run temperature, and the gas allowed to equilibrate; they were then pushed into the hot spot where they were held in the experimental gas stream for four days. Quenching was accomplished by pulling the thermocouple well back into the cooler part of the furnace, during which the temperature of the sample fell from the run temperature to under ~300 °C in less than three minutes. The furnace was then brought to room temperature over several hours, and the sample removed. Zero time experiments, in which samples were quenched immediately after the thermocouple indicated that the run temperature had been achieved, demonstrated that significant reaction occurred neither during introduction of the sample nor during quenching for run temperatures between 700 and 1000 °C. Holzheid and O’Neill (1995) used electrochemical measurements to determine the chemical potential of O2(g) in equilibrium with Cr(c) and Cr2O3(c) as a function of temperature. Setting RT ln fO2 equal to their chemical potential leads to the conclusion that log fO2 in the experimental gas mix was greater than −24.1 at 908 °C and less than −23.0 at 951 °C. Using these limits on the equilibrium log fO2 yields 6400 < PH2/PH2O < 9900. For such a gas mix, −19.4 < log fO2 < −19.1 at 1500 K. Phase compositions in run products were determined with an automated ARL-EMX-SM three-spectrometer electron microprobe operated at a voltage of 15 keV and a beam current of either 0.25 to 0.6 μA for wavelength dispersive analyses or 0.1μA for energy dispersive analyses. Standards used were Ilm-A128 for Fe and Ti, An 100 glass for Ca and Si, En 100 glass for Mg, and Al2O3 for Al. On-line data reduction employed the program MAGIC (JW Colby, Bell Laboratories) for wavelength dispersive analyses, and a procedure modified from that of Reed and Ware (1973) for energy dispersive analyses.

Results Results for all successful runs are summarized in Table 2, in which log fO2 is based on the PH2/PH2O ratio of the gas mix determined from Cr- Cr2O3 equilibrium.

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Average glass compositions for five of the runs are shown in Table 3. The relatively low MgO and high TiO2 contents of the ETEG bulk composition relative to those of Type B inclusions are reflected in the differences in glass compositions between the ETEG runs and those of the other bulk compositions studied. Average melilite compositions and their standard deviations are shown for each of the run products in Table 4. Melilite is a solid solution between gehlenite (Ge), Ca2Al2SiO7, and åkermanite (Åk), Ca2MgSi2O7. As seen in Table 4, the average mole fraction åkermanite, XÅk, ranges from 0.36 to 0.67 in the run products, with melilite from the bulk composition with the lowest MgO/Al2O3 ratio, B1-1, having the lowest XÅk. In the B2-1 runs, the average XÅk of melilite increases with decreasing equilibration temperature, but this relationship is not found in the ETEG runs. Synthetic clinopyroxene has a pale green to blue-green pleochroism which is more intense in the pyroxene from ETEG and B1-1 runs than in that from B2-1 runs, which is lower in Ti content. The fact that clinopyroxene synthesized at low fO2, and containing negligible amounts of multivalent elements except for Ti, exhibits pleochroism suggests the presence of both Ti3+ and Ti4+. This is further indicated by the similarity of the pleochroism and compositions of the

Table 2. Run conditions and results. Run

Bulk Composition

T (°C)

log fO2 (atm)

Time (hr)

Mineral Assemblage

83L-37 83L-33 83L-34 83L-17 83L-18 83L-19

B1-1 B2-1 B2-1 ETEG ETEG ETEG

1229 1244 1236 1229 1221 1209

−19.2 ± 0.2 −19.1 ± 0.2 −19.2 ± 0.2 −19.2 ± 0.2 −19.3 ± 0.2 −19.5 ± 0.2

63 17 13 22 26 24

Sp, Mel, Cpx, An, gl Sp, Mel, Cpx, An, gl Sp, Mel, Cpx, An, gl Cpx, Mel, An, Pv, gl Cpx, Mel, An, Pv, gl Cpx, Mel, An, Pv, gl

Sp: spinel; Mel: melilite; Cpx: fassaite; An: anorthite; gl: glass; Pv: perovskite.

Table 3. Compositions of synthetic glass. Run

83L-37

83L-33

83L-17

83L-18

83L-19

MgO (wt%) Al2O3 SiO2 CaO TiO2 FeO Total

5.80 ± 0.11 21.96 ± 0.57 40.53 ± 0.79 28.60 ± 0.02 1.86 ± 0.15 0.03 ± 0.00 98.78 ± 0.49

8.50 ± 0.02 18.72 ± 0.37 41.08 ± 0.29 29.50 ± 0.24 2.13 ± 0.18 0.01 ± 0.01 99.94 ± 0.75

4.67 ± 0.01 16.66 ± 0.29 40.13 ± 0.12 32.47 ± 0.16 5.12 ± 0.11 0.03 ± 0.01 99.08 ± 0.49

5.94 ± 0.14 15.05 ± 0.09 40.38 ± 0.86 31.94 ± 0.61 6.82 ± 0.51 0.03 ± 0.02 100.16 ± 0.90

3.82 ± 0.07 14.45 ± 0.13 39.86 ± 0.34 35.29 ± 0.65 5.97 ± 0.09 0.01 ± 0.01 99.40 ± 0.50

Mg Al Si Ca Ti Fe

0.081 ± 0.001 0.242 ± 0.007 0.378 ± 0.006 0.286 ± 0.001 0.013 ± 0.001 0

0.116 ± 0.001 0.202 ± 0.002 0.377 ± 0.000 0.290 ± 0.000 0.015 ± 0.001 0.001 ± 0.001

Cation Proportions

Uncertainties represent one standard deviation.

0.066 ± 0.000 0.187 ± 0.003 0.380 ± 0.001 0.330 ± 0.001 0.037 ± 0.001 0

0.083 ± 0.002 0.167 ± 0.004 0.380 ± 0.006 0.322 ± 0.003 0.048 ± 0.004 0

0.054 ± 0.001 0.162 ± 0.002 0.380 ± 0.004 0.361 ± 0.005 0.043 ± 0.001 0

98

Grossman et al. Table 4. Compositions of synthetic melilite.

Run MgO (wt%) Al2O3 SiO2 CaO TiO2 FeO Total

83L-37

83L-33

83L-34

83L-17

83L-18

83L-19

5.10 ± 0.35 24.47 ± 1.74 30.44 ± 1.40 40.66 ± 0.40 n. d. n. d. 100.67 ± 1.10

8.31 ± 0.77 17.31 ± 1.16 33.27 ± 0.67 40.98 ± 0.23 n. d. n. d. 98.87 ± 0.66

7.76 ± 0.28 15.23 ± 0.59 35.91 ± 0.18 40.63 ± 0.52 n. d. n. d. 99.53 ± 0.69

8.30 ± 0.28 15.92 ± 0.57 34.42 ± 0.43 40.55 ± 0.31 n. d. n. d. 99.19 ± 0.55

9.69 ± 1.43 12.19 ± 3.73 35.69 ± 2.89 40.69 ± 0.31 0.27 ± 0.13 0.08 ± 0.00 98.61 ± 0.40

7.14 ± 0.33 18.01 ± 0.95 33.89 ± 0.94 41.21 ± 0.80 0.17 ± 0.02 0.01 ± 0.01 100.43 ± 0.93

Cations per 7 Oxygen Anions Mg Al Si Ca Ti Fe Total

0.342 ± 0.021 1.300 ± 0.090 1.372 ± 0.062 1.964 ± 0.021 − − 4.978 ± 0.113

0.565 ± 0.050 0.931 ± 0.066 1.518 ± 0.026 2.003 ± 0.001 − − 5.017 ± 0.087

0.526 ± 0.022 0.817 ± 0.026 1.634 ± 0.015 1.981 ± 0.016 − − 4.958 ± 0.041

0.566 ± 0.018 0.859 ± 0.032 1.578 ± 0.014 1.989 ± 0.006 − − 4.992 ± 0.040

0.663 ± 0.094 0.665 ± 0.208 1.651 ± 0.122 2.017 ± 0.030 0.010 ± 0.005 0.001 ± 0.000 5.007 ± 0.261

0.482 ± 0.022 0.961 ± 0.050 1.534 ± 0.042 1.999 ± 0.031 0.006 ± 0.001 0 4.982 ± 0.052

XÅk

0.357 ± 0.050

0.531 ± 3.1

0.600 ± 1.3

0.572 ± 0.016

0.668 ± 0.102

0.523 ± 0.030

n. d.: not detected. Uncertainties represent one standard deviation.

synthetic pyroxene to those of natural pyroxene, called fassaite, in refractory inclusions from the Allende meteorite, which has been shown to contain both Ti3+ and Ti4+ on the basis of optical spectroscopy (Dowty and Clark 1973; Burns and Huggins 1973) and X-ray absorption near-edge structure (XANES) spectroscopy (S. Simon et al. 2005). In Table 5, each column shows the average composition of the fassaite in a particular run product, with one standard deviation on the mean of the analyses quoted as the uncertainties. TiO2tot is the wt% TiO2 when all Ti is calculated as TiO2. The entries for TiO2 and Ti2O3 are the weight percentages of these oxides when the pyroxene formula is calculated on the basis of four total cations, including exactly 1.00 Ca ion, per six oxygen atoms, the result of which is given in the middle section of the table. From such formulae, the pyroxene composition is resolved into the mole fractions of five pyroxene end-members by assigning all Mg to diopside, Di, CaMgSi2O6; all Fe to hedenbergite, Hd, CaFeSi2O6; all Ti4+ to Ti4+ -bearing pyroxene, T4P, CaTiAl2O6; all Ti3+ to Ti3+ -bearing pyroxene, T3P, CaTiAlSiO6; and the remainder to Ca-Tschermak’s molecule, CaTs, CaAl2SiO6. The mean composition of synthetic fassaite crystals from each of the experimental runs in Table 1 is plotted in Figure 1, where they are compared to analyses of multiple individual fassaite points in the natural inclusion analogs of B1-1 and B2-1, TS 33 and TS 20, respectively. All electron microprobe data for fassaite and melilite from natural inclusions used in this work are taken from Simon and Grossman (2006), or are unpublished analyses performed by the same technique and on the same instrument at the University of Chicago as in that work. In general, data for the synthetic fassaite, including that from the ETEG bulk composition, fall along the composition trends of fassaite in the natural refractory inclusions. The fassaite compositions in B2-1 lie entirely within the range of fassaite compositions in its natural analog. The B1-1 fassaite, however, is lower in MgO than the lowest MgO data point from TS33 and higher in TiO2 + Ti2O3 than the highest point for TS33, reflecting the mismatch in bulk composition between this synthetic material and its natural analog. TS33 actually contains 10.3 wt% MgO, 33.6% Al2O3, 27.4% SiO2, 26.6% CaO and 2.06% TiO2 (Simon and Grossman 2004), ~25%

Redox Conditions in the Solar Nebula

99

Table 5. Compositions of synthetic fassaite crystals. Run

83L-37

MgO (wt%) 6.7 ± 0.32 Al2O3 22.90 ± 0.45 SiO2 33.72 ± 1.32 CaO 24.72 ± 0.12 12.89 ± 1.65 TiO2tot 7.86 ± 0.97 Ti2O3 TiO2 4.09 ± 0.94 FeO 0.01 ± 0.01 Total 100.01 ± 0.97

83L-33 10.68 ± 1.03 16.92 ± 2.15 40.64 ± 1.14 25.08 ± 0.31 6.56 ± 0.55 2.98 ± 0.82 3.23 ± 0.44 0.01 ± 0.00 99.53 ± 0.83

83L-34 12.24 ± 0.14 14.54 ± 0.83 44.08 ± 1.55 25.70 ± 0.27 4.36 ± 0.52 1.83 ± 0.79 2.32 ± 1.39 0 100.71 ± 2.18

83L-17 8.94 ± 0.71 18.84 ± 1.57 36.69 ± 1.66 24.72 ± 0.36 10.79 ± 0.94 5.57 ± 0.60 4.55 ± 0.61 0.02 ± 0.02 99.33 ± 0.57

83L-18

83L-19

10.76 ± 1.43 17.44 ± 3.06 40.11 ± 3.07 25.52 ± 0.80 7.49 ± 1.67 2.66 ± 1.31 4.52 ± 0.07 0.03 ± 0.03 101.04 ± 0.89

9.19 ± 0.62 19.27 ± 1.65 35.99 ± 1.44 25.28 ± 0.29 9.73 ± 0.95 3.66 ± 0.52 5.76 ± 0.64 0.01 ± 0.02 99.15 ± 0.66

Cations in Formula with 4 Total Cations, including 1 Ca Cation, per 6 Oxygen Anions Mg Al Si Ti3+ Ti4+ Fe

0.374 ± 0.019 1.007 ± 0.029 1.258 ± 0.056 0.246 ± 0.031 0.116 ± 0.027 0

0.586 ± 0.059 0.734 ± 0.096 1.497 ± 0.061 0.092 ± 0.025 0.090 ± 0.013 0

Di CaTs T3P T4P Hd

0.374 ± 0.019 0.265 ± 0.008 0.246 ± 0.031 0.115 ± 0.027 0

0.586 ± 0.059 0.232 ± 0.030 0.092 ± 0.025 0.090 ± 0.013 0

0.662 ± 0.018 0.622 ± 0.039 1.598 ± 0.069 0.056 ± 0.024 0.063 ± 0.038 0

0.498 ± 0.041 0.829 ± 0.073 1.370 ± 0.074 0.175 ± 0.019 0.128 ± 0.018 0.001 ± 0.001

0.584 ± 0.083 0.749 ± 0.137 1.461 ± 0.136 0.081 ± 0.040 0.124 ± 0.007 0.001 ± 0.001

0.516 ± 0.037 0.855 ± 0.077 1.354 ± 0.065 0.114 ± 0.017 0.161 ± 0.019 0.001 ± 0.001

0.584 ± 0.083 0.210 ± 0.038 0.081 ± 0.040 0.124 ± 0.007 0.001 ± 0.001

0.515 ± 0.037 0.209 ± 0.019 0.114 ± 0.017 0.162 ± 0.019 0.001 ± 0.001

Mole Fractions of Pyroxene Components 0.662 ± 0.018 0.219 ± 0.014 0.055 ± 0.024 0.064 ± 0.038 0

0.498 ± 0.041 0.199 ± 0.017 0.174 ± 0.019 0.128 ± 0.018 0.001 ± 0.001

Di: CaMgSi2O6; CaTs: CaAlAlSiO6; T3P: CaTi3+ AlSiO6; T4P: CaTi4+ Al2O6; Hd: CaFeSi2O6. Oxide totals include values for TiO2 and Ti2O3, not TiO2tot. Uncertainties are one standard deviation.

higher in MgO and a factor of two lower in TiO2 than B1-1. Note that all of the synthetic fassaite have Ti3+/Titot ratios between 0.4 and 0.7, the same range as seen for the vast majority of the natural fassaite grains. Grains of synthetic spinel, anorthite and perovskite large enough to be analyzed without significant contamination from surrounding phases were found to be stoichiometric MgAl2O4, CaAl2Si2O8 and CaTiO3, respectively. Spinel contains > 18O ~ 17O. The first calculations by Lyons and Young (2005a) suggested that a slope of 1.0 is to be expected, but mass-dependent fractionation was not included in those calculations. The subsequent calculations (e.g., Fig. 13) suggest that a slope of 1.0 is produced even with mass fractionation. In addition, one would wish to include other competing MIF mechanisms, such as that proposed by Marcus (2004) and described in the previous section of this chapter. An important criterion for judging success or failure for matching models to the meteorite record will be the degree to which derived timescales can be squared with the meteorite data. For example, transport times of the 16O-poor H2O produced in models (e.g., Fig. 18) can be compared with timescales determined from decades of cosmochemical research characterizing 16 18 O O−1 and 16O17O−1 exchange reactions evidenced in objects like CAIs. If, for example, the timescales of 16O18O−1 and 16O17O−1 exchange in CAIs as determined using short-lived nuclides (primarily radiogenic 26Mg produced by decay of 26Al) are too short or too long in comparison to the model timescales for shifting Δ17O, then we will have to reconsider either the models or the meaning of the CAI data. A collateral consequence of this, and indeed any, CO self-shielding model may be that N2 should also exhibit the isotopic consequences of FUV illumination, as suggested previously (Kitamura and Shimizu 1983; Clayton 2002a). N2 is an important N-bearing species in the insterstellar medium and, apparently, in the distal regions of disks (Knauth et al. 2004). Isotopespecific photodissociation of N2 may lead to N isotope effects in more readily observable species like N2H+ and HCN in disks. Future observations of the isotope ratios for these N-bearing species may serve as tests of the importance of photodissociation in the isotope systematics of circumstellar disks in general and in the early Solar System in particular. This is important because, in the absence of photochemistry, large 15N/14N ratios in meteoritic materials would seem to require inheritance from the interstellar medium, with profound implications for the origins of organic materials in the Solar System (Busemann et al. 2006).

SUMMARY A simple galactic evolution model can account for the observed variations of oxygen isotopes in primitive meteorites. In such a scenario, silicate dust in the ISM has a life-time of about 0.5 to 1.5 b.y., values that are consistent with the estimates of Clayton et al. (1989) for coarse

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dust. The Galactic evolution model is consistent with the solar oxygen isotope composition of Ireland et al. (2006) but inconsistent with that of Hashizume and Chaussidon (2005). There are two fundamentally different chemical classes of mechanisms that are candidate progenitors of mass-independent oxygen isotope fractionation in the early Solar System. One is symmetry-induced intramolecular vibrational disequilibrium of vibrationally excited reactant oxygen-bearing molecules. The other is isotope-selective photodissociation of CO coupled with self-shielding and formation of H2O. Symmetry-induced fractionation should only have resulted in preservation of oxygen MIF effects if mediated by dust grain surfaces. Experiments, some of which are suggested here, are needed to establish the viability of surface-mediated MIF processes relevant to the early Solar System. The principle attraction of this hypothesis is that it is intrinsic to the rock-forming process itself and would therefore be expected to pervade all primitive rock materials. CO self-shielding is an attractive hypothesis for the origin of mass-independent oxygen isotope fractionation in the early Solar System because it appeals to a process known to occur in the interstellar medium and, possibly, in disks. Three astrophysical settings for CO selfshielding are proposed as sites for generation of Δ17O variability in the early Solar System. One is the inner annulus of the protostellar disk at relatively high temperature. Another is at the surfaces of the disk high above the midplane, where light from the central star grazes the gas and dust of the disk, resulting in a zone of active CO predissociation and self-shielding. Interstellar light illuminating the disk at high incident angles causes a similar horizon of CO photodestruction. The last site for CO self-shielding relevant to Solar System formation is in the molecular cloud that gave rise to the protosun. The overall consequence of CO selfshielding is conversion of CO gas to 16O-poor H2O. Timescales for transferring this signal from the outer disk to the inner disk in the region of terrestrial planet formation would have been on the order of one million years. Experiments are required to understand in detail the influence of temperature on CO self-shielding. If CO self-shielding is a natural consequence of circumstellar disk evolution, we should be able to test the hypothesis directly by searching for C16O excesses in other protostellar disks with new, high-spatial-resolution observations of young stellar objects. A key difference between galactic evolution, chemically-induced MIF effects, and CO self-shielding is the predicted relative oxygen isotopic compositions of primeval dust and the Sun. Therefore, the oxygen isotopic composition of the Sun will be a crucial arbiter that may permit us to narrow the list of possible origins for oxygen MIF in the early Solar System.

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 219-246, 2008 Copyright © Mineralogical Society of America

Oxygen and Other Volatiles in the Giant Planets and their Satellites Michael H. Wong Astronomy Department, University of California Berkeley, California 94720-3411, U.S.A. [email protected]

Jonathan I. Lunine Lunar and Planetary Laboratory, University of Arizona 1629 E University Blvd., Tucson, Arizona 85721, U.S.A.

Sushil K. Atreya Department of Atmospheric, Oceanic, and Space Sciences, University of Michigan Ann Arbor, Michigan 48105-2143, U.S.A.

Torrence Johnson Jet Propulsion Laboratory, California Institute of Technology 4800 Oak Grove Dr., Pasadena, California 91109, U.S.A.

Paul R. Mahaffy NASA Goddard Space Flight Center Greenbelt, Maryland 20771, U.S.A.

Tobias C. Owen Institute for Astronomy, University of Hawaii 2680 Woodlawn Dr., Honolulu, Hawaii 96822, U.S.A.

Thérèse Encrenaz LESIA, Observatoire de Paris Meudon 92195, France ABSTRACT Giant planet atmospheric composition and satellite densities provide insights into protoplanetary disk conditions. Abundances of condensable species and noble gases in wellmixed atmospheres can distinguish among several giant planet formation scenarios, and satellite densities are first order measurements of ice:rock ratios. Recent work on protosolar abundances, relying on three-dimensional spectroscopic modeling of the solar photosphere, provides the framework for the interpretation of measurements. Model densities of protoplanetary disk condensates are shown as a function of carbon partitioning between CO, CH4 and organics. Comparison with observed satellite densities shows that Saturn’s icy satellites are inconsistent with solar composition, and must either have formed in a water-rich environment or have suffered a complex collisional history. The larger satellites of the giant planets are consistent with solar composition, with densities that speak of variation in the partitioning of carbon. 1529-6466/08/0068-0010$05.00

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Wong et al. Thermochemical equilibrium calculations predict water as the deepest tropospheric cloud on Jupiter, the planet with the best-constrained bulk water abundance. Yet cloud base pressure levels, remote spectroscopic water vapor measurements, and in situ mass-spectral measurements have all been unable to distinguish conclusively between subsolar and supersolar Jovian bulk water abundances, due to modeling assumptions and/or the spatially-variable water vapor distribution in Jupiter’s troposphere. Modeling of images of lightning flashes is consistent with supersolar water abundances. Galileo probe measurements are consistent with an enrichment factor of 4±2 over the protosolar values for most volatiles other than water (C, N, S, and the noble gases Ar, Kr, and Xe). With that of oxygen unknown, Jupiter’s enrichments of other volatiles could be explained in terms of enrichment by heretofore unidentified solar composition icy planetesimals, by planetesimals containing volatiles trapped in water ice clathrates, or by enriched gas in the evolved disk. All models involving delivery of elements by planetesimals require planetesimal formation at temperatures below 40 K, to trap argon and molecular nitrogen. Although atmospheric C/H ratios have been measured for all four giant planets, a conclusive test of the competing formation scenarios cannot be made until O/H is measured on all four planets (extremely difficult on Uranus and Neptune), and abundances of the other volatiles and noble gases are measured for the outer three.

INTRODUCTION Oxygen-based insights from the outer planets and their moons As the dominant solid-forming element in a gas of solar composition, oxygen is a primary tracer of early protoplanetary disk conditions. Condensation of gaseous H2O at low temperature trapped other volatiles along with it, and the overall density of condensates—including water, silicates, and solid carbon—was affected by the partitioning of carbon between CH4, CO, and carbonaceous forms. Any spatial variability of protoplanetary disk conditions therefore would also have been imprinted on the condensates. This chapter reviews the extent to which early solar conditions can be reconstructed from observations of the oxygen abundances and related volatiles in the giant planets and their moons. Unfortunately, it is highly likely that the bulk abundance of water in the atmosphere of a giant planet has not yet been measured. Jupiter’s water abundance is the best constrained of the outer planets, and a large part of this chapter is a review of attempts to determine the bulk water abundance in this planet. Water in the outer planets was delivered by accretion of solid planetesimals, and these planetesimals indeed fit the profile of “dirty snowballs:” contaminants in the accreted ice resulted in the presence of other volatile gases in the present-day atmospheres of these planets. Temperature and pressure affect both the phase of ice condensed from the disk, as well as the relative abundances of the volatile gases incorporated into this ice, so trace gas abundances illuminate when, where, what, and how the planetesimals formed before their accretion into the giant planets, even if the present-day planetary water abundances are not known. In much the same way, a diner confronted by a bowl containing just tofu cubes and seaweed might infer that it had previously contained miso soup, even if the broth itself had already been consumed. The densities of outer planet satellites can be interpreted to first order as measurements of their ice-to-rock ratios. A given definition of “protosolar composition gas” permits a certain range of condensate densities consistent with protosolar composition, as a function of the dominant carbon reservoir. Past discussions of solid material formed in the outer Solar System have focused on differences expected between material formed near giant planets, where carbon is generally expected to be in the reduced form (CH4, with oxygen as H2O), and material formed in the outer protoplanetary disk, where CO is believed to be the dominant form, with less oxygen available in the form of water due to the solar C/O ratio of about 0.5 (Prinn 1993). The

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bulk densities of outer planet satellites thus allow condensate source regions to be inferred, but it seems likely that the carbon budget of the protoplanetary disk also included a high proportion of solid carbon in the form of amorphous carbon grains, macromolecular carbon, and ices of compounds such as methanol, formaldehyde, and carbon dioxide. Solid carbon complicates the interpretation of satellite densities, both because it directly affects condensate density and because it reduces the amount of CO available to sequester oxygen in the gas phase. In addition to abundance ratios, we also refer to mixing ratios and mole fractions in this chapter. In terms of a number density, nX, of species X, an abundance ratio such as X/H is simply defined as nX/nH; an outer planet atmospheric mixing ratio is defined as nX/nH2 because H2 is the major atmospheric component in the observable parts of these atmospheres; and a mole fraction is defined as nX/nT, where nT is the total atmospheric number density.

The protosolar abundances Our ability to interpret giant planet volatile gas inventories and satellite bulk densities is limited by our knowledge of the protosolar composition. In the case of outer planet moons, the protosolar C/O ratio is critical, since the partitioning of carbon between CH4 and CO controls the amount of water and thus the density of condensed solids. For the outer planets themselves, the oxygen abundance is not directly measured. Measurements of the abundances of volatile elements originally included with water ice, such as C, N and some noble gases, can be interpreted only with the aid of accurate estimates of the protosolar abundance ratios of these elements. Most early discussions of the composition of the protoplanetary disk were based on the abundances compiled by Cameron (1981). With Cameron’s C and O abundance values, the expected uncompressed density of condensates ranges from 1300 to 1900 kg m−3, depending on the partitioning of carbon between CO and CH4. This agreed reasonably well with the range of outer planet satellite densities known at the time, particularly for Ganymede and Callisto (Schubert et al. 1986). At the same time, measurements of the methane abundances in the atmospheres of the outer planets were being reported, but oxygen and other volatile gases were still unconstrained. The increasing abundance of methane in these planets with increasing distance from the Sun has been known for several decades (Slipher 1933; Owen 1967; Encrenaz et al. 1974; Lutz et al. 1976). A major revision to the solar values was proposed by Anders and Grevesse (1989), and a review of carbon chemistry by Simonelli et al. (1989) adopted similar values. The decrease in the C/O ratio relative to the previous standard resulted in a smaller range of condensate densities (1400 to 1600 kg m−3). The Anders and Grevesse (1989) standard was used to interpret the deep nitrogen and sulfur abundances discovered by the Galileo probe upon its descent into Jupiter’s atmosphere (e.g., Folkner et al. 1998; Niemann et al. 1998). Recent advances in 3-dimensional spectroscopic modeling of the solar photosphere have resulted in updated abundances of C, O, N, and other elements (Grevesse et al. 2005; Asplund et al. 2004; Allende Prieto et al. 2002). The revised oxygen abundance is about 50% lower than the previously accepted Anders and Grevesse (1989) value and, along with a change in carbon abundance, results in both greater overall condensate densities as well as a greater range (1500 to 2200 kg m−3) as a function of CO/CH4 ratio. Protosolar abundances are slightly greater than these photospheric values, due to the effects of gravitational settling (Turcotte et al. 1998; Turcotte and Wimmer-Schweingruber 2002; Lodders 2003). Throughout this review, we interpret enrichments using both Anders and Grevesse (1989) and Grevesse et al. (2005) solar compositions, to illustrate the very large effect of uncertainties in the solar abundance determinations. The new abundances, although based on improved solar photospheric modeling, raise issues between solar structure modeling and helioseismic results. A recent assessment of this “solar model problem” was made by Drake and Testa (2005) in light of new measurements of the Ne/O ratio in the coronae of active solar-like stars with the Chandra X-ray Observatory.

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They conclude that, if the higher neon/oxygen ratio of 0.4 in these stars is adopted for the Sun, the solar models can match helioseismic sound speeds using the new O and C abundances. However, Grevesse et al. (2005) prefer a solar Ne/O ratio of 0.15, based on spectroscopic observations of the Sun’s corona and on direct sampling of solar wind material. They note that the lower Ne/O value is in better agreement with the Ne/O in the local galactic medium and with theories of type II supernova nucleosynthesis. In any case, the sound speed discrepancies between helioseismic and solar model results, which are at the 3% level, are mostly confined to the interface between the radiative and convective zones of the Sun, a region that is difficult to accurately characterize with a one-dimensional model like the standard solar model (Christensen-Dalsgaard 2006). Thus, we adopt the Grevesse et al. (2005) protosolar O/H and C/H values of 5.13×10−4 and 2.75×10−4, respectively, in this work.

MEASURING OXYGEN IN JUPITER’S ATMOSPHERE Oxygen in the observable part of Jupiter’s atmosphere is present almost exclusively in the form of H2O. Two sources of atmospheric water have been identified: an internal reservoir; and an external source, first identified by the Infrared Space Observatory (ISO) satellite (Feuchtgruber et al. 1997; Coustenis et al. 1998). An external source of water has also been found in the stratospheres of the other giant planets and Titan. The external source of oxygen might have two origins: a local contribution from satellites and/or rings; or an interplanetary origin, from comets, including comet Shoemaker-Levy 9 in the case of Jupiter (Lellouch et al. 2002), and/or micrometeoritic infall. In order to constrain formation models of the giant planets, we need to determine the oxygen content of the internal reservoir. Attempts to measure the bulk abundance of water in Jupiter have been unsuccessful to date. Water condenses deep below Jupiter’s upper cloud decks, and although breaks in the cloud layers allow glimpses of deeper levels, local meteorology also reduces the water column abundance in these areas, so that the deep well-mixed water mixing ratio is inaccessible to both remote sensing and to the Galileo Probe. In this section, we review Jupiter’s cloud structure, in order to understand its effect on measurements of water. We look at the results of the Galileo Probe Mass Spectrometer (GPMS), which provided the only in situ measurement of water on Jupiter, and then review methods of inferring water abundances using remotely sensed data sets.

Structure of the cloud layers Unlike the Earth’s troposphere, in which only water clouds condense, Jupiter’s troposphere hosts clouds of at least three different compositions (e.g., Weidenschilling and Lewis 1973). The top two layers, which are most accessible to remote sensing, are composed of ammonia ice and ammonium hydrosulfide (or some other combination of ammonia and hydrogen sulfide). The water cloud is seldom glimpsed, since it lies deeper than these higher clouds. The exact pressure level at which water condenses depends on the pressure-temperature structure of the atmosphere, as well as the abundance of water vapor in the atmosphere. Figure 1 shows the maximum depth of the water cloud base (otherwise known as the lifting condensation level) as a function of water enrichment relative to solar. We first calculated the saturation vapor pressure of water for the temperature-pressure structure measured by the Galileo Probe Atmospheric Structure Instrument (Seiff et al. 1998). The saturation vapor pressures are expressed in Figure 1 in terms of solar enrichment, where the black curve corresponds to solar O/H from Anders and Grevesse (1989), and the grey curve corresponds to the Grevesse et al. (2005) protosolar O/H. Thus, for solar O/H, the deepest possible cloud would form at about 5 bar. Cloud bases at lower pressures (higher altitudes) would also be expected, since condensation and circulation can act to deplete water locally. But in an atmosphere with solar O/H, clouds could not form deeper than about 5 bar, because localized water enrichments over the well-mixed value would be difficult to create and maintain.

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Figure 1. Cloud base pressure (or lifting condensation level) as a function of bulk water abundance. Black curve is for solar O/H from Anders and Grevesse (1989), grey curve is for protosolar O/H from Grevesse et al. (2005). We used the observed temperature-pressure profile from the Galileo Probe Atmospheric Structure Instrument (Seiff et al. 1998), and determined the saturated water vapor pressure over an aqueous ammonia solution at each level (Atreya and Romani 1985) based on laboratory data for the 273-363 K range (Wilson 1925; Linke 1965).

Galileo Probe Mass Spectrometer water mixing ratio measurements One of the primary goals of the Galileo Probe Mass Spectrometer (GPMS) experiment was to descend below the expected water cloud base near 5 bar and measure the deep, well-mixed water mixing ratio (Niemann et al. 1996). However, the probe entry site was in the vicinity of a 5-μm hotspot, a region of atypical meteorology characterized by reduced mixing ratios of condensable volatiles. The deepest water mixing ratio measured by the GPMS was (4.9 ± 1.6) × 10−4 in the 17.6-20.9 bar pressure region (Wong et al. 2004). Evidence from numerous sources (see below) strongly suggests that the bulk water abundance remains unmeasured due to the probe’s entry into a 5-μm hot spot. These 5-μm hot spots are regions of “unusual clarity and dryness” (Orton et al. 1998). Early photometry of Jupiter in the 5-μm wavelength region found that the north equatorial belt was much brighter than the rest of the planet, with fluxes too high to be reflected sunlight (Westphal 1969). Higher resolution studies revealed that even within the restricted latitude range of the north equatorial belt (approximately 8-18° N planetographic latitude in Smith et al. 1979), localized hot spots were responsible for much of the radiation (e.g., Keay et al. 1973; Armstrong et al. 1976). The radiation is due to thermal emission from as deep as 5 bar in Jupiter’s atmosphere, within a spectral window of low gas opacity bracketed by NH3 absorption longward of 5.2 μm and CH4 absorption shortward of 4.5 μm (Atreya 1986). Clarity therefore modulates Jupiter’s 5-μm brightness, and a strong correlation with dryness was demonstrated by the longitude-resolved variation of the ammonia mixing ratio presented in Sault et al. (2004). Figure 2 summarizes the water vapor mixing ratio measurements of the GPMS. Each data point is derived from a single measurement of counts at the atomic mass-to-charge ratio of 18, characteristic of singly-ionized H2O. To convert these counts to an atmospheric mixing ratio, we divided the counts at mass 18 by the counts measured for a gas with constant mixing ratio, such as helium or methane. Other considerations that affect the data are the contribution to

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Figure 2. Galileo Probe Mass Spectrometer (GPMS) water measurements, from Wong et al. (2004). Individual data points have been averaged to give mixing ratios of (4.7 ± 1.5) × 10−5 over the 11-11.7 bar pressure range and (4.9 ± 1.6) × 10−4 over the 17.6-20.9 bar range. The horizontal line shows the protosolar water abundance using the Grevesse et al. (2005) O/H ratio, and the vertical line shows the lifting condensation level for a solar water abundance. The order of magnitude increase from 11 to 20 bar, much deeper than the lifting condensation level, indicates the strong effect of local meteorology on the measured water mixing ratio.

mass 18 from doubly ionized argon; pressure variation of the calibration constant relating count ratios and mixing ratios; signal nonlinearity; and instrumental background signals. Although the corrections for these effects are described more completely in Niemann et al. (1998) and Wong et al. (2004), the instrumental background signal is of particular note for this discussion. Shaded boxes in Figure 2 show water measurements for two pressure ranges: 1111.7 bar (called DL2a) and 17.6-20.9 bar (called DL2b). These pressure ranges correspond to times during the GPMS experiment when the mass spectrometer was directly sampling Jupiter’s atmosphere through one of two independent gas pressure reduction systems. Between the DL2a and DL2b experiments, the mass spectrometer was exposed to a special cell in which gases were enriched with respect to hydrogen, mainly for the purpose of measuring isotopic ratios. During this enrichment cell experiment, count rates for enriched gases (including H2O) shot up, and a decaying residual background signal was observed in the DL2b experiment just afterwards. Because of this large background signal, Niemann et al. (1998) reported only an upper limit for the 17.6-20.9 bar H2O mixing ratio, but the current value is based on detailed calibration experiments done at NASA Goddard Space Flight Center, using the flight spare version of the GPMS, in order to simulate descent conditions and model the decaying background component (Wong et al. 2004). The error bars in Figure 2 show the estimated uncertainties in each mixing ratio data point due to all sources of error, including this background signal. The decrease in the central values of the 17.6-20.9 bar points with increasing pressure implies that there is still a residual background signal (memory from the enrichment cell operation) that has not been entirely removed. However, within the estimated uncertainties of the three data points between 17.620.9 bar, we cannot distinguish between a factor of two increase over this pressure range, a factor of two decrease, or a constant value over the whole pressure range. Increases or decreases of greater than a factor of two over this range do not fall within the estimated uncertainty. The data at 17.6-20.9 bar are statistically consistent with a linear increase in water that extrapolates

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to 1.7 × solar abundance1 at 20 to 40 bar, depending on calibration constant uncertainties (Wong and Mahaffy 2001). Due to the magnitude of the estimated uncertainties, we interpret the GPMS H2O measurements as two single data points: one at about 11 bar, and one at about 19 bar, demonstrating an order of magnitude increase over the 11-20 bar range. It is likely that the water mixing ratio continues to increase even further with depth, well beyond the level from which the last signals were received from the probe at 22 bar, given the unusual meteorology of the probe entry site.

The probe entry site: A 5-μm hot spot A look at the pressure-dependent variation of condensable volatile mixing ratios in the probe entry site strongly suggests that the deep GPMS measurement does not represent Jupiter’s bulk oxygen abundance. Mixing ratios of all three condensable volatiles were observed to increase with depth, with the ammonia and H2S mixing ratios eventually leveling off at their deep well-mixed values. For ammonia, this constant mixing ratio was achieved at about 8 bars (Folkner et al. 1998). This 8-bar equilibration level is at a much greater pressure than the lifting condensation level of about 1 bar for the ~5-6 × solar bulk nitrogen enrichment relative to Grevesse et al. (2005) values measured by the probe’s radio signal attenuation (Folkner et al. 1998) and by the GPMS (Niemann et al. 1998; Mahaffy et al. 1999; Atreya et al. 2003; Wong et al. 2004). The GPMS also measured the increase of the H2S mixing ratio with depth. This gas reached its equilibration level somewhere within the 12-16 bar region (Niemann et al. 1998; Atreya et al. 1999, 2003; Wong et al. 2004), where the GPMS was not sensitive to the ambient atmosphere due to the enrichment cell experiment. The main loss process for H2S is condensation into the NH4SH cloud, which should occur at about 2.6 bar for an approximately 3 × solar H2S enrichment (Atreya et al. 1999). Thus, the staggered equilibration levels for ammonia and for H2S occur in the same order as the staggered condensation levels for the clouds formed by these gases. Water should then reach its equilibration level at the highest pressure of all, since its lifting condensation level is deepest. No quantitatively precise explanation exists for the probe entry site condensable volatile profiles, with their equilibration levels occurring much deeper than their expected condensation levels. There is, however, an obvious pattern: based on lifting condensation levels, the ammonia clouds should form at about half the pressure of the NH4SH clouds, and the pressure of the observed ammonia mixing ratio equilibration level was also about half the pressure of the H2S mixing ratio equilibration level. Given this pattern, and the fact that the water lifting condensation level is in the 5-6 bar range for water in the 1-3 × solar range (Fig. 1), it is reasonable to expect that the water mixing ratio would not reach a constant value until pressures much deeper than the H2S equilibration level, somewhere in the 12-16 bar region. A simple scaling based on the fact that the pressure at the water condensation level is about three times higher than the pressure at the NH4SH condensation level would lead to the expectation that the water equilibration level in the probe entry spot is as deep as 30-45 bar, roughly the same level at which the GPMS data would extrapolate to a solar water mixing ratio (Wong and Mahaffy 2001). There is, however, no physical basis on which to expect the same scaling law to apply to the lifting condensation levels and the probe entry site mixing ratio equilibration levels for all three condensable volatiles. In fact, a comparison of condensable volatile mixing ratios at the ammonia and H2S equilibration levels in the probe entry site shows that these mixing ratios do not scale this simply (Wong et al. 2004). Two models of 5-μm hotspots provide qualitative agreement with the condensable volatile mixing ratio profiles at the probe entry site. The 3-dimensional, nonlinear model of Showman and Dowling (2000) simulated long-lived structures with drift rates and longitudinal spacings Using Grevesse et al. (2005) solar abundances, or 1 × solar abundance using Anders and Grevesse (1989).

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that matched 5-μm hot spot observations, and it successfully reproduced the cloud-top velocity field measured for a Jovian 5-μm hot spot. But, due to computational limitations, this code could only model vertical shifts of condensable volatile profiles to about 10 bar. The Rossby wave hypothesis of Friedson (2005) agrees with the spacing between the 5-μm hotspots and the intervening spectroscopically identifiable ammonia ice clouds of Baines et al. (2002), finding pressure variations along isentropes of up to a factor of 40 within the propagating wave. Although neither of these models accurately matches all the observed characteristics of 5-μm hot spots, they are sufficiently realistic to provide convincing evidence that the water mixing ratio sampled by the Galileo probe is not characteristic of the well-mixed water abundance in Jupiter’s atmosphere.

Spectroscopic measurements of Jovian water Remote sensing measurements tell us about the water abundance in Jupiter through two main methods: direct spectroscopic analysis of water lines; and measurements of cloud depths. Unfortunately, neither method has revealed the bulk oxygen abundance of Jupiter. We briefly review remote sensing results here, since these results set valuable lower limits on the water abundance and may also inform future attempts to measure water mixing ratios in Jupiter. Modeling of water lines in the 4.5-5 μm region of Jupiter’s thermal infrared spectrum allows the retrieval of water vapor mixing ratios in the troposphere. Spectral data from the Voyager Infrared Radiometer Interferometer and Spectrometer (Voyager IRIS; Bjoraker et al. 1986a; Drossart and Encrenaz 1982; Kunde et al. 1982; Lellouch et al. 1989a,b), the Kuiper Airborne Observatory (KAO; Bjoraker et al. 1986a,b), ISO’s Short Wavelength Spectrometer (ISO/SWS; Roos-Serote et al. 1999), and Galileo Near Infrared Mapping Spectrometer (Galileo NIMS; Roos-Serote et al. 1998, 1999, 2000, 2004; Irwin et al. 1998, 2001; Nixon et al. 2001) have been interpreted to yield estimates that vary significantly with data set and with modeling approach, but there is a broad consensus that subsaturated water vapor mixing ratios are widespread. No evidence is found for water vapor mixing ratios in excess of the solar abundance at pressures from which the 5-μm emission emanates. Within Jupiter’s 5-μm window, the deepest level of penetration (where τ=1) is limited to about 6 bar by pressure-induced hydrogen absorption (e.g., Bjoraker et al. 1986b). But these measurements are not signs of a low bulk oxygen abundance, because they are influenced by correlated meteorological patterns and spectral opacity in Jupiter’s atmosphere. The overall brightness of Jupiter’s 5-μm spectral window is strongly modulated by cloud opacity near 1.5 to 2 bar (e.g., Bjoraker et al. 1986b; Roos-Serote et al. 1998; Irwin et al. 1998; Nixon et al. 2001), so spatially-averaged spectra will be weighted more towards regions with low 2-bar cloud opacity. Longitudinally resolved thermal microwave images of Jupiter demonstrate that this low cloud opacity is correlated with local depletions of ammonia (Sault et al. 2004; Wong et al. 2006a), and GPMS data show that all condensable volatile gas abundances (not just ammonia) were found to be depleted to very deep levels in at least one 5-μm hotspot. Roos-Serote et al. (2000) used a multispectral NIMS image of a 5-μm hotspot to determine that water relative humidity was very low (< 1%) across most of the region, but isolated regions nearby had very high water relative humidity, approaching saturation. These humid regions had much lower 5-μm radiances, by at least an order of magnitude, and also showed evidence of vigorous convective activity in the Galileo imaging data. Thus, it is clear that the lower spatial resolution Voyager IRIS and KAO observations are biased toward relatively cloud-free regions with low water relative humidity, and we cannot determine the amount of water vapor in the more humid regions without independently knowing the filling factor between humid and dry regions. Lunine and Hunten (1987) found that this filling factor must be 2% or less to match the 5-μm observations. Jupiter’s spectrum in the 5-μm window is also affected by the presence of a water cloud. Carlson et al. (1993) modeled IRIS 5-μm spectra in the North Equatorial Belt with a detailed

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radiative transfer code including multiply-scattering, thermally-emitting clouds with spectrallydependent extinction, and cloud base levels determined by the assumed condensable volatile abundances. They could only fit the observed spectra (which included 5-μm hot spot spectra) by including a water cloud at the pressure level corresponding to a 2 × solar water enrichment. Roos-Serote et al. (1999) compared the IRIS spectra used in Carlson et al. (1993) with spectra from Galileo NIMS and ISO/SWS. They found an anomalous slope in the IRIS spectra, which could have tricked the Carlson et al. (1993) model into requiring a water cloud. Further study of water cloud effects on the NIMS 5-μm spectral characteristics (Roos-Serote et al. 2004) revealed that this spectral region could be fit with either (a) no clouds between 2.5 and 8 bars, in which case the water vapor relative humidity never exceeds 10%, resulting in subsolar O/ H; or (b) clouds, as predicted by thermochemical models for a given composition, but with no significant opacity between 2.5 and 5 bars allowed. This finding is weakly suggestive of supersolar water abundances, since only such abundances could result in cloud condensation at pressures deeper than 5 bar. Imaging results are not compatible with a complete absence of clouds deeper than 3 bar. In Galileo imaging data, deep clouds were seen adjacent to the much higher tops of optically thick plumes associated with strong convection (Banfield et al. 1998; Gierasch et al. 2000). With cloud-top pressures in excess of 4 bar, these deep, thick clouds can only be composed of water. But since only the tops of these optically thick clouds could be seen at pressures of 4 bar or more, these clouds cannot directly imply water mixing ratios greater than those measured by the GPMS. Fifteen clouds were tracked as they moved through clearings associated with 5-μm hotspots in Cassini imaging data (Li et al. 2004). It seems very likely that these clouds were truly water clouds, because their measured speeds indicate that these clouds are at different altitudes from the opening in the 2-bar cloud layer that defines the 5-μm hotspots, and because the filter set used in the study ruled out the possibility that these clouds were instead located above the cloud deck framing the 5-μm hotspots. Although this study was unable to determine the pressure level of the deep clouds, it describes clouds at P > 4 bar. In combination with the previous result that cloud opacity is not significant between 2.5 and 5 bar (Roos-Serote et al. 2004), these clouds therefore should be located at 5 bar or deeper, again suggesting solar or supersolar water abundances. Spectroscopic measurements of CO may also be used to infer the bulk oxygen abundance in Jupiter indirectly. As mentioned before, in the observable troposphere, water is the thermochemical equilibrium species of oxygen. But at temperatures exceeding 1000 K, the CO/H2O ratio increases beyond the part per billion level. Observations of CO at the part per billion level in the troposphere (e.g., Bjoraker et al. 1986b; Noll et al. 1988, 1997; Bézard et al. 2002) therefore indicate that CO is being brought up from Jupiter’s deeper interior faster than it can reach chemical equilibrium with H2O. The observed mixing ratio of CO is then a function of the time constant for mixing, the reaction rates and pathways for chemical exchange between CO and H2O, and the total abundance of oxygen. Fegley and Lodders (1994) found that the observed tropospheric CO content was in good agreement with an O/H ratio of 2 × 10−3, which is 2.3 × solar using Anders and Grevesse (1989) values, or 3.4 × solar using Grevesse et al. (2005) values. Bézard et al. (2002) also modeled diffusion of CO from the deep interior, but used different reaction pathways and a different treatment of eddy mixing. They find 0.2-9 × solar O/H using Anders and Grevesse (1989), or 0.3-13 × solar O/H using Grevesse et al. (2005).

Lightning on Jupiter Imaging studies of Jovian lightning provide possibly the best evidence that Jupiter’s oxygen abundance is at least solar. Since reflected sunlight drowns out lightning flashes on Jupiter’s dayside, the lightning data set consists of images from spacecraft positioned to observe the planet’s nightside: the Voyagers (e.g., Cook et al. 1979; Smith et al. 1979; Borucki et al. 1982), Galileo (Little et al. 1999; Dyudina et al. 2002), and Cassini (Dyudina et al. 2004). Lightning

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on Jupiter is almost certainly generated in water clouds. Levin et al. (1983) argued that the ammonia cloud probably cannot generate sufficient charge separation, partly due to the low mass loading of the cloud. Charge separation is also inhibited by the reduced conductivity of ammonia ice, mainly due to the much lower temperature at the ammonia cloud level compared to temperatures in the deeper and warmer NH4SH and water cloud levels. Thermodynamic models predict that the NH4SH cloud layer should have a similar mass loading to the ammonia cloud, but the water cloud should be about an order of magnitude more massive (e.g., Atreya et al. 1999). Levin et al. (1983) conclude that the mass loading, simultaneous existence of liquid and solid phases, and vigorous convection combine to make the water clouds the ideal candidates for lightning generation. Lightning strikes in the water cloud are blurred in spacecraft images due to scattering by intervening cloud particles. Models relating the size (width at photometric half-maximum) of the lightning spots to the depth of the lightning strokes beneath the cloud tops in Voyager and Galileo images agree that the lightning must be located at pressures of 5 bar or more (Borucki and Williams 1986; Little et al. 1999; Dyudina et al. 2002). Since Jovian lightning strokes must be intracloud, or possibly extending upwards from the cloud (Levin et al. 1983), clouds must be present at the depths found for the lightning strokes. According to Figure 1, lightning at 5 bar or deeper therefore implies a water abundance greater than or equal to solar. The deepest flashes found by Dyudina et al. (2002) occur at 10 bar or deeper, which would imply an oxygen enrichment of at least 9 times solar. Although the degree of the enrichment depends on the accuracy of the photometric model, the agreement between multiple models that the lightning is at 5 bar or deeper is, nevertheless, a compelling result. Voyager and Galileo lightning images were taken with wideband filters, but the Cassini lightning observations were taken in a narrow filter optimized to detect Hα emission. Simulated Jovian lightning spectra (Borucki et al. 1996) show strong emission at this wavelength, with broader line emission produced at a pressure of 5 bar than at 1 bar. Dyudina et al. (2004) used the lightning frequency and brightness distributions determined from Voyager and Galileo observations, along with the expected increase in lightning signal-to-noise ratio due to the narrow Hα filter on Cassini, to predict the number of lightning strikes expected to be found in Cassini imaging data. They found 10-100 times less lightning in Cassini images than expected. Their favored explanation for this discrepancy was that lightning production at depths greater than the 5 bar used in the Borucki et al. (1996) study resulted in greater spectral broadening of the Hα emission, rendering most of the lightning strikes undetectable by the Cassini imager. Although it is possible that lightning stroke frequency was reduced by one or two orders of magnitude during the Cassini encounter compared to the Voyager and Galileo observations, it seems more likely that the unexpectedly low Hα emission is an indication of lightning production at pressures deeper than 5 bar, which in turn implies a supersolar water abundance.

Oxygen isotopes in Jupiter Oxygen isotopic ratios for Jupiter are completely unknown. The GPMS was unable to resolve oxygen isotopic signatures due to mass interference from other gases. The H217O mixing ratio could not be determined from the signal at the mass-to-charge ratio of 19, due to interference from 38Ar++ as well as a relatively high level of background signal (Fig. A.3 in Wong 2001). Although solar oxygen isotopic ratios would suggest a stronger and more easily measured signal from H218O, any signal from this molecule, at mass 20, would have been overpowered by neon.

Summary of Jovian oxygen The most important conclusion from this review of all the available data is that the oxygen abundance on Jupiter remains unknown. We have an excellent lower limit from the Galileo Probe mass spectrometer, but it is only a limit. Table 1 summarizes the oxygen abundance

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estimates determined from the numerous investigations detailed above. All of the inferred oxygen abundances are, unfortunately, heavily model-dependent, heavily influenced by meteorology and spatial inhomogeneity, or both. The best way to overcome these obstacles would be to return to Jupiter with multiple probes entering the atmosphere at different locations, and operating to depths of 50-100 bar (Atreya and Wong 2005). Being technologically challenging and expensive, such a mission is far into the future. In the meantime, help is on the way in the form of Juno! The Juno Polar Orbiter spacecraft will use passive microwave radiometry to map water on the planet to atmospheric pressure levels exceeding 100 bars. The findings of the Juno mission will be helpful in guiding the more ambitious multiprobe mission to Jupiter, where simultaneous measurements of all other key elements could also be carried out.

OUTER PLANET VOLATILE GASES Although knowledge of the oxygen abundance on Jupiter is crucial for construction of a reasonable formation scenario for Jupiter and the outer planets, information about other volatile gases on the outer planets can inform the discussion. For all the outer planets, the atmospheric mixing ratio of methane, the primary reservoir of carbon, has been measured. For Jupiter, the Galileo Probe also yielded determinations of Jupiter’s abundances of nitrogen, sulfur, and noble gases. In this section, we discuss measurements of volatiles on Jupiter and the outer planets, and finally we relate this information to conditions in the early solar system and to the formation of the giant planets.

Oxygen and other heavy element enrichments in Jupiter As discussed in the previous section, there is no evidence that the GPMS measured Jupiter’s bulk water abundance. Figure 3 summarizes the GPMS measurements of Jupiter’s volatile and noble gases. Except for water, the data are consistent with a 3 × solar enrichment within the estimated uncertainties, using the solar composition of Anders and Grevesse (1989). The enrichments are more uneven if Grevesse et al. (2005) values are used, with C/H = 4.3 ± 1.0 × solar, N/H = 4.9 ± 1.9 × solar, S/H = 2.9 ± 0.7 × solar, and the noble gases Kr and Xe are 2-2.5 × solar. The revised protosolar Ar abundance results in a Jovian argon enrichment of 5.4 ± 1.1 × solar. Despite this non-uniformity, it is safe to say that the heavy elements are enriched

Figure 3. Galileo Probe Mass Spectrometer measurements of volatile and noble gases (from Mahaffy et al. 2000 and Wong et al. 2004). Error bars show uncertainties in the GPMS mixing ratio values, and comparison between Anders and Grevesse (1989) solar abundances (squares) and Grevesse et al. (2005) protosolar abundances (circles) gives an estimate of the effect of uncertainties in solar abundance values.

>9

Depth of lightning

> 15

> 1.2

3.8, or 0.3-15

Fegley and Lodders (1994), Bézard et al. (2002) Little et al. (1999), Borucki and Williams (1986) Dyudina et al. (2002), Dyudina et al. (2004), see text

× ×

2

Banfield et al. (1998), Gierasch et al. (2000), Li et al. (2004)

Roos-Serote et al. (2004), also see text

e.g., Bjoraker et al. (1986a,b), RoosSerote et al. (1999, 2000), also see text

× ×

Niemann et al. (1998), Wong et al. (2004)

References

×

Strongly affected by H2O spatial distribution?3

×

×

Strongly affected by model assumptions?2

Solar O/H = 8.53×10−4 in Anders and Grevesse (1989) and protosolar O/H = 5.13×10−4 in Grevesse et al. (2005). See text for details about how these measurements are modeled to infer water abundances. 3 Observations are biased towards regions with lower water vapor mixing ratios, so the deep well-mixed abundance of water is not measured.

> 0.7

Depth of lightning

1

2.3, or 0.2-9

Internal CO source

subsolar, or ≥ 1.7

subsolar, or ≥ 1 > 0.3

0.03

0.02

> 0.2

0.48 ± 0.16

Grevesse et al. (2005)1

0.29 ± 0.09

Anders & Grevesse (1989)1

Clouds at P ≥ 4 bar

5-μm spectroscopy (clouds)

5-μm spectroscopy (gas)

GPMS

Measurement technique

Jovian O/H relative to solar

Table 1. Summary of measurements relating to the determination of Jupiter’s water abundance.

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at Jupiter, by a factor of 4 ± 2 × solar, even using the latest values of the protosolar elemental abundances (Owen and Encrenaz 2006). Although GPMS measurements for many of these abundances are unique, independent confirmation exists for ammonia and methane mixing ratios in Jupiter’s atmosphere. Ammonia’s mixing ratio is highly spatially variable on Jupiter, so only in the probe entry site has the deep well-mixed ammonia abundance been measured. Folkner et al. (1998) modeled the attenuation of the probe-to-orbiter radio signal as the probe descended, finding that the ammonia mixing ratio increased with depth until about the 8-bar pressure level, where it reached a mole fraction of 700 ± 100 ppm. The enrichment over solar abundance corresponding to this value is 3.6 ± 0.5 (Anders and Grevesse 1989) or 6.0 ± 0.9 (Grevesse et al. 2005). The result from the probe signal attenuation is slightly larger than the GPMS measurement quoted above (see also Fig. 3 and Wong et al. 2004). Jupiter’s methane abundance is much easier to constrain by remote sensing, since the atmosphere is too warm for methane condensation, leaving the tropospheric mixing ratio constant and spatially homogeneous. The GPMS-derived methane mixing ratio of 2.4 ± 0.6 × 10−3 (Wong et al. 2004) compares well with numerous methane mixing ratios derived from remote sensing. Gautier et al. (1982) derived a mixing ratio of 1.95 ± 0.22 × 10−3 from Voyager IRIS data, and Knacke et al. (1982) used ground-based observations in the 1100-1200 cm−1 spectral range to obtain a mixing ratio of 2.5 ± 0.4 × 10−3. Kunde et al. (2004) presented the first detection of CH4 rotational lines in spectra acquired by Cassini’s Composite Infrared Spectrometer (CIRS), confirming previous methane abundance determinations.

Volatile enrichments in the other outer planets Like Jupiter, Saturn is composed primarily of hydrogen and helium, so its composition can also be usefully discussed in terms of elemental enrichments with respect to hydrogen. Since methane does not condense on this planet, its mixing ratio should be constant within the troposphere. Numerous spectroscopic studies have yielded methane abundances (Buriez and de Bergh 1981; Courtin et al. 1984; Karkoschka and Tomasko 1992; Kerola et al. 1997). The Cassini CIRS investigation improved on these estimates by allowing a derivation of tropospheric CH4 without a priori haze and temperature profile assumptions (Flasar et al. 2005; Orton et al. 2005). The CIRS methane mixing ratio of 5.1 ± 1.0 × 10−3 corresponds to a supersolar enrichment of 7.0 ± 1.4 (Anders and Grevesse 1989) or 9.3 ± 1.8 (Grevesse et al. 2005), about a factor of two larger than for Jupiter. Modeling of Saturn’s thermal microwave spectrum (de Pater and Dickel 1991) yields more modest enrichments of ammonia of 2.2 × solar (Anders and Grevesse 1989) or 3.1 × solar (Grevesse et al. 2005), for pressures of 3 bar or more. However, 3 bar in Saturn’s atmosphere is expected to be within the condensation region for NH4SH, where ammonia mixing ratios may be variable. Thus, the microwave-derived NH3 mixing ratio should be regarded as a lower limit for Saturn’s bulk nitrogen component. Tropospheric water has been detected at 5 μm in Saturn with ISO (de Graauw et al. 1997), but, as in the case of Jupiter, the very low measured abundance (H2O/H2 = 2 × 10−7) suggests a strong undersaturation effect, probably associated with dynamical effects within dry regions of subsidence. Convection in Saturn’s atmosphere is not well understood, however, primarily because Saturn’s cooler atmosphere (compared with Jupiter) means that its clouds condense at higher pressures. Obscured by overlying haze, the clouds are more difficult to observe and therefore less useful as tracers of dynamics. Hydrogen sulfide and the noble gases have not been measured for Saturn, so the methane mixing ratio and the upper limit for ammonia provide the best constraints on the volatile enrichments in Saturn’s atmosphere. The only other heavy elements detected at Saturn—P, As, Ge and Si—are all disequilibrium species, and thus are not good indicators of the heavy element enrichment factor. Unlike Jupiter and Saturn, Uranus and Neptune have compositions that are not dominated by hydrogen and helium. Based on observed gravitational moments and other basic parameters,

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Podolak et al. (2000) conducted a Monte Carlo search of acceptable density profiles for these planets. In terms of the basic Solar System building blocks of rock, ice, and gas (or refractory, volatile, and gas components), estimates of the proportions of gas present in Uranus and Neptune depend on assumptions about the planets’ overall compositions. Thus, Podolak et al. (2000) found an absolute maximum proportion of gas (by mass) of about 30% for Uranus and Neptune, under the unrealistic assumption that the planets were composed entirely of rock and gas. Models constructed assuming a rocky core surrounded by a mantle composed of ice and gas, and constrained by available equation of state data, are reviewed by Hubbard et al. (1995) and Guillot (2006). These models generally agree that the proportion of gas (by mass) in these planets is around 5-15%, so the O/H ratio in these planets should be very large. The only heavy element abundance measured at Uranus and Neptune is that of carbon, whose value relative to hydrogen ranges from 18-50 × solar at Uranus and 28-63 × solar at Neptune, derived from Voyager radio occultations (Lindal et al. 1987; Lindal 1992) and from ground-based optical spectroscopy (Baines et al. 1995). On both planets, methane condenses at approximately the 1-bar level (Atreya and Wong 2005), which is consistent with the Voyager observations of a cloud deck at this level (Tyler et al. 1986, 1989). Oxygen in the form of CO has been detected on both Uranus and Neptune. Millimeter-wave spectroscopy of Neptune yields about 1 ppm of CO in both the troposphere and stratosphere of Neptune (Rosenqvist et al. 1992; Marten et al. 1993). Encrenaz et al. (2004) found a CO mixing ratio of 3 × 10−8 in the lower stratosphere of Uranus from CO fluorescence in the 4.65.0 μm interval, along with a CO tropospheric upper limit of 2 × 10−8. This measurement is consistent with previous millimeter-wave stratospheric CO upper limits of 3 × 10−8 (Rosenqvist et al. 1992, Marten et al. 1993). Modeling of CO diffusion from the deep interiors of these planets by Lodders and Fegley (1994) assumed tropospheric CO mixing ratios of < 10−8 on Uranus and about 1 ppm on Neptune (Rosenqvist et al. 1992, Marten et al. 1993), resulting in supersolar O/H enrichment factors for Uranus and Neptune, respectively, of < 226 and 382 (Anders and Grevesse 1989), or < 331 and 560 (Grevesse et al. 2005). Although these values are strongly dependent on the model assumptions made for mixing rates, reaction rates, and reaction pathways, they are broadly consistent with the low gas complements based on gravity data for Uranus and Neptune. However, the vertical distribution of CO in Uranus inferred by Encrenaz et al. (2004) is more consistent with an external origin of CO—possibly contributed by meteorites or icy satellites—in which case CO is not diagnostic of the bulk oxygen enrichment of the planet. For Neptune, the extreme oxygen enrichments suggested by Lodders and Fegley (1994) are challenged by the D/H ratio measured in atmospheric methane, assuming equilibration of deuterium between HDO and CH3D (Owen and Encrenaz 2006). A lower oxygen enrichment, comparable to that of methane, would satisfy the deuterium constraint.

OXYGEN IN OUTER PLANET SATELLITES To assess the effects of O and C abundances on the material that condensed in the outer Solar System, the expected condensate density as a function of carbon partitioning for both the historical and newly proposed solar abundances is calculated on a uniform basis. For these calculations, the four components are: anhydrous rock (3360 kg m−3); metallic sulfide/oxide phase (4800 kg m−3); refractory organics (1700 kg m−3); and water ice (940 kg m−3). The results are plotted in Figures 4 and 5. In both figures, the vertical axis gives the material density of condensed protosolar material, including contributions from all four components mentioned above. Figure 4 shows that the condensate density increases with the fraction of carbon in the form of CO, because CO sequesters O in the gas phase and reduces the amount of water ice in the protosolar condensate. This effect is reduced if solid organics are plentiful

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Figure 4. Model uncompressed densities of condensates in the protoplanetary disk as a function of carbon partitioning in the gas phase between CO and CH4. Models for historical values of solar photospheric oxygen and carbon abundances are shown along with the adopted protosolar values from Grevesse et al. (2005), based on recent spectral analysis and models for photospheric compositional evolution and gravitational settling. Determinations of uncompressed density for the major icy satellites of Jupiter and Saturn (Ganymede, Callisto, and Titan) are shown as well as the outer Solar System objects Pluto and Triton. The wide and puzzling range of densities for the smaller objects in the Saturn system is illustrated by values from 1000 kg m−3 (Tethys) to 1600 kg m−3 (Enceladus) with a mass averaged value of a little over 1200 kg m−3. Phoebe (probably a captured outer Solar System object) may have a density of 1600 (0% porosity) to over 2400 kg m−3 (~ 30% porosity) depending on its bulk porosity.

Figure 5. Protosolar condensate density versus fraction of carbon in the form of refractory organics. The refractory organics are assumed to have a mean density of 1700 kg m−3, equivalent to amorphous carbon. The black line illustrates the case in which the rest of the carbon is in the form of carbon monoxide, while the shaded line illustrates the case in which it is in the form of methane. For a heavy organics fraction in excess of 0.5, the difference in density between the two cases is less than 10%. Material density is equivalent to the object density for a non-porous body small enough that compression can be neglected.

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(black line in Fig. 5). But if all gas-phase carbon is in the form of CH4 (shaded line in Fig. 5), the organic fraction has a small effect on the total condensate density, because the amount of ice in the protosolar condensate is not affected.

Jupiter’s satellites Jupiter’s large Galilean satellites exhibit a strong density gradient from the innermost, Io, which is rocky and volcanic, to the outermost, Callisto. This radial density gradient is generally believed to result from the temperature gradient in the circum-Jupiter disk in which the Galilean satellites formed. Callisto’s uncompressed density of 1420 kg m−3 (based on a modestly differentiated internal structure) belies an icier composition than its next inner neighbor, Ganymede, whose density is 1570 kg m−3 (McKinnon 1997; Schubert et al. 2004). Callisto’s relatively undifferentiated interior also indicates a more gentle accretion history compared to Ganymede, so Callisto may be a better measure of the oxygen content of solid planetesimals in the circum-Jupiter disk. Callisto’s density, which lies somewhat lower than the GanymedeCallisto-Titan average shown in Figure 4, is consistent with solar-composition material formed in a CO-rich environment under the older solar composition tabulations (shaded curves in Fig. 4). Using the Grevesse et al. (2005) solar abundances, Callisto’s density now matches that of solar composition material condensed in very CO-poor conditions, implying efficient conversion of CO to CH4 in the circumplanetary disk. Inward from the Galilean satellites lies the much smaller Amalthea, with a remarkably lower density of 857 ± 99 kg m−3. It has been suggested that Amalthea, with its small size and low gravity, may have a large bulk porosity, similar to that attributed to several small asteroids. For a body of Amalthea’s size, porosities of less than 0.4 might be reasonable and stable against compaction (Belton et al. 1995), resulting in material densities of 860 to 1400 kg m−3. For material densities within this range, the moon could plausibly be constituted from material with the same uncompressed density as the outer icy satellites, requiring that Amalthea formed in a colder environment than the inner Galilean satellites, either at a different original position or at a later time when the “gas-starved” inner circumplanetary disk had cooled enough to allow ice to exist near Amalthea’s current position (Anderson et al. 2005; McKinnon 2006). No evidence of water ice is seen in Amalthea’s near infrared reflectance spectrum (Takato et al. 2004; Wong et al. 2006b), but this does not rule out an icy interior composition, because only the surface layer is spectroscopically sampled. The Trojan asteroids, like Jupiter’s satellites, may also share a dichotomy of origins. Densities are known for two Trojans, based on binary orbit determinations. At 2200 kg m−3, the density of 624 Hektor is consistent with bulk chondritic composition, suggesting formation at its current heliocentric location (Marchis et al. 2006b). However, 617 Patroclus has a much lower bulk density, 800 kg m−3 (Marchis et al. 2006a). Interestingly, the few known transNeptunian and Kuiper belt object densities also seem to demonstrate both low- and high-density populations (see below). The densities of Hektor and Patroclus therefore may support the idea that Trojans originated in the Kuiper belt, were gravitationally scattered during the passage of Jupiter and Saturn through their 1:2 mean-motion resonance, and finally became captured in Jupiter’s Lagrange points (Morbidelli et al. 2005; Tsiganis et al. 2005).

Saturn’s satellites Saturn’s satellite system consists of one planet-sized moon, Titan, a collection of small and medium-sized objects usually referred to as the icy satellites, and a retinue of distant, irregular, presumably captured objects, of which Phoebe is the largest. The range in density among these objects suggests origins in regions of differing carbon chemistry and/or significant fractionation of ice and rock from solar equilibrium values. Titan has a density that is virtually identical to that of Ganymede and Callisto, and thus is consistent with an equilibrium condensate from a reducing circumplanetary disk.

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However, the rest of what are usually referred to as Saturn’s “icy satellites” present even more serious problems for simple equilibrium condensation models than does Amalthea in the Jupiter system. The masses and sizes of the medium to small icy satellites are now well enough known from combined spacecraft and astrometric data to confirm the large variations in their densities hinted at by early Voyager measurements and the lack of a simple radial variation in density such as is seen in the Jupiter system. Enceladus and Dione have densities similar to the uncompressed Titan value, but Iapetus, Rhea, and Tethys have significantly lower values. Estimates of the densities for the small co-orbital and ring-related satellites are even lower, below that of Amalthea. The mass-weighted average density of the icy satellites (excluding Titan) is a little over 1200 kg m−3 (Jacobson 2004). It is interesting to note that the original density of Enceladus decreases from Titan-like to icy satellite-like, if its current H2O loss rate of 150 ± 30 kg s−1 (Tian et al. 2007) is assumed to have remained constant over the age of the Solar System. Although this assumption is highly questionable, it is clear that changes in satellites after their formation may complicate our attempts to use them to define conditions in the early Solar System. Given the historical solar abundance values, material with the icy satellite average density could, in principle, be consistent with a CH4-rich equilibrium chemistry, but would require fractionation or redistribution of the rock and ice to form the high- and low-density members of the group. For most of these satellites, the current abundance values are inconsistent with formation from a solar-composition source without later alteration. Saturn’s local environment was possibly enriched in water compared with the protoplanetary disk at the time that icy satellites formed, with Titan’s formation in a different environment. It is possible that the formation of Saturn somehow led to a more oxygen-rich or water-rich circum-Saturnian disk relative to the protoplanetary disk (Mosqueira and Estrada 2003). Some mechanism for redistribution of the rock-rich and ice/carbon-rich fractions is again required to produce the range of observed densities. In addition, there is no direct evidence for such large amounts of solid carbon in the satellites, whose surfaces are all extremely ice-rich. Finally, there is the possibility that Saturn’s icy satellites represent the debris from past collisions, in which the components of the impactors were distributed unevenly among the resulting fragments. Canup and Ward (2006) raised the possibility that other large satellites were present early in Saturn’s history and migrated inward as a result of interactions with gas in the circum-Saturnian disk. Such a situation would ensure collisional interactions, but collisions could also take place after dissipation of the circumplanetary disk if the resulting satellite system was not sufficiently stable (Chambers et al. 1996). Thus, Saturn’s icy satellites may have been chipped off of Titan or another differentiated satellite that was subsequently lost to collision with Saturn.

Outer Solar System satellites and Kuiper Belt Objects Data for other objects in the outer Solar System suggest a complex mixture of condensation and modification by planetary formation processes. The larger of Uranus’ satellites have densities suggesting an average composition similar to Ganymede, Callisto and Titan, consistent with equilibrium condensation in a CH4-dominated circumplanetary disk. As with the Saturn and Jupiter systems, however, there is at least one anomalously low-density moon, in this case Miranda, suggesting a more complex history. Neptune’s large moon, Triton, is believed to be a captured object, formed in the outer parts of the protoplanetary disk. Its uncompressed density is ~1900 kg m−3, making it considerably rock-rich compared with even the large satellites of Jupiter and Saturn. This is consistent with an equilibrium condensate in a CO-rich disk given the new solar C and O values (Fig. 4). A caveat is that the processes involved in Triton’s supposed capture by Neptune may have altered its original volatile composition (McKinnon et al. 1995).

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The uncompressed density of Pluto is also ~1900 kg m−3 (Stern et al. 1997), again consistent with equilibrium condensation from a CO-rich disk for the new C/H and O/H ratios. A significant enrichment of rock would be required, however, to explain this density in terms of the earlier Anders and Grevesse (1989) abundances. Eris (2003 EL61), possibly the most massive KBO known, has a density of 2200 ± 300 kg m−3 (Brown 2006). This density means that Eris, like Pluto, could have formed from solar-composition material in a CO-rich environment. The high densities of these objects argue against a large fraction of carbon occurring in the form of solid organics (Fig. 5). The irregular orbit of Saturn’s moon Phoebe is consistent with its origin as a captured object, so we discuss its composition here rather than in the Saturn section above. The Cassini flyby of Phoebe in June of 2003 yielded a mean density determination of 1630 ± 33 kg m−3 (Jacobson et al. 2004; Porco et al. 2005). This mean density would correspond to Phoebe’s material density only if the satellite had zero porosity; the shaded region in Figure 4 demonstrates that higher porosities would correspond to higher material densities. For moderate porosities, Phoebe’s material density overlaps with uncompressed densities of other objects formed in the outer parts of the protoplanetary disk, i.e. Pluto and Triton. These densities are consistent with disk chemistry from moderately reducing (CO ~ 0.25) to very CO rich values, with about 30% or less carbon in the form of organic solids (Fig. 5). As more Kuiper Belt Objects are discovered, many with satellites, more data on densities are accumulating. Preliminary indications are that there may be both high-density and lowdensity objects represented, although some of the very low densities are for bodies so small that porosity makes an accurate estimate of the actual sample density difficult. Some of the small low-density objects may be fragments of a differentiated parent body, as suggested by work demonstrating that Eris may be a member of a collisional family of small objects with strong spectroscopic water features (Brown et al. 2007).

FORMATION OF THE OUTER PLANETS Although there is considerable debate concerning many details of the formation of the outer planets, it is generally accepted that the three-phase core instability model of Pollack et al. (1996) is a good description of the formation mechanism of the giant planets. In the very rapid first phase, planetary embryos grow by solid accretion, until they have depleted the solids within their feeding zones. The planets then accrete solid and gaseous material at a slower and relatively constant rate, for periods lasting up to several million years. Finally, in the third stage, the planets reach a critical mass at which disk gas hydrodynamically collapses onto the core. The much lower gas fractions of Uranus and Neptune, compared with Jupiter and Saturn, are explained by their failure to reach the critical mass for runaway gas accretion before the dissipation of the protoplanetary disk. Thus, the study of oxygen in the giant planets is primarily the study of the solid planetesimals accreted by these planets in the first two phases of their formation. But none of the bulk water abundances of the giant planets are known. Knowledge of the enrichments of the other volatile gases is an important avenue for investigating the O/H ratios of the giant planets, since the enriched abundances of these gases suggest that they were brought to the giant planets by solid planetesimals. The difficulty with this line of inquiry lies in finding a robust and plausible mechanism for trapping the volatiles in solids while providing an equally compelling explanation of the varying degrees of volatile enrichment in the giant planets. Although the precise mechanism of volatile enrichment has not been isolated, arguments have been made in favor of accretion of planetesimals composed chiefly of amorphous or crystalline ice, or even carbonaceous material. Another possibility is that the atmospheric volatile enrichments resulted from the accretion of enriched, processed gas in an evolved protoplanetary disk.

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Volatile enrichment by icy planetesimals The commonly accepted hypothesis that Jupiter accreted planetesimals composed mainly of ice seeks to explain the abundances of other volatiles through their trapping in either amorphous ice (Atreya et al. 1999; Owen et al. 1999) or crystalline ice clathrates (Gautier et al. 2001; Hersant et al. 2004). For a given heliocentric distance in the protoplanetary disk, Gautier et al. and Hersant et al. calculated the evolutionary temperature-pressure track as a function of time. As conditions changed, they described a progressive clathration of each volatile gas remaining in the disk, based on the thermodynamic review of clathration given in Lunine and Stevenson (1985). For Jupiter’s case, they postulated that the disk cooled to about 35 K after 5.6 million years, sufficient to form clathrates of argon, before Jupiter reached critical mass for runaway gas accretion. Details of their disk evolution and clathration model could only match Jupiter’s sulfur abundance if H2S gas in the protoplanetary disk were reduced to 0.57 × solar (using Anders and Grevesse 1989 definitions). Hersant et al. (2004) attempt to justify the H2S depletion by invoking inner disk processing to remove some of the gaseous H2S, but Lodders (2004) pointed out that this mechanism would result in virtually all of the H2S being lost to FeS, so the required factor of 0.57 selected by Hersant et al. (2004) to match Jupiter’s sulfur abundance is questionable. Additionally, S/O in Comet Halley was about 0.035 in gas and dust (Mumma et al. 1993), S/O in multiple comets comes out to be about 0.028 (Irvine et al. 2000), and solar S/O is 0.03 (Grevesse et al. 2005), so evidence is lacking for subsolar sulfur in icy planetesimals. The depletion of H2S gas prior to clathrate formation was imposed for all the giant planets studied by Hersant et al. (2004). Because each clathrate guest atom/molecule is enclosed by 5.66 or 5.75 water molecules (depending on the crystalline structure of the ice), enrichment of volatiles in Jupiter by clathrate ice accretion would imply an O/H enrichment 10.5 × solar (Anders and Grevesse 1989) or 15 × solar (Grevesse et al. 2005), at the absolute minimum. The implied O/H would be even higher if clathration were less than 100% efficient. Amorphous ice is also capable of trapping volatiles. It is more likely that the planetesimals which formed the outer planets were themselves formed of amorphous ice grains, since this is the form taken by ice condensed from vapor at temperatures less than 130 K (Petrenko and Whitworth 1999). Wang et al. (2005) also note that the inelasticity of amorphous ice provides the “stickiness” necessary to facilitate the growth of planetesimals from grains. Amorphous ice formed at temperatures less than 35 K traps Ar, CO, CH4, and N2 with equal efficiency (Bar-Nun et al. 1988), while at higher temperatures the gases are less easily trapped, with different trapping efficiencies. Roughly equivalent Jovian enrichments of Ar, N, and C thus imply very low ice formation temperatures, if trapping in amorphous ice took place (Owen et al. 1999; Atreya et al. 1999; Owen and Encrenaz 2003, 2006). A low temperature of formation was also required by the scenario of Hersant et al. (2004), but the two proposed forms of ice imply differing values of the bulk water abundance in Jupiter, based on the observed volatile enrichments. Since low-temperature amorphous ice traps gas extremely efficiently, the minimum Jovian water abundance is a factor of three less than the minimum Jovian water abundance required by the clathrate mechanism. Owen and Encrenaz (2003, 2006) took a step back from the discussion of the mechanics of volatile trapping in icy planetesimals, and simply started with the assumption that these planetesimals included everything (except hydrogen, helium, and neon) in protosolar proportions. They found that the degree of volatile enrichment in the atmospheres of all of the outer planets could be explained by the accretion of approximately 10 earth masses of these solar composition icy planetesimals (SCIPs) by each planet. This mass of solar composition condensate in each giant planet would explain observed values of D/H on Uranus and Neptune, Jupiter’s heavy element enrichments, and the enrichments of carbon in all the giant planets. Some fraction of this icy mass could have built the cores of these planets—prior to directly capturing gas from the disk (e.g., Mizuno 1980, Pollack et al. 1996)—but some fraction could have been

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accreted later into the gaseous envelopes (Pollack et al. 1996). If the SCIPs contributed mainly to the giant planet cores rather than to the envelopes, then this scenario requires the further assumption that volatile species are not retained in the cores, but are mixed into the gaseous envelope—an assumption that is well-justified by the similar levels of enrichment among the noble gases and the volatiles (Fig. 3). The Owen and Encrenaz (2003, 2006) scenario was not compatible with the low methane abundance for Saturn found by Kerola et al. (1997). Instead, Owen and Encrenaz (2003) predicted a higher methane abundance that was later validated by Cassini CIRS methane retrievals (Orton et al. 2005, Flasar et al. 2005). Enrichment of heavy elements via SCIPs calls for the same enrichment for oxygen (and all other heavy elements) on Saturn as is observed for carbon. Owen and Encrenaz (2003, 2006) propose that Kuiper Belt Objects may be representative of SCIPs accreted by the giant planets, provided they have remained at temperatures below 25K since the Solar System formed.

Volatile enrichment by carbonaceous planetesimals The tarry planetesimal idea (Lodders 2004) hinges on the hypothesis that the deep water mixing ratio sampled by the GPMS was indeed characteristic of Jupiter’s well-mixed water abundance. Although the considerations presented in the previous sections argue strongly against this possibility, the Lodders (2004) model at least shows that a wide diversity of conclusions can be drawn from the relatively few outer planet compositional data available. Lodders (2004) claims that during Jupiter’s formation, temperatures in the local protoplanetary disk were too high for water condensation. The rapid accretion of Jupiter’s core was instead driven by the accumulation of carbonaceous (or tarry) planetesimals. Because she determined that sulfur would have been accreted by Jupiter entirely in the solid phase, but has been completely converted to H2S gas in Jupiter’s troposphere, Lodders (2004) normalized abundance ratios of volatile elements in Jupiter to sulfur. This normalization yields abundances of H, He, Ne, and O that are subsolar with respect to S, while Ar, Kr, Xe, and P are solar, and C and N are supersolar.2 Subsolar H, He, and Ne (with respect to S) are conveniently explained by sequestering them into Jupiter’s metallic hydrogen region, where Lodders (2004) posits other elements are not soluble. Although the lack of a good experimental or theoretical description of the molecular to metallic hydrogen transition makes the insolubility argument difficult to resolve, one weakness of the argument is the principle that heavier things sink, so any of these heavier elements should have rained onto Jupiter’s core instead of defying gravity by bubbling up into the observable atmosphere. The solar (with respect to S) abundances of the other noble gases are used as evidence that they were directly captured as gas. Supersolar C/S is the basis for the tarry planetesimal hypothesis, but supersolar N is dismissed because the N/S ratio found for Jupiter is marginally consistent, within uncertainty, with solar N/S3. One major problem with a subsolar O/S ratio is not addressed in this scenario, however. If oxygen was not brought to Jupiter in the form of ice (due to higher than previously suggested disk temperatures), then it should have been present, and therefore accreted, as gas. If the noble gases were accreted directly as gas while maintaining solar ratios (with respect to sulfur), then oxygen should have followed the same pattern. The tarry planetesimal idea of Lodders (2004) thus appears to be an insoluble paradox: it is based on the unsubstantiated premise that Jupiter’s O/S ratio is subsolar, yet a rigorous analysis of the mechanics of the hypothesis 2

Note that Lodders (2004) used a unique set of protosolar abundances (Lodders 2003), which featured C, N, O, Kr, and Xe protosolar abundances very similar to those in Grevesse et al. (2005), and S and Ar protosolar abundances slightly lower than the solar abundances of Anders and Grevesse (1989). However, the set of protosolar abundances chosen does not alter the major conclusion of her work. 3

Jupiter’s N/S = 7.5 ± 3.4 using GPMS mixing ratios from Wong et al. (2004), or 9.1 ± 2.1 substituting the ammonia mole fraction derived from probe radio signal attenuation by Folkner et al. (1998). Solar N/S = 4.4 using either Lodders (2003) or Grevesse et al. (2005) tabulations, which is marginally consistent with GPMS-derived N/S but not consistent with the ammonia abundance of Folkner et al. (1998).

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demand that Jupiter end up with a solar O/S and supersolar O/H ratios, although the minimum consistent O/H ratio would be slightly reduced by condensation of deep silicate clouds, as in Fegley and Lodders (1994). The O/H paradox, combined with the problems discussed above concerning buoyancy of insoluble elements in metallic hydrogen and the supersolar N/S ratio, renders the tarry planetesimal hypothesis untenable, although there is no reason to rule out a carbonaceous component within the icy planetesimals favored by most other researchers.

Volatile enrichment by disk evolution Evolutionary processes in the protoplanetary disk—including radial motion of solids, turbulent diffusion, and evaporation of the disk atmosphere—may have had strong effects on the giant planet volatile inventories. The “snow line” near 5 AU is a key landmark in this discussion. In a turbulent protoplanetary disk, diffusion of water vapor outside of the snow line would result in increased condensation near Jupiter’s orbit, enhancing the concentration of ice there (Morfill and Völk 1984; Stevenson and Lunine 1988). Cuzzi and Zahnle (2004) modeled a turbulent protoplanetary disk, considering the effect of an inward flux of solid material due to gas drag on meter-sized particles (Weidenschilling 1977). For water, they concluded that the inward flux of meter-sized particles was much greater than the outward diffusive vapor flux, resulting in a snow line that behaved more like an evaporation front than the condensation front examined by Morfill and Völk (1984) and Stevenson and Lunine (1988). Large enrichments of water vapor are produced inside the snow line. With the addition of a planetesimal sink (such as Jupiter) to stop the influx of solids near the snow line, water instead becomes depleted in the inner Solar System due to outward diffusion. Although enhancements of ice abundance near the snow line help satisfy the constraint that Jupiter formed rapidly enough to capture its gas before dissipation of the protosolar disk, these scenarios do little to accelerate the formation of the other outer planets. Other volatile species would be very poorly trapped in ice formed at 5 AU by outward diffusion, so abundances of nitrogen and the noble gases would be nearly solar, at odds with Jupiter’s composition. Cuzzi and Zahnle (2004) suggest that the meter-sized particles drifting inward would release other volatiles at their respective evaporation fronts, perhaps explaining Jupiter’s volatile enrichments. The variation in evaporation temperatures of the noble gases, however, would result in very different enrichment factors for each gas at Jupiter, rather than the relatively constant enrichment factor observed for the noble gases (Guillot and Hueso 2006). Instead, Guillot and Hueso (2006) proposed that photoevaporation of the protoplanetary disk would result in enriched midplane gas, which was then captured directly by the giant planets. This model is based on the protoplanetary disk evolution model of Hueso and Guillot (2005), but it includes photoevaporation driven both by extreme ultraviolet radiation from the central star in the inner (< 10 AU) disk region as well as by far ultraviolet radiation from stellar neighbors in the star-formation region. Noble gases condense on grains in the cold outer disk. Although N2 is not discussed by Guillot and Hueso (2006), due to its low condensation temperature it would behave exactly as a noble gas in their scenario. The grains settle to the midplane and drift inward due to gas drag, releasing noble gases as they move to warmer regions of the disk. Guillot and Hueso (2006) claim that the noble gases will remain at the midplane of the disk, due to the negative vertical temperature gradient, and photoevaporation will preferentially remove H, He, and Ne, species that are gaseous at all temperatures, only from the disk atmosphere furthest from the midplane. This method enriches the midplane in noble gases, which are then directly accreted by the giant planets, and makes the potentially testable prediction that the noble gas mixing ratios should be the same on both Jupiter and Saturn. Because the accretion of enriched gas implies that the giant planets formed relatively late, this evaporation scenario also nicely limits the amount of inward migration experienced by the giant planets.

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Oxygen, primarily in the form of water ice, is widely regarded as the milk that fed the young outer planets until they grew big enough to directly accrete disk gas. We have not yet determined the bulk oxygen abundance in any giant planet atmosphere, leaving giant planet formation models somewhat unconstrained, but future measurements may allow distinction between alternative formation scenarios. Numerous attempts to measure Jupiter’s bulk water abundance are summarized in Table 1. The GPMS measurement established that it is at least 50% of the solar value, in a region of atypical meteorology, with volatile depletions to great depths. Taking into account the meteorological and model limitations of the investigations summarized in Table 1, there is broad consistency with a solar or greater water abundance. Although the lightning depth analyses are very model-dependent, multiple investigations by separate teams all suggest supersolar water abundances. Confidence in this result is bolstered by the fact that supersolar water would ensure both liquid and solid cloud particles, greatly facilitating lightning generation. Other volatile gases besides water have been successfully measured in Jupiter, and the methane mixing ratios are known to increase with distance from the Sun in all four outer planet atmospheres. These measurements have stimulated discussions of the details of outer planet formation. The observed methane abundances are consistent with the ideas of Owen and Encrenaz (2003, 2006), in which about ten earth masses of solar composition icy planetesimals were accreted by all the giant planets, leading to the prediction that, in each planet, the enrichments with respect to hydrogen of all the volatiles (nitrogen, oxygen, sulfur, noble gases other than neon) should be equal to the carbon enrichments. Alternately, if water enrichments are found to be at least three times greater than the enrichments of other volatiles, then accretion of volatiles trapped in water ice clathrates would be a reasonable explanation (Gautier et al. 2001, Hersant et al. 2004). If we someday find noble gas mixing ratios to be identical on all the giant planets, than accretion of volatiles in the gas phase from a chemically evolved disk would be the obvious mechanism (Guillot and Hueso 2006). The recently updated protosolar abundance tabulations of Grevesse et al. (2005) make interpretation of Jupiter’s volatile inventory (Fig. 3) more difficult than with the older Anders and Grevesse (1989) values. The exceptionally high argon enrichment is particularly challenging to account for. As the most volatile species in Figure 3, argon’s larger enrichment compared to the other gases defies any enrichment process discussed in this chapter. One possible interpretation would be to consider the factor of two revision in protosolar argon to be indicative of a factor of two uncertainty in all the protosolar volatile abundances, leading to the conclusion that the GPMS measurements yield an enrichment of 4 ± 2 × solar for all the volatiles (except water). This conclusion should hopefully be robust against future improvements to the protosolar abundances. The newly proposed values for the solar abundances of carbon and oxygen also result in a significant increase in the expected density of condensates from a solar composition disk, regardless of the state of carbon in the system. The densities of objects formed in the outer Solar System, either in the protoplanetary disk or in circumplanetary disks, are generally consistent with the density of equilibrium condensates expected for a range of carbon chemistry in these environments. A notable exception is the Saturn system, where the presence of very low-density satellites is inconsistent with equilibrium expectations and seems to require a more complex scenario with processes resulting in depletion of silicate materials, collisional disruption of differentiated satellites, or enrichment of water ice.

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ACKNOWLEDGMENTS Portions of this work were done at the University of California with support from NASA grants NAG 5-12062 and NNG 05GF09G, and at the Jet Propulsion Laboratory, Caltech, under a contract from NASA. We thank Matthew Browning for helpful discussions, and Glenn MacPherson for organizing the fascinating and interdisciplinary “Workshop on Oxygen in Earliest Solar System Materials and Processes” just outside of the Great Smoky Mountains National Park. W.B. McKinnon provided a helpful and professional review.

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 247-272, 2008 Copyright © Mineralogical Society of America

Oxygen in Comets and Interplanetary Dust Particles Scott A. Sandford,1 Scott Messenger,2 Michael DiSanti,3 Lindsay Keller,2 and Kathrin Altwegg4 1

NASA Ames Res. Center, Astrophysics Branch, Moffett Field, CA 94035, U.S.A. NASA Johnson Space Center, Astromaterials Res. Office, Houston, TX 77058, U.S.A. 3 NASA Goddard Space Flight Center, Greenbelt, MD 20771, U.S.A. 4 University of Bern, Physikalisches Institut, Bern 3012, Switzerland

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e-mail (corresponding author): [email protected]

ABSTRACT Comets are thought to have accreted in the cold, outer portions of the protosolar nebula when our Solar System was forming, and they clearly contain a higher proportion of volatiles than materials formed closer in to the Sun. Storage of most comets in the cold, outer regions of the Solar System since their formation has probably helped minimize any subsequent parent body processing of their contents. As a result, cometary materials may represent samples that best preserve the original components from which our Solar System was made. Comparison of cometary and asteroidal materials (which formed much closer to the Sun) can also provide insights into large-scale heterogeneity and transport of materials in the early solar nebula. Comets are also of considerable interest since they may have delivered volatiles, like H2O, to the early cooling Earth, thereby playing a key role in making the Earth habitable. Comets should also have delivered organic materials to the surface of the early Earth and these, depending on their nature, may have played a role in the origin of life. Thus, the study of the composition of comets has the potential to provide important insights into the formation and evolution of our Solar System (and by extension, other planetary systems), and the creation of an inhabited Earth. The study of the chemistry, mineralogy, and isotopic distributions of oxygen in cometary materials can provide unique information that addresses these issues. Our current knowledge of the nature of oxygen in comets is based on several different lines of evidence, including remote telescopic and spacecraft observations of comets, direct laboratory analyses of extraterrestrial samples, and in situ measurements of a small number of individual comets. Information derived from these different approaches suggests that the chemical, mineralogical, and isotopic state of oxygen in these primitive bodies is extremely variable, i.e., comets appear to be made up of a wide range of very different components that are considerably out of equilibrium with each other. This is consistent with the idea that cometary materials should have largely escaped extensive parent body processing. The study of samples recently returned from Comet 81P/Wild 2 by the Stardust spacecraft has immensely improved our understanding of cometary materials. Of particular interest is the observation that these samples contain intimate mixtures of both volatile and refractory oxygenbearing materials. This suggests that while comets formed in the outer reaches of the protosolar disk, they were constructed of materials that had originally formed and evolved in a wide variety of locations that spanned essentially the entire radial extent of the protosolar nebula.

INTRODUCTION Comets have enormous scientific significance because they represent samples from the most distant regions of the Solar System and probably best preserve the original starting materials. Comets are rich in volatiles and ices because they formed in cold regions of the 1529-6466/08/0068-0011$05.00

DOI: 10.2138/rmg.2008.68.11

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solar nebula (Boss 1998), well beyond the “snow line” where water ice condensed. These objects derive from two distinct source regions: the Kuiper Belt—a flattened disk (i < 30°) of plutons and other small Solar System bodies beyond the orbit of Neptune (30-50 AU); and the Oort cloud—a spherically distributed population of bodies residing very far from the Sun (103-105 AU). Whereas Kuiper Belt objects (KBOs) are thought to have formed in place (Luu and Jewitt 2002), the Oort cloud comets are thought to have been gravitationally scattered from the neighborhood of the giant planets (Oort 1950). Owing to the large range in radial distances over which comets formed (5-50 AU), cometary materials have experienced correspondingly wide ranges of thermal histories, and may incorporate different proportions of presolar and Solar System components. Overall, comets are likely to preserve the most pristine material to be found in the Solar System, and thus represent the closest remaining link to the dense molecular cloud from which our Solar System formed. Comparison of cometary and asteroidal materials also provides insight into large-scale heterogeneity and transport of materials in the early solar nebula. The chemistry, mineralogy, and isotopic distributions of oxygen in cometary materials can provide unique insights into these issues. However, comets remain enigmatic objects and many of their basic properties are not well understood. Our current knowledge of the chemical distribution and isotopic composition of oxygen in comets is based on several different lines of evidence, including: remote observation of comets using telescopes and spacecraft; direct laboratory analyses of interplanetary dust particles thought to come from comets; inferences from meteoritic materials; theoretical considerations and inferences from interstellar materials; and in situ measurements at comets. It should be noted that, despite their uncertain provenance, studies of stratospherically collected interplanetary dust particles (IDPs) have provided some of the most detailed information so far on the distribution of O in the non-volatile fraction of cometary materials. Anhydrous IDPs have been linked to short-period comets by their unequilibrated mineralogy, fragile structure, fine grain size (50-1,000 nm), high volatile element and C content, and their high abundances of presolar materials (Bradley et al. 1988; Nier and Schlutter 1990, 1992; Thomas et al. 1993; Zolensky and Barrett 1994; Messenger and Walker 1997; Messenger 2000; Messenger et al. 2003). These IDPs have also been dynamically linked to comets from their inferred high atmospheric entry velocities that reflect eccentric orbits prior to Earth encounter (Joswiak et al. 2000). Owing to the stochastic nature of orbital evolution of interplanetary dust particles (see, for example, Liou and Zook 1996, 1997), the origins of specific IDPs are uncertain, but it is highly likely that many, if not most, of the anhydrous IDPs have cometary origins. In the discussions that follow, we will therefore consider their chemical, mineralogical, and isotopic compositions to have bearing on the nature of cometary materials. It should be noted that samples recently returned to Earth from Comet 81P/Wild 2 by the Stardust Mission are generally consistent with this basic contention (Brownlee et al. 2006). Each of the approaches described above provides unique insights into the nature of cometary materials and their oxygen carrier reservoirs. Each approach also suffers from different “selection effects” in terms of limitations as to what can be learned from them. A complete understanding of the nature of carriers of oxygen in comets requires a careful synthesis of evidence gathered from these different approaches. Since comets are known to consist of mixtures of both volatile and refractory components (cf. Festou et al. 2004) and these two components are often amenable to study using very different techniques, the discussion is divided into sections that deal with these reservoirs of material somewhat independently. Our current understanding of the chemical forms of oxygen within both the volatile and refractory components of comets is considered first, followed by a review of what is known about the oxygen isotopic ratios in these different cometary components.

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THE CHEMICAL FORM OF OXYGEN IN THE INTERSTELLAR MEDIUM, “COMETARY” INTERPLANETARY DUST PARTICLES, AND COMETS Oxygen carried by carbonaceous materials in the interstellar medium, meteorites, cosmic dust, and cometary samples The history of oxygen begins with its nucleosynthetic production in stars (D. Clayton 2003; Meyer et al. 2008). Newly synthesized oxygen is ejected into interstellar space from asymptotic giant branch (AGB) stars, novae, and supernovae in mineral grains, free atoms, and gas phase molecules. The nature of the ejecta is dependent on a number of variables including gas density, temperature, and composition (cf. Gail and Sedlmayr 1987). One of the most critical variables that defines the nature of stellar ejecta is the C/O ratio of the material in the outflow. If C/O > 1 in a local outflow region, then most of the O can be incorporated into the very stable CO molecule. This leaves the excess C free to become incorporated into SiC and other oxygen-free materials. If, however, C/O < 1 then most of the C is incorporated into CO and the excess O is free to form other materials, principally molecular oxides, silicates, and oxide minerals (Tsuji 1986; Tielens 1991). Thus, the principal O-bearing components that are injected into the diffuse interstellar medium (ISM) are CO gas and a variety of refractory oxygen-bearing minerals. Once injected into the galaxy’s diffuse interstellar medium, these materials are repeatedly cycled through a wide variety of environments (Fig. 1). These include the diffuse ISM, dense molecular clouds, and star-formation regions, from which most of the material is ultimately recycled back into the diffuse ISM, although some of the material is incorporated into newly formed stars and planetary systems. The conditions in these different environments vary over a wide range of temperatures, pressures, radiation fluxes, etc. Transitions between these environments can radically transform the state of the oxygen. The diffuse interstellar medium. The diffuse ISM is filled with a high radiation field. This radiation is sufficient to destroy most gas phase molecules. In addition, shock waves from energetic events like supernova explosions frequently traverse the diffuse ISM. These shock waves can destroy dust particles by a variety of processes (sputtering, grain-grain collisions, etc.) (cf. Seab 1987; Shull and Draine 1987; McKee 1989; Jones et al. 1994). As a result of these destructive processes, the majority of O in the gas phase portion of the diffuse ISM is in the form of atomic O (Cartledge et al. 2001). The main carrier of O among solids in the diffuse ISM is silicates; infrared absorption spectra taken on lines of sight passing through the diffuse ISM suggest that the silicates are largely non-crystalline (Kemper et al. 2004). The diffuse ISM also contains C-rich materials. The principal organic component of the gas phase consists of polycyclic aromatic hydrocarbons (PAHs) (Onaka et al. 1996), while the main organic component of the solid phase consists of C-rich grains that contain both aromatic and aliphatic structures (Sandford et al. 1991; Pendleton et al. 1994). Current infrared spectroscopic evidence, however, suggests that neither of these C-rich populations carry very much oxygen (Pendleton and Allamandola 2002). Dense interstellar molecular clouds. Shocks and the radiation field of the diffuse ISM can penetrate the edges of dense molecular clouds, so it is perhaps not surprising that oxygen found in the edges of clouds is still largely present in the form of atomic O and amorphous silicates (cf. Whittet et al. 1996; Vastel et al. 2000). However, the greater density of gas and dust found deeper in dense interstellar clouds promotes the synthesis of a wide variety of new molecular materials. In addition, the interiors of these clouds have optical depths that are sufficient to absorb most of the radiation field present in the ambient diffuse ISM. This allows newly created molecular species to survive over appreciable time scales. These large optical depths also result in very low gas and grain temperatures (10 K < T < 50 K). Under these conditions, molecular species can be formed by a variety of different chemical processes and survive.

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Figure 1. Materials in the galaxy are repeatedly cycled through a wide variety of environments, including the diffuse ISM, dense molecular clouds, and star-formation regions. Most of the material in star–formation regions is redistributed back into the general diffuse and dense ISM, but a small fraction ends up in stars and planets. Some of this material is returned to the ISM at the end of the stars’ lifecycles. Comets are thought to be relatively pristine reservoirs of the material from which the Solar System formed.

One of the principal chemical processes that occurs in the gas phase is ion-molecule reactions (cf. Herbst 1987; 2003). These reactions create a variety of O-bearing species (van Dishoeck and Blake 1998; Ehrenfreund and Charnley 2000). The resulting compounds are fairly simple and relatively volatile; these species are not particularly “meteoritic.” Further chemical processing is required to “set” the oxygen into forms that can survive in meteoritic materials. At the low temperatures that occur in dense clouds, however, most gas phase species will condense onto the first dust grain they collide with (Fig. 2). This results in additional gas-grain chemistry. In this case the H/H2 ratio plays a critical role in determining what will be produced by gas-grain reactions (cf. Tielens and Hagen 1982; d’Hendecourt et al. 1985; Brown and Charnley 1990; Hasegawa et al. 1992; Charnley 1997). In environments where H/H2 > 1, the principal chemical reactions add hydrogen to other atomic species on the ice surface, resulting in simple hydrides like H2O, CH3OH, NH3, and CH4. This results in highly polar ices dominated by H2O. If, however, the local gas has H/H2 < 1, gas-grain reactions yield ices dominated by less-polar molecules like CO, CO2, O2, and N2. Infrared spectra of the ices in dense clouds support this basic concept (Sandford et al. 1988; Tielens et al. 1991). Provided they were not greatly warmed during their incorporation into comets during formation of the Solar System, it is possible that these icy materials could be directly preserved in comets. As with the products of gas phase ion-molecule reactions, however, these icy compounds would not have efficiently survived in meteoritic materials; further chemical processing was required to “set” the oxygen in these ices into the more refractory forms that survived in meteoritic materials. One process that may play a role in generating more refractory materials is irradiation chemistry (Fig. 2). The ice mantles surrounding dust grains in dense clouds and the protosolar

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Figure 2. A large fraction of the volatile materials in dense interstellar clouds condense onto grain mantles. These ices are generally rich in oxygen in the form of molecules like H2O, CH3OH, CO, and CO2. Cosmic rays and UV photons can break down these molecules into ions and radicals that can recombine to produce considerably more complex, and more refractory, O-bearing organic species.

nebula can be further processed by cosmic rays, UV radiation from the attenuated diffuse ISM field, or UV produced by nearby stars and cosmic ray interactions (Norman and Silk 1980; Prasad and Tarafdar 1983). These forms of radiation can break bonds within molecules in the ice and generate ions and radicals that can subsequently react. This results in additional chemistry that forms a wide variety of more complex, largely C-rich, species. Laboratory studies show that some of the species produced by charged particle and UV irradiation contain oxygen and resemble those found in meteoritic and cometary materials. These include molecular species like amphiphiles, amino acids, and aromatic ketones and alcohols (cf. Dworkin et al. 2001; Bernstein et al. 1999, 2002a, 2003). Of particular interest is the irradiation of PAHs in mixed molecular ices, since this results in the addition of excess H and a variety of chemical side groups (=O, -OH, -NH2, -CN, -CH3, -OCH3, etc.) derived from other species in the ice (Bernstein et al. 2002b, 2003). As shown in Figure 3, in H2O-rich ices, i.e., the ices seen to dominate most dense clouds, the principal additions to PAHs are =O and -OH groups linked to peripheral C atoms on the PAH, and bridging oxygen spanning “bay” regions (aromatic ethers) (Bernstein et al. 1999). Many of these species resemble those found in meteorites. For additional, more detailed discussion of oxygen in the interstellar medium, see the chapter by Jensen et al. (2008). The oxygen in organics in meteorites, cosmic dust, and comet samples. Some of the oxygen in C-rich carriers in primitive meteorites resides in soluble species like amino acids, amphiphiles, carboxylic acids, etc. (cf. Cronin et al. 1988; Krishnamurthy et al. 1992; Huang et al. 2005). However, a major C-rich carrier of oxygen in meteorites is an insoluble macromolecular material (“kerogen”) of uncertain origin (Cronin et al. 1988; Cody et al. 2002). In the Murchison carbonaceous chondrite, this material has relative abundances of C, O, and N of 100:18.3:3.8, respectively (Cody et al. 2002). The oxygen in this macromolecular material is found in both aromatic and aliphatic moieties (Gardinier et al. 2000; Cody et al. 2002). The meteoritic macromolecular material appears to be much more O-rich than the Crich solid-state material seen in the diffuse ISM, but considerably less O-rich than the average of organics in collected IDPs (Flynn et al. 2006) and the organics seen in cometary grains returned from comet 81P/Wild 2 by the Stardust mission (Sandford et al. 2006).

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Figure 3. Polycyclic aromatic hydrocarbons (PAHs) are abundant in space. Their exterior carbon rings are easily functionalized with oxygen when they are irradiated in O-bearing ice mantles in dense interstellar clouds. The mix of functional groups depends somewhat on the composition of the ice, but =O, -OH, -O(ether), and –O-CH3 groups are commonly seen in laboratory simulations.

Because of their small sizes, considerably less detail is available about the nature of the organics in IDPs. They contain abundant aromatic species (see, for example, Allamandola et al. 1987; Wopenka 1988; Quirico et al. 2005), and a variety of specific PAHs have been identified in them (Clemett et al. 1993; 1998). The total oxygen content of the organics in IDPs is generally much higher than that seen in the macromolecular material in meteorites (Flynn et al. 2006; Sandford et al. 2006), but the nature of the molecular carriers is still unclear. Given the abundance of aromatic species in these particles, aromatic ketones and alcohols may be possible carriers. Amino acids have been identified in larger Antarctic micrometeorites (Matrajt et al. 2004), but these are unlikely to be a major reservoir of the O found in IDPs since they are minor components in meteorites and micrometeorites. Furthermore, these amino acids may be products of aqueous processing (Strecker synthesis), and the parent bodies of anhydrous IDPs have not experienced hydrous alteration. The organics returned from comet 81P/Wild 2 by the Stardust spacecraft are generally considerably richer in O than meteoritic organics, but are comparable, on average, to those of measured stratospheric IDPs (Sandford et al. 2006). X-ray Absorption Near Edge Spectroscopy (XANES) studies of these cometary samples demonstrate that the O is present in a wide variety of bonding states, including aldehydes, alcohols and esters. While the 81P/Wild 2 samples have O/C ratios somewhat similar to those seen in IDPs, it is clear that the two types of samples are not identical. The Wild 2 samples appear to contain a labile organic component that is missing from anhydrous IDPs (Sandford et al. 2006). If anhydrous IDPs have a cometary origin, this suggests that they may have lost a more labile fraction of their original organics during atmospheric entry or during their transit through interplanetary space from their parent body to Earth. Also, as noted above, the organics in IDPs are generally dominated by aromatic materials. In contrast, many Wild 2 samples contain little or no aromatic materials. The organics in these aromatic-poor grains contain large amounts of O, indicating that the O in Comet Wild 2 organics cannot be solely associated with aromatic species. In these particles, the carrier may be dominated by a “polymeric” material akin to polyoxymethylene or related “irregular” molecular structures similar to those made when mixed molecular ices are processed by high energy radiation (Schutte et al. 1993; Bernstein et al. 1995; Sandford et al. 2006). The generally increasing O content of organics from the diffuse ISM, to meteorites, to IDPs and cometary samples suggests that meteoritic, IDP, and cometary organics do not simply consist of unaltered organics from the diffuse ISM, but must contain contributions from either dense cloud, protosolar nebula, and/or parent body processes. The fact that material in

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IDPs and from comet 81P/Wild 2 generally shows higher O contents than meteoritic organics further suggests that dense cloud and/or protosolar nebular processes (radiation processing, gas and gas-grain chemistry, etc.) must generally act to drive up the overall O content of organics more than do parent body processes (aqueous alteration, heating, etc.) In summary, oxygen is seen in a variety of forms in stellar outflows, the diffuse ISM, and dense molecular clouds. Silicates and other oxides are the main carriers of solid-state oxygen in the diffuse ISM and are major carriers in dense clouds. Most of the gas phase O in the diffuse ISM is in the form of atomic O; very little oxygen seems to be associated with the C-rich materials seen in the diffuse ISM. Most of the non-mineral oxygen seen in the dense ISM is in relatively volatile, “non-meteoritic” forms, and much of it is in the form of ices. “Fixing” interstellar oxygen into C-rich forms that can survive incorporation into asteroidal and cometary parent bodies requires significant processing in dense cloud and protostellar nebular environments. These chemical processes result in organics that contain oxygen in a diverse set of molecules ranging from small, soluble species to macromolecular “kerogens.” In the specific case of comets, the organics (at least for comet 81P/Wild 2) are extremely O-rich and the oxygen is present in a wide variety of bonding states.

Direct detection of oxygen-bearing volatiles in comets Whipple (1951) was the first to recognize the importance of comets for the history of our Solar System. Late in the last century it was accepted that comets represent the best-preserved material from the early Solar System. Because they retain their volatiles, comets also appear to have retained their full complement of solar nebular oxygen (Fig. 4). Furthermore, since some molecules found in comets, e.g., C4H (Geiss et al. 1999), can be traced back to the dark molecular cloud from which our Solar System formed, we can use cometary volatiles and their abundances to study the processes that led from the molecular cloud through accretion into the solar nebula, to the present constituents of cometary nuclei (e.g., Altwegg et al. 1999). Until about 1980, it was assumed that comet nuclei primarily contained frozen H2O, NH3, and CH4. However, models based on this assumption, even when supplemented with a few other minor species, failed to explain the radicals and ions identified in spectra of cometary

Figure 4. Elemental abundances of carbon (C), oxygen (O) and nitrogen (N) in different bodies of our Solar System as a function of hydrogen/silicon ratio (after Geiss 1988).

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comae. One of the first models to come close to explaining the observations was based on the assumption that cometary nuclei consisted of frozen interstellar molecules and grains (Biermann et al. 1982; Greenberg 1982). Since that time, considerable progress has been made in finding new molecular species in comets and in modeling and understanding their compositions in this context (e.g., Altwegg et al. 1999). Figure 5 shows the abundances of cometary molecules compared with those of molecules in molecular clouds having C/O ratios near solar. Such data are becoming increasingly available (e.g., Ehrenfreund and Charnley 2000). The relative abundances of most cometary molecules measured so far are similar to those seen in interstellar dense clouds; the ranges of observed molecular abundances, relative to H2O, generally overlap in these objects. This suggests that much of the volatile component in comets has been preserved from the preceding cold molecular cloud stage. However, the degree of processing experienced by their ices is a fundamental question in cometary science, and indeed this is the primary driver for building a new taxonomy of comets based on composition. One of the defining characteristics of comets is their high abundance of volatiles (ices). As mixtures of nebular and interstellar ices, their compositions may be sensitive indicators of spatial and temporal variations in the nebular thermal environment. The dominant cometary ice phases (polar and apolar) are major reservoirs of primordial volatile oxygen. Measuring the compositions of cometary native ices (i.e., those contained in the nucleus) can therefore provide unique constraints on the origin and history of O-bearing icy materials in the Solar System. These data reflect the degree to which the composition of organic pre-cometary ices varied with distance from the young Sun (and with time) in the early solar nebula. Their structure and composition depend on local conditions (chemistry, temperature, degree of radiation processing) prevalent when and where they formed (Mumma et al. 1993; Irvine et al. 2000; Bockelée-Morvan et al. 2004). Measured cometary ice compositions can be compared with interstellar ices, with laboratory-processed analogs, and with formation models of comets and the proto-solar nebula.

Figure 5. A comparison between the relative abundances of various molecular species in ices found in dense molecular clouds (triangles), in gases in “hot cores” within dense molecular clouds (squares), in the gas phase in cometary comae (circles), and the range of abundances found in gas phase molecules in dense molecular clouds (cross-hatched bars). All abundances are normalized to the abundance of H2O, which has been assigned a value of 100. Arrows indicate upper limits. After Crovisier (1998).

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Oxygen plays a prominent role in the volatile composition of comets. Most of the volatile oxygen is in the form of H2O, and since this is the most abundant ice in comets, it is generally used as the reference against which the abundances of other ices are measured. In terms of oxidized carbon, abundances of CO (relative to H2O) are highly variable among comets, ranging from less than 1% to nearly 20%. Abundances of H2CO and CH3OH as high as ~3% and ~7%, respectively, have been reported (see, e.g., Fig. 12 of Bockelée-Morvan et al 2004). CO2, although observed in only a few comets, appears to be present at levels of several percent. By contrast, hydrocarbon molecules (CH4, C2H2, C2H6) provide an overall lower contribution to the total budget of volatile carbon. Abundances of the simplest fully reduced one-carbon molecule, CH4, although variable among comets, are at most CH4/H2O ~ 2% (see Gibb et al. 2003), and C2H2 and C2H6 are even less abundant (e.g., C2H6/H2O ~ 0.6% and C2H2/H2O ~ 0.2% in most Oort cloud comets observed to date). This is surprising in view of the very hydrogenrich composition of the solar nebula, in which comets formed. It is consistent, however, with material formed in an interstellar cloud or in a kinetically controlled solar nebula, where CO can either condense directly in polar (H2O-rich) or apolar (H2O-poor) ices (see below). Cometary nuclei warm when approaching the Sun, causing their ices to sublime, releasing volatiles into their comae (i.e., atmospheres), where they can be sensed spectroscopically at infrared wavelengths (principally between 2.8 and 5.0 μm) and sub-millimeter wavelengths. A fundamental challenge for cometary observations is distinguishing direct release by the nucleus (parent volatiles) from sources of extended release in the coma (e.g., by thermal degradation of grains or by chemistry). The spectral signature of volatiles consists of emission lines arising from fluorescent vibrational excitation by incident solar radiation. Although the presence of such lines in cometary spectra was predicted previously (Mumma 1982; Crovisier and Encrenaz 1983; Weaver and Mumma 1984; Bockelée-Morvan and Crovisier 1987), their detection required the development of astronomical spectrometers having sufficiently high sensitivity and spectral resolving power (e.g., λ/Δλ ~ 2 × 104 or higher). Modern IR spectrometers have small (sub-arc-second) pixels, and so are well-suited for measuring molecular abundances of volatiles in comets. They also provide spatial coverage of the sky, thereby permitting an accurate measure of the spatial distribution of emission in the coma. The emission intensity for direct release is highly peaked at the nucleus and decreases approximately as inverse projected distance from the nucleus, whereas distributed release gives rise to a flatter spatial profile of emission. This is particularly relevant to CO and H2CO, for which significant, or even dominant, distributed source contributions have been observed in some comets (Eberhardt et al. 1987; Meier et al. 1993; Wink et al. 1997; Eberhardt 1999; DiSanti et al. 1999, 2003). Advances in instrumentation over the past decade now enable the routine detection of multiple parent volatiles in comets, primarily through resonant (i.e., fundamental-band) rovibrational transitions occurring between excited and ground vibrational states. Detected molecules include CO (Fig. 6A), monomeric formaldehyde (H2CO, Fig. 6B), and methyl alcohol (CH3OH, Fig. 6C). Together with CO2, these represent the principal reservoirs of volatile oxidized carbon in comets. Ground-based detection of H2O requires sensing non-resonant (“hot band”) transitions (Figs. 6A, 6D) that occur between two excited vibrational states that are not significantly populated in the terrestrial atmosphere (Dello Russo et al. 2000). Prompt emission from OH was also proposed for sounding H2O in comets (Mumma 1982; Bockelée-Morvan and Crovisier 1989; Mumma et al. 2001). In contrast to the relatively flat spatial distribution of emission from fluorescent OH, the spatial profile of OH prompt emission faithfully traces that of parent (H2O). Through simultaneous measurement with H2O, a method has been developed for using OH prompt emission lines to quantify H2O production in comets (Bonev et al. 2004; Bonev 2005). These OH lines can therefore be used to establish both the spatial distribution and production rate of H2O when emissions from H2O itself are not available for observation within the band pass used.

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Figure 6. Spectra of oxygen-bearing molecules in comet C/2004 Q2 (Machholz) obtained with the Near IR Spectrometer (NIRSPEC) at the Keck Observatory. In each panel, the upper trace shows the extracted spectrum (solid) including dust continuum emission, and the superimposed modeled atmospheric transmittance (heavy dashed trace). The bottom trace shows residual cometary emission in excess of the continuum. A. Simultaneous measure of CO fundamental and H2O hot band emission near λ = 4.7 μm. CO lines are labeled by their rotational designations, and are blue-shifted relative to their terrestrial counterparts due to the geocentric velocity of the comet (~ −20 km s−1). B. Region of H2CO emission near 3.6 μm, showing the ν1 Q-branch plus four lines of OH prompt emission, used as a proxy for H2O. C. CH3OH ν3 band emission, centered near 3.52 μm, with P-, Q-, and R-branches indicated. D. H2O hot band emissions near 2.9 μm. Detailed treatment of these and other comet Q2 spectra can be found in Bonev (2005).

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The abundances (i.e., the production rates) of parent volatiles are determined by applying a quantum mechanical fluorescence model to the observable ro-vibrational lines. In most cases the analysis requires development of new (or extension of existing) fluorescence models appropriate to the low temperatures (20-150 K) typical of cometary comae. Such fluorescence models need to be formulated from high-resolution laboratory spectra, synthesized at the desired rotational temperature, and convolved to the instrumental resolution. Interpretation of infrared emission from H2CO (Reuter et al. 1989; DiSanti et al. 2006), CH3OH (Reuter 1992), and H2O (Dello Russo et al. 2000, 2005) has been approached in this manner. Because CO is a linear molecule having a relatively simple spectral signature, its rotational temperature can be measured directly by comparing the flux (W m−2) contained in lines spanning a range of rotational energies (Herzberg 1950; DiSanti et al. 2001). More recently, a general method has been developed for measuring rotational temperatures of any molecular species for which a fluorescence model exists over a range of temperatures (Dello Russo et al. 2005; DiSanti et al. 2006). Besides H2O, CO, CH3OH, and H2CO, the other main O-containing gas phase species in cometary comae is CO2 (e.g., Feldman et al. 1986; Crovisier et al. 1997), which has typical abundances relative to H2O of 1-10% (Bockelée-Morvan et al. 2004). A host of other O-containing species have also been identified in comets, including HCOOH, HCOOCH3, CH3CHO, NH2CHO, HNCO, OCS, and SO2, but these species are generally present at abundances well below 1% that of H2O (see Bockelée-Morvan et al. 2004 for a review). The discovery of cometary CO during spacecraft ultraviolet observations of Comet West (1976 VI) (Feldman and Brune 1976; Feldman 1978) established this molecule as an important component in these primitive Solar System objects. Since then, CO has been observed extensively at UV (Festou et al. 1982; Feldman et al. 1997), IR (Mumma et al. 2003, and references therein), and radio (Biver et al. 2002) wavelengths, and it has become the cornerstone for studies of oxidized carbon in comets. Pure CO ice has the lowest sublimation temperature (~25 K) of any parent molecule, so fractionation of CO-rich, pre-cometary ices should depend strongly on local temperatures. Alternatively, if trapped in H2O-rich ices, CO can be captured at higher temperatures (~50 K, or perhaps even to 150 K; Sandford and Allamandola 1988, 1990; Crovisier and Encrenaz 2000). In the IR, CO emission has been detected in all Oort cloud comets observed since 1996, and the large variation in its native abundance relative to H2O (1–20%) suggests that pre-cometary ices experienced a range of temperatures (indeed, Oort cloud comets formed in the giant planets’ region, ~5-30 AU from the Sun, over which local temperatures would have varied greatly). However, the abundances of CO and CH4 (the next most volatile parent molecule in comets, after CO) are not correlated among the comets studied (Gibb et al. 2003), suggesting that thermal effects alone cannot explain volatile abundances in cometary nuclei. A plausible means of converting CO to H2CO and H2CO to CH3OH involves hydrogen atom addition reactions on the surfaces of pre-cometary grains. This process is analogous to the H-atom addition (e.g., to C2H2) proposed to explain the high abundance ratio of C2H6/CH4 observed in C/1996 B2 (Hyakutake) (Mumma et al. 1996) and subsequent comets. The relative abundances of native CO, H2CO, and CH3OH to each other and to H2O in comets provide a measure of the efficiency of hydrogenation of CO on grain surfaces. UV and proton irradiation of laboratory analogs of pre-cometary ice (mixed H2O and CO) demonstrated the production of HCO, H2CO, CH3OH, and formic acid (Bernstein et al. 1995; Hudson and Moore 1999). Hydrogen-atom irradiation, both of polar (mixed H2O, CO) ice and of apolar (pure CO) ice, showed the conversion efficiency to be highly dependent on temperature in the 10–25 K range and on the density of H atoms (Hiraoka et al. 2002; Watanabe and Kouchi 2002; Watanabe et al. 2004). Hydrogen-atom addition to CO proceeds through the highly-reactive formyl radical (HCO). This subsequently converts to formaldehyde polymers such as polyoxymethylene (POM) and related derivatives (Huebner et al. 1987; Meier et al. 1993; Schutte et al. 1993), and also to H2CO and then to CH3OH. The relative abundances of CO, H2CO and CH3OH test their inter-relationship in comets, and the laboratory yields provide a basis for comparison in assessing natal conditions. ,1

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The presence of formaldehyde and related compounds in comets may also have been important for the origin of life. For example, H2CO has been proposed as the principal one-carbon molecule capable of generating complex organics of biological importance (Weber 2000; 2002). H2CO may also have played a central role in the formation of amino acids in primitive meteorite parent bodies during aqueous alteration via Strecker-cyanohydrin synthesis. Alternative mechanisms for producing amino acids in pre-cometary ices without liquid water have also been identified (Bernstein et al. 2002a; Muñoz Caro et al. 2002). These studies either employ HCN, NH3, and H2CO directly, or generate them in situ (Moore and Hudson 2003). Cometary organic molecules may also have contributed to the formation of nucleobases (Oró 1960), sugars (Weber 2002), possible pre-RNA backbones (Nelson et al. 2000), and a host of biochemical intermediates (Oró et al. 1990) either on the icy body or after accretion to the early Earth.

The oxygen-bearing minerals in “cometary” IDPs and samples from comet 81P/Wild 2 The remote detection of oxygen-bearing minerals in comets. The presence of O-bearing minerals in comets can be detected at infrared wavelengths, primarily through the use of spectroscopy in the 8-13 μm region where the characteristic Si-O stretching vibrations of silicate minerals fall, but silicate features can sometimes be detected out to 40 μm. When small cometary grains are ejected from a cometary nucleus, they are heated by solar radiation. This energy is subsequently re-radiated in the infrared, resulting in a quasi-blackbody continuum upon which a superimposed silicate emission feature is sometimes seen. To produce a strong feature in emission, the emitting silicate particles must be smaller than 1 μm in radius; particles larger than this will, for the most part, only contribute to continuum emission, and their mineralogy is difficult to assess in this manner. The strength and profile of the emission features are therefore dependent on grain composition, size and albedo, and emission models must be used to interpret these features (see Hanner and Bradley 2004 for a review). The infrared emission features of dust in cometary comae are generally interpreted to be largely dominated by olivine, pyroxene, and “glassy” silicates (e.g., Hanner et al. 1994; Wooden et al. 1999; Crovisier et al. 2000; Hanner and Bradley 2004). Models of the silicate emission feature suggest that only a minor fraction of the silicates (15-30%) are in crystalline form; the remainder of the silicates may be amorphous. A unique opportunity to study cometary dust remotely occurred when the Deep Impact mission sent a 364 kg impactor into the nucleus of comet 9P/Tempel 1 at 10.2 km/sec. The impact produced large amounts of ejecta from the surface layers of the comet, and infrared spectra of the ejecta were collected by the Spitzer Space Telescope (Lisse et al. 2006). Emission signatures due to amorphous and crystalline silicates, amorphous carbon, carbonates, phyllosilicates, polycyclic aromatic hydrocarbons, H2O gas and ice, and sulfides were reported. Many of these materials are seen in chondritic porous IDPs and samples from comet 81P/Wild 2 (see below), but others, for example phyllosilicates and carbonates, have yet to be confirmed in Wild 2 samples. The reason for these apparent differences has yet to be fully resolved. Overall, the spectral features observed in cometary infrared emission spectra imply a fairly complex mineralogy that includes both amorphous and crystalline grains, with olivine and pyroxene being the dominant crystalline phases. The mineralogical mixture is largely consistent with the composition of chondritic porous IDPs and samples returned from comet 81P/Wild 2, although it is difficult to assess the abundance of glassy silicates in the Wild 2 samples (see the following two sections). Oxygen-bearing minerals in “cometary” IDPs. Interplanetary dust particles are collected in the stratosphere at 20-25 km altitude and are typically 5-15 μm in diameter. There are two major types of IDPs: chondritic porous (CP) IDPs that are typically anhydrous; and chondritic smooth (CS) IDPs that have undergone aqueous alteration that formed clay minerals and carbonates. The CP IDPs have been linked to cometary sources from their inferred orbital characteristics (Brownlee et al. 1995), fine-grained mineralogy (Bradley and Brownlee 1986), infrared spectral properties (Sandford 1991), and high abundances of presolar materials

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(Messenger et al. 2003). These particles have escaped the thermal metamorphism and aqueous alteration that affected even the most primitive meteorites, and are characterized by high carbon and nitrogen abundances (Thomas et al. 1993; Keller et al. 2004; Flynn et al. 2006), unequilibrated mineralogy (Keller and Messenger 2005), and the presence of non-solar hydrogen and nitrogen isotopic signatures and abundant presolar silicates (Messenger et al. 2003). Typical CP IDPs are highly porous particles that consist of fine-grained crystalline silicates, GEMS (glass with embedded metal and sulfides, Bradley 1994) grains, and Fe-Ni sulfides, all bound together by an organic-rich carbonaceous matrix. The constituent grains in IDPs are much smaller ( 1), as might occur in deep melting zones. The curves sweep clockwise as r increases, representing a change from catazonal to mesozonal to epizonal emplacement conditions. The crystal-melt fractionation factor was taken as 0.9995 for all curves except one of the two simple fractional crystallization curves in the lower right, for which a value of 1.0005 was used. Modified after Criss (1999).

Real magmas While our ability to explain magmatic differentiation in terms of simple models needs improvement, the fact remains that the total observed range of δ18O values in terrestrial magmas, at least −2 to +16‰, is large and interesting (Taylor and Sheppard 1986). The low values are most important because most crustal processes tend to progressively increase rather than decrease magmatic 18O contents. This phenomenon of 18O-poor magmas has been reviewed by Taylor (1986), and is clearly associated with the assimilation of large quantities of hydrothermally-altered, 18O-poor country rock that commonly occur in regions of subaerial explosive volcanism. Similarly, assimilation of 18O-rich metamorphic or sedimentary material appears to be required at the other extreme. A key problem for explaining natural igneous suites in terms of simple models is that scattered relationships typically result when data for one chemical or isotopic system are plotted against another. A partial explanation is that the observed compositional variations for

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many components are actually quite small, and appear to be important largely because of the great precision of modern instruments. Even when a good correlation is found for a given data pair, a (pessimistic) rule of thumb is that, for each geochemical variable added to the data set, an additional compositional end-member is required in a model that “explains” the results. For example, while two reasonable end-members might be found that appear to explain a given pair of isotopic compositions for a given igneous suite (e.g., oxygen and Sr isotopes), the same end-members would probably not accurately explain Pb isotope data for the same rocks. This “Murphy’s Law of Geochemistry” is likely rooted in two realities. First, both the compositions and the ambient physical conditions are heterogeneous over the large vertical and lateral scales of magmatic systems; and second, different processes and different exchange and transport rates govern the various chemical elements or physical subsystems. The unsurprising result is that simple geochemical models can provide beautiful descriptions of small, simple systems, but often provide inadequate descriptions of complex, natural systems.

Subsolidus fractionation processes One of the most interesting and unexpected outcomes of oxygen isotopic studies of igneous rocks is the discovery that they are commonly and profoundly modified by subsolidus processes. In particular, oxygen isotopes prove that many rock suites have undergone extensive recrystallization and exchange with external reservoirs, even though they may appear to be petrographically fresh. This disconnect arises in part because low-temperature processes, such as weathering and oxidation, generally have a severe impact on rock appearance, yet these effects are typically not pervasive, so relict minerals that simply retain their original δ18O values normally dominate the assemblage. In contrast, if interaction and exchange between rocks and fluids occur at higher temperatures, within or near the stability fields of many important igneous minerals, rock appearance need not be degraded much. These conditions foster profound recrystallization of the silicate framework; isotopic exchange with ambient fluids; and production of new alteration assemblages that are faithfully recorded by the δ18O values and by the readjusted mineral fractionation relationships. Many different exchange, recrystallization and degassing processes modify the oxygen isotopic compositions of rocks. The best tool for identifying and quantifying such effects is the δ-δ plot, whose characteristics were defined by Gregory and Criss (1986) and Gregory et al. (1989). In particular, this diagram provides a topologically straightforward way to distinguish rocks that have undergone isotopic exchange in closed- or open- system environments. Basically, closed-system processes cannot alter the whole-rock δ18O values, but in open systems, where an oxygen-bearing liquid or gas infiltrates and/or escapes the rock volume, the whole-rock δ18O values can change, and all constituent minerals typically become richer or poorer in 18O. These closed-system and open-system effects tend to produce trends having negative or positive slopes, respectively, on δ-δ plots (Fig. 4). New techniques, such as ion probe isotopic analysis on small spatial scales, will probably yield even greater insights into these processes. Slow cooling of deep-seated crystalline rocks approximates exchange in a closed system, and commonly produces anomalously large oxygen isotopic differences among coexisting igneous minerals, which become obvious whenever “absurd” temperatures are calculated from the observed isotopic fractionations. Such effects are especially well expressed in wet, deeply emplaced granitic plutons. An example is the Idaho batholith, where even the freshest specimens exhibit inter-mineral 18O fractionations that are 0.5 to 2‰ too large to correspond to magmatic temperatures. In marked contrast, rapidly quenched, dry magmas, as exemplified by lunar basalts, commonly exhibit small but regular inter-mineral 18O differences that are appropriate for magmatic temperatures (e.g., Mayeda et al. 1975). To the extent that slow cooling is a closed-system process, the oxygen isotopes are merely transferred from one mineral to another, with some minerals gaining 18O at the expense of others. Thus, closed-system processes mostly produce negative slopes on δ-δ plots, which is mathematically required when major phases are

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Figure 4. A δ-δ plot of coexisting quartz and feldspar, showing infinite temperature (α = 1.0) and magmatic (α = 1.001) isothermal lines, and the quadrants available to a rock having initial δ18O values of +9 and +10 for feldspar and quartz, respectively, undergoing exchange in open and closed systems. Phenocrysts from ash flow tuffs were quenched and thus preserve a record of their equilibration at magmatic temperatures. Also shown are trends for granitic rocks that were infiltrated by sea water (plagiogranites; Gregory and Taylor 1981) and low-18O meteoric-hydrothermal fluid (Idaho suites; Criss and Taylor 1986), which exhibit the positive slopes indicative of open system behavior. After Gregory and Criss (1986).

plotted against each other. Scattered relationships occur more commonly because the effects tend to be small, and the rates of isotopic exchange differ greatly among the various minerals, as does the dependence of these rates on the progressively changing temperatures. Far more profound are the effects caused by infiltration of rocks by fugitive fluids in open systems, as occurs in many metasomatic, hydrothermal or metamorphic environments (Gregory et al. 1989). Both high and low 18O values can be produced, but because normal igneous processes tend to progressively increase δ18O values, the low values that can be generated in open systems are very unusual and most amenable to clear-cut interpretation. Such effects are exemplified by interactions between igneous rocks and hydrothermal fluids derived from ordinary, 18O-poor meteoric waters in continental interiors or high-latitude locations. The western USA has more examples of such effects than any other large region in the world, due to its high heat flow, regional extension, extensive Tertiary magmatism, and the low δ18O values of −10 to −20‰ typical of its meteoric waters. These regional characteristics collectively foster the deep penetration and convective circulation of meteoric-hydrothermal fluids and enhance their potential to undergo oxygen isotope exchange with host rocks, which mostly occurs at temperatures of 150 to 350 °C. Notable isotopic effects produced by meteoric-hydrothermal activity include: 1) reduction of whole-rock δ18O values, commonly to values of +4 to −6‰; 2) the development of low-18O terranes featuring regular 18O zonation around volcanic centers, epizonal igneous intrusions, or faults; 3) the persistence of such low-18O patterns over vertical distances of 1 to 10 km, and over lateral distances up to 10s of kilometers; 4) the development of extremely large and highly anomalous oxygen isotopic fractionations between coexisting minerals, producing highly elongated, positive-sloped trends on δ-δ plots (Fig. 4); and 5) the association of anomalous

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δ18O values with anomalous δD values, largely due to the fact that 18O-poor meteoric fluids are proportionately depleted in deuterium. Criss and Taylor (1986) provide many examples and details. From a global perspective, the effects of hydrothermal alteration about oceanic spreading centers are more important than effects due to meteoric fluids. Pervasive interactions of the spreading-center igneous suite with heated, circulating seawater produces both 18O enrichments and 18O depletions, depending on the temperature of interaction, as this controls the effective fluid-rock fractionation factor (Fig. 5). The net result of these interactions is the production of a vertical gradient in the δ18O values of the ocean floor, from relatively low to relatively high values with decreasing depth, with essentially no net change in the bulk δ18O value of the affected rock sequence. Basically, the δ18O value of ocean water is buffered by fluid-rock interactions at oceanic spreading centers, as pointed out by Muehlenbachs and Clayton (1976) and Gregory and Taylor (1981). 8

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OXYGEN ISOTOPE ZONATION AND HETEROGENEITY IN PLANETARY LITHOSPHERES Processes producing 18O zonation Fractional crystallization. The similarity between Bowen’s series and the relative order of 18O incorporation among common igneous minerals (Eqn. 2; Fig. 2) explains why igneous suites tend to evolve toward progressively higher δ18O values. This similarity is a physiochemical effect related to increasing silica polymerization, and stronger cation-oxygen bonds, in the progressively later members of Bowen’s series. Because magmas commonly rise as they cool, this process can produce vertical 18O gradients in affected crust. Such effects are rather small, as discussed above, though they can be magnified by AFC processes.

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Density effects. On the most basic level, vertical zonation of 18O in planetary lithospheres occurs because the tendency of minerals to incorporate 18O is strongly and negatively correlated with their densities. Figure 6 shows the relationship between mineral density and all of the various calcite-mineral fractionation constants given in Table 5 of Chacko et al. (2001). Clearly, dense minerals tend to have a smaller affinity for 18O than light minerals. Practically any process that tends to stratify rock density will therefore tend to stratify 18O as well. Water, with a density of 1, is off the scale of Figure 6 and deviates strongly from the indicated mineral trend. Convective 18O pumping. Convective processes also have great potential to produce 18O zonation. Vertical 18O increases are produced in oceanic hydrothermal systems (Fig. 5) and are evident in the wall rocks hosting meteoric-hydrothermal systems (Criss and Taylor 1986). Strong vertical 18O increases are also commonly observed in hydrothermal vein deposits, especially those precipitated by ascending fluids following the hydrostatic boiling curve (Larson and Taylor 1987). It is useful to consider convection as a process capable of redistributing 18O as well as heat. Imagine fluid circulating in a closed cell in an initially homogeneous, permeable rock reservoir subjected to a temperature gradient. The same temperature gradients that produce the buoyancy differences that induce the fluid to convect also generate differences in the rockfluid fractionations, with the latter being smallest where temperatures are highest (Fig. 7). Any fluid-rock exchange will therefore result in 18O being systematically transferred from the hottest to the coolest zones, creating a series of 18O iso-surfaces that tend to parallel the ancient isothermal surfaces. Such iso-surfaces have been mapped in a large, 3D block of hydrothermally altered volcanic rock in the Comstock Lode mining district, and can be visualized in sections and stereopairs (Singleton and Criss 2004). The redistribution of 18O by convective processes is not restricted to heterogeneous systems where fluids flow through permeable rock substrates; it can, theoretically, occur in 1 Silicates Oxides and Carbonates

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Figure 6. Graph illustrating the strong negative correlation between the constant C3 of Equation (1b) and mineral density, where the relevant fractionation equations are between the indicated mineral and calcite. All values of C3 were taken from Table 5 of Chacko et al (2001). Dense minerals tend to have low values of C3 and accordingly have the least affinity for 18O.

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O, ‰ Figure 7. Conceptual models showing the effects of hydrothermal fluid circulation and exchange on the distribution of 18O in host rocks with an initial value of +6 (vertical line). Temperature increases linearly with depth and controls the effective rock-fluid fractionation factor. The “open” curve depicts exchange effects in a partially-equilibrated open system in which permeability decreases exponentially with depth; compare Figure 5. The “equilibrated” curve shows a system that everywhere has equilibrated with a circulating fluid having a uniform composition.

homogenous, viscous convective systems such as the upper mantle. Significant effects are predicted for systems undergoing some transformative process such as crystallization at the upper or lower boundary layers, as would occur in a magma ocean. Example calculations for a crystallizing magma chamber mostly yield 18O increases with increasing depth (Fig. 8).

Processes producing 18O heterogeneity Temperature effects. Oxygen isotopic heterogeneity is mostly produced by fractionation processes that are primarily temperature-driven. Fractionation factors generally increase as temperature decreases, tend to be large between solids and liquids or gases, and can be magnified by non-equilibrium or biological processes. Thus, at the high temperatures in the mantle, the δ18O values of important phases are within about 3‰ of each other. In marked contrast, a researcher standing on South Pole snow with a δ18O of −50‰ would breathe normal air with a δ18O of +23‰, containing small amounts of atmospheric CO2 and water vapor with values near +41 and −75‰, respectively. Even larger 18O differences probably occur on icy moons and in other cold environments in the outer solar system. Figure 9 shows the relationship between the δ18O of common terrestrial materials and their temperatures of occurrence or formation. Reservoir effects. Oxygen isotopic heterogeneity also depends on the mass balance relationships in the systems hosting the phases. Given a particular fractionation factor, greater isotopic heterogeneity can be produced in open systems than in closed systems. Open-system conditions are exemplified by Rayleigh fractionation behavior, which can produce extreme isotopic enrichments or depletions in the volumetrically-small terminal stages. A simple rule of thumb is that, the more unusual the isotopic composition, the less such material will be present. Figure 10 shows the relationship between the δ18O values of common terrestrial materials and their aggregate mass in the Earth.

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Figure 8. Conceptual models showing the effects of homogeneous convection in a crystallizing magma chamber on the vertical distribution of 18O. A. Rayleigh crystallization of a mafic, 18O-poor phase progressively removed by crystal settling at the bottom of the magma chamber. B. Rayleigh crystallization of a low-density, felsic, 18O-rich phase removed by crystal flotation to the top of the magma chamber. M,F. Rayleigh evolution of a magma chamber featuring crystallization and settling of a dense mafic phase (M) coupled with simultaneous crystallization and flotation of a felsic phase (F), such that the last residual magma occurs in a thin “sandwich horizon” (SH) at the center of the magma chamber. Arrows indicate the direction of progress. All curves but F illustrate an upward increase in 18O.


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Figure 10. Graph of the δ18O values of common terrestrial materials vs. the common log of the aggregate mass of oxygen in the indicated reservoir. Isotopic heterogeneity increases with decreasing reservoir size.

Bulk 18O composition of the continents Several lines of evidence point to upward enrichment of 18O in the Earth. In oceanic crust, the upward 18O increases produced by hydrothermal interactions are well documented (Fig. 5). Similarly, several processes working in concert have produced continents with higher average δ18O values than the +5.7‰ value of the bulk Earth. The first task is to establish and quantify this latter effect. Owing to its high average SiO2 content of about 59 ± 1%, the mean δ18O value of continental crust must be significantly higher than that of the mafic and ultramafic rocks of the ocean floor and the upper mantle. The average density of about 2.75 g-cm−3 and other continental properties suggest a mean composition similar to granodiorite, quartz diorite, or andesite (e.g., Henderson 1982), so the 18O values typical for such lithologies provide a useful estimate of the bulk continental δ18O value. Upper crustal materials such as sedimentary and metamorphic rocks generally have much higher δ18O values, though their contribution to the average is probably small because such lithologies likely constitute only 5-20% of the total continental mass. For example, Valley et al. (2005) believe that sediments represent 14% of the continental crust and have an average δ18O of +17. Low-18O rocks produced by meteoric-hydrothermal activity are even less abundant and cannot significantly offset the high values of the latter lithologies. All things considered, the average δ18O value for the continental crust is probably about +7.5 ± 0.5. Note, however, that Valley et al. (2005) estimate that the average δ18O of continental crust is +9 to +10. Regardless, it appears that the continental crust is 18O-enriched relative to the Earth’s upper mantle. A useful question is, from what source was the high 18O content of continental material derived? Considering that the upper mantle is at least 30× more massive than continental crust, the seemingly obvious answer is that the continental 18O “excess” was derived through a process producing a trivial complementary 18O reduction, amounting to < 0.1‰, in the upper mantle. The only other possibility is that a smaller reservoir, having a very low δ18O value,

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was generated to counterbalance the high 18O value of the continents. Earth’s hydrosphere is the only reservoir that is massive enough, and sufficiently low in 18O (−1‰), to be significant in this regard. Given that the ocean covers 71% of Earth’s surface and has a mean depth of 3.8 km, the mass ratio of hydrosphere to continental crust has been variously estimated as being between 1:9 and 1:15, depending in large part on the assumed mean thickness of continental crust. While normal rocks are about 50% oxygen by weight, water is 89% oxygen, however, so in terms of oxygen content, these relative mass ratios become about 1:5 to 1:8. Now, equation 1.11b of Criss and Farquhar (2008) may be used to define the bulk δ18O value of the “system” comprising the hydrosphere plus continental crust, as illustrated below for the 1:5 case: δbulk = (1/6)(−1) + (5/6)(7.5 ± 0.5) = +6.1 ± 0.4

(3)

Use of the 1:8 relative ratio would produce a higher average, +6.6 ± 0.4, but considering that this high continental mass is derived by assuming that the continents are very thick (~50 km), the bulk δ18O of average continental material would probably be lower, +7.0 ± 0.5‰, and the average 18O value calculated for the continent-ocean “system” would again be ~ +6.1 ± 0.4. Several surprising points emerge from these elementary considerations. First, the hydrosphere is an important part of the 18O balance of the uppermost layers of the solid Earth, not just a miniscule reservoir whose composition is controlled by the latter. Second, the net 18Oenrichment in the outermost part of Earth is small, probably being only +0.4 ± 0.4‰ higher than the bulk upper mantle value of +5.7‰. Third, the genetic link between continental crust and the hydrosphere, long suspected from the absence of andesites and granites in meteorite and lunar collections, is quantitatively established by 18O balance. Finally, while small net 18Oenrichment of the uppermost layers occurs, the primary effect of dynamic planetary processes is the magnification of 18O differences among interacting reservoirs at progressively lower temperatures.

Isotopic changes over geologic time Plots of δ18O of a given material as a function of stratigraphic depth or geologic age provide important information about systematic changes over geologic time. Two types of changes are evident. Diffusion, recrystallization, mixing and related processes attenuate 18O differences among isotopically distinct but otherwise identical materials over the time of their mutual contact. In contrast, secular changes tend to magnify differences between samples formed today and similar materials formed progressively further back in time. Regarding the first effect, the observed 18O heterogeneity among similar, proximal materials tends to decrease with the length of the record examined. A few examples will illustrate this point. At the brief end of this scale, over a normal year, the δ18O values of ordinary meteoric precipitation at many single locations in continental interiors vary by 10 to 25‰. Such differences are attenuated in shallow ground waters, which over annual to decadal time scales vary by 1 to 5‰, forming the basis for a quantitative estimate of ground water residence time (Criss 1999). Similarly, the seasonal 18O differences recorded in high-latitude ice cores become progressively attenuated with increasing age, from ~15‰ today to 2.5 Ga), which preserve a record of the existence of 18O-rich lithosphere ranging back to the earliest surviving vestiges of crustal material formed at 4.4 Ga, but Valley et al. (2005) believe that, in younger zircons, the range and average δ18O of magmas appear to have increased significantly. They attribute this to the influence of recycled continental crust.

CONCLUSIONS Oxygen isotopes provide key data on Earth’s formation and evolution and on the interactions between its lithosphere, hydrosphere, biosphere and atmosphere. The δ18O values of major lithospheric reservoirs such as MORB closely approximate those of lunar igneous rocks (+5.7 ± 0.3‰). This similarity suggests that Earth and Moon have a common primordial affinity and that +5.7‰ approximates the bulk oxygen composition of the EarthMoon system, but only surficial materials representing the outermost shells of these bodies have been sampled. Many processes foster enrichment and increased heterogeneity of 18O contents with decreasing depth in planetary lithospheres. Upward 18O enrichment is promoted by convection, hydrothermal alteration, fractional crystallization, and the tendency for the least dense silicate and oxide minerals to concentrate 18O relative to more dense minerals. Sediments in general and, more importantly, the continents and the uppermost layers of the ocean floor, therefore tend to be much richer in 18O than average upper mantle material. Increased 18O heterogeneity is promoted by the strong tendency for isotopic fractionation factors to increase as temperature decreases, and by the increased potential for complex interactions to occur between solid, liquid and gaseous phases near planetary surfaces. Thus, a large total range is observed in sediments (at least −4 to +42‰), terrestrial magmas (−2 to +16‰), and terrestrial igneous rocks (−10.5 to > +16‰). Only small effects (< 2‰) can be attributed to pure fractional crystallization, as isotopic fractionations are small at high temperatures; this partly explains why felsic rocks are systematically higher in 18O by 1-4‰ than mafic and ultramafic rocks, but cannot explain the total ranges in terrestrial rocks. Earth’s large 18O ranges require interaction or exchange of rocks and magmas with oxygen reservoirs located or formed at or near Earth’s surface, where large enrichments or depletions in 18O are possible. Key processes that have been recognized include formation of minerals in equilibrium with the hydrosphere or atmosphere; subsolidus exchange of rocks with infiltrating fluids; and the incorporation into magmas of wallrocks having high or low δ18O values generated under low-temperature conditions. Subsolidus alteration effects are easily identified on δ-δ plots.

REFERENCES Bowen NL (1928) The Evolution of the Igneous Rocks. Princeton University Press, Princeton Chacko T, Cole DR, Horita J (2001) Equilibrium oxygen, hydrogen and carbon isotope fractionation factors applicable to geologic systems. Rev Mineral Geochem 43:1-81 Criss RE (1999) Principles of Stable Isotope Distribution. Oxford University Press, Oxford Criss RE, Farquhar J (2008) Abundance, notation, and fractionation of light stable isotopes. Rev Mineral Geochem 68:15-30 Criss RE, Taylor HP Jr (1986) Meteoric-hydrothermal systems. Rev Mineral 16:373-424 Eiler JM (2001) Oxygen isotope variations of basaltic lavas and upper mantle rocks. Rev Mineral Geochem 43:319-364

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Foulger GR, Natland JH, Presnall DC, Anderson DL (eds) (2005) Plates, plumes, and paradigms. Geol Soc Am, Special Paper 388. Geological Society of America, Denver, Colorado Gregory RT, Criss RE (1986) Isotopic exchange in open and closed systems. Rev Mineral 16:91-127 Gregory RT, Criss RE, Taylor HP Jr (1989) Oxygen isotope exchange kinetics of mineral pairs in closed and open systems: Applications to problems of hydrothermal alteration of igneous rocks and Precambrian iron formations: Chem Geol 75:1-42 Gregory RT, Taylor HP Jr (1981) An oxygen isotope profile in a section of Cretaceous oceanic crust, Samail ophiolite, Oman: Evidence for δ O-buffering of the oceans by deep (>5 km) seawater-hydrothermal circulation at mid-ocean ridges. J Geophys Res 86:2737-2755 Henderson P (1982) Inorganic Geochemistry. Pergamon Press, New York Hoefs J (2004) Stable Isotope Geochemistry. Springer, New York Johnsen SJ, Dansgaard W, Clausen HB, Langway CC (1972) Oxygen isotope profiles through the Antarctic and Greenland ice sheets. Nature 235:429-434 Keith ML, Weber JN (1964) Carbon and oxygen isotopic composition of selected limestones and fossils. Geochim Cosmochim Acta 28:1787-1816 Knauth LP, Lowe DR (1978) Oxygen isotope geochemistry of cherts from the Onverwacht group (3.4 billion years), Transvaal, South Africa, with implications for secular variations in the isotopic composition of cherts. Earth Planet Sci Lett 41:209-222 Kyser TK (1986) Stable isotopic variations in the mantle. Rev Mineral 16:141-164 Kyser TK (ed) (1987) Stable Isotope Geochemistry of Low Temperature Fluids. Mineralogical Society of Canada, Quebec Larson PB, Taylor HP Jr (1987) Solfataric alteration in the San Juan Mountains, Colorado: Isotopic variations in a boiling hydrothermal environment. Econ Geol 82:1019-1036 Lodders K, Fegley BJ (1998) The Planetary Scientist’s Companion. Oxford University Press, Oxford Mayeda TK, Shearer J, Clayton RN (1975) Oxygen isotope fractionation in Apollo 17 rocks. Proc Lunar Sci Conf 6th:1799-1802 Muehlenbachs K, Clayton RN (1976) Oxygen isotope composition of the oceanic crust and its bearing on seawater. J Geophys Res 81:4365-4369 Savin SM, Yeh HW (1981) Stable isotopes in ocean sediments. In: The Sea. Emiliani C (ed) John Wiley & Sons, New York 7:1521-1554 Sharp Z (2007) Principles of Stable Isotope Geochemistry. Pearson Prentice Hall, Upper Saddle River, New Jersey Singleton MJ, Criss RE (2004) Symmetry of flow in the Comstock Lode hydrothermal system: Evidence for longitudinal convective rolls in geologic systems. J Geophys Res 109:B03205 Spicuzza MJ, Day JMD, Taylor LA, Valley JW (2007) Oxygen isotope constraints on the origin and differentiation of the Moon. Earth Planet Sci Lett 253:254-265 Taylor HP Jr (1980) The effects of assimilation of country rocks by magmas on 18O/16O and 87Sr/86Sr systematics in igneous rocks. Earth Planet Sci Lett 47:243-254 Taylor HP Jr (1986) Igneous rocks: II. Isotopic case studies of circumpacific magmatism. Rev Mineral 16:273317 Taylor HP Jr, Sheppard SMF (1986) Igneous rocks: I. Processes of isotopic fractionation and isotope systematics. Rev Mineral 16:227-271 Urey HC (1947) The thermodynamic properties of isotopic substances. J Chem Soc (London) 562-581 Valley JW (1986) Stable isotope geochemistry of metamorphic rocks. Rev Mineral 16:445-489 Valley JW, Cole DR (eds) (2001) Stable Isotope Geochemistry. Reviews in Mineralogy and Geochemistry, Vol. 43. Mineralogical Society of America, Chantilly, Virginia Valley JW, Lackey JS, Cavosie AJ, Clechenko CC, Spicuzza MJ, Basei MAS, Bindeman IN, Ferreira VP, Sial AN, King EM, Peck WH, Sinha AK, Wei CS (2005) 4.4 billion years of crustal maturation: Oxygen isotopes in magmatic zircon. Contrib Mineral Petrol 150:561-580 Valley JW, Taylor HP, O’Neil JR (eds) (1986) Stable Isotopes in High Temperature Geological Processes. Reviews in Mineralogy, Vol. 16. Mineralogical Society of America, Chantilly, Virginia Wiechert U, Halliday AN, Lee DC, Snyder GA, Taylor LA, Rumble D (2001) Oxygen isotopes and the Moonforming giant impact. Science 294:345-34

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 527-553, 2008 Copyright © Mineralogical Society of America

Basalts as Probes of Planetary Interior Redox State Christopher D. K. Herd Department of Earth and Atmospheric Sciences 1-26 Earth Sciences Building University of Alberta Edmonton, Alberta, T6G 2E3, Canada [email protected]

ABSTRACT Whether the redox state, quantified as oxygen fugacity, recorded in a planetary basalt is an accurate representation of the redox state of the planetary interior from which it was derived through partial melting, ascent, eruption and emplacement is a fundamental question in planetary geology. In the absence of mantle xenoliths in samples from the Moon, Mars and differentiated asteroids, the basalt-mantle source relationship must be extrapolated from what is known about the Earth in order to probe the redox state of these planetary interiors. A review of current knowledge regarding the basalt-mantle source relationship for the Earth provides insights into the advantages and pitfalls of determining mantle redox state. The range of currently available oxybarometers, including thermodynamic models based on ferrous-ferric mineral equilibria and multivalent cation analysis are surveyed and their limitations presented. The result is a basis for the informed interpretation of the oxygen fugacity of planetary basalts, and new insights into the role of C-H-O volatiles in the terrestrial planets.

INTRODUCTION For the purpose of elucidating the redox evolution of terrestrial planet interiors, the Earth represents a natural, if not ideal, laboratory in which to examine the relationship between the oxygen fugacity (fO2) of a basaltic sample and the redox state of its mantle source. Partial melting of terrestrial planet interiors to produce basaltic eruptives is a fundamental process that was initiated on the terrestrial planets shortly after their formation and has continued in some cases to the present day. The redox characteristics of a basalt are the direct results of the physical and chemical conditions of partial melting, ascent and eruption. The oxygen fugacity of a basalt from the Earth can be determined by the judicious selection of one or more pertinent oxybarometers based on ferrous-ferric mineral equilibria or multivalent trace element characteristics; comparison of its oxygen fugacity with that of a related mantle xenolith, complemented perhaps by insights from laboratory experiments, places constraints on the behavior of multivalent elements (and redox-sensitive volatiles) during partial melting. In this way, the influence of the physical and chemical conditions of partial melting of the mantle source on the oxygen fugacity of the basaltic product can be quantified, or at least characterized. The same cannot be done for samples of the Moon, Mars and differentiated asteroids (e.g., 4 Vesta), due to the apparent absence of mantle xenoliths from our samples of these suites. Instead, the basalt-mantle source relationship, as quantified from studies of terrestrial samples, must be adapted for and extrapolated to the suite of basaltic samples from these other planetary bodies. For this reason, summaries of the oxidation state of the Earth’s mantle, the array of available oxybarometers applicable to basaltic samples, and a discussion of whether the 1529-6466/08/0068-0019$05.00

DOI: 10.2138/rmg.2008.68.19

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oxygen fugacity of a basalt is reflective of the oxidation state of its mantle source are important precursors to the presentation of data for planetary basalts, and are essential to their proper interpretation. Oxygen fugacity is not a straightforward concept. A summary of the origin of the concept and a discussion of common misconceptions is provided by Frost (1991). The misconception most relevant to the following discussion is that oxygen is present as a fluid species in natural systems. Although we express oxygen fugacity in units of gas pressure (for example, 10−9.3 bars of O2), oxygen fugacity does not represent the partial pressure of a gas, but instead monitors the chemical potential, and can therefore be used to describe a condensed system in which no free oxygen is present (Frost 1991). Therefore, the “O2” term in many of the equations presented in this chapter does not necessarily indicate involvement of free oxygen. Oxygen fugacity is most often expressed relative to an assemblage of pure phases that define a specific oxygen fugacity for a given temperature. These assemblages are known as solid oxygen buffers or petrologic buffers; they are not necessarily present in natural systems, but represent convenient references to describe the oxygen fugacity of a natural system. Common examples are iron-wüstite (IW), fayalite-magnetite-quartz (FMQ), nickel-nickel oxide (NNO), and magnetite-hematite (MH). A number of studies have been carried out over the past several decades to define the buffer equations, resulting in a number of formulations. A selection is given in Table 1, and other compilations can be found in Frost (1991) and Chou (1987). The oxygen fugacity defined by the buffers increases with increasing temperature, and the slopes of the buffers are nearly the same in all cases. An oxygen fugacity estimate calculated using an oxybarometer (see below) can be expressed relative to one of the buffers, obviating the need to state both the absolute oxygen fugacity and temperature. For example, an oxygen fugacity estimate of log fO2 = −9.3 at a temperature of 1200 °C and a pressure of 1 bar corresponds to 1 log unit below FMQ, or “FMQ − 1”. The most recent data for solid oxygen buffers come from O’Neill and Pownceby (1993); regression of their data for the IW and NNO buffers is provided in Table 1. Also provided are regressions of data from O’Neill (1987a) for NNO, O’Neill (1988) for IW, and O’Neill (1987b) for FMQ. The latter equation should be used instead of the erroneous equation of Holloway et al. (1992) for FMQ based on the same data. Many of the formulations do not explicitly account for pressure effects, and often the buffers are not experimentally calibrated for high pressures; additional buffer equations that include pressure terms are provided by Ballhaus et al. (1991). The use of different buffer formulations can result in small discrepancies; for example, the use of the NNO formulation according to Schwab and Küstner (1981) for log fO2 = −9.3 and T = 1200 °C yields NNO − 1.8, whereas use of the regression of the O’Neill and Pownceby (1993) data (Table 1) yields NNO − 1.6. Although the discrepancies tend to be small, typically less than the uncertainty on an oxybarometer estimate, an explicit statement of the formulation used to calculate and express oxygen fugacity provides the reader with the ability to compare oxygen fugacity estimates without introducing added uncertainties.

THE OXIDATION STATE OF THE EARTH’S MANTLE The differences between oxygen fugacity, oxidation state and oxygen content, as outlined above, are fundamentally important for understanding the Earth’s mantle. Iron is the most abundant element that exists in more than one oxidation state in planetary interiors; for this reason the oxidation state of the Earth’s mantle is expressed in terms of iron oxidation state, i.e., the relative proportions of the different valence states of iron (Fe0, Fe2+ and Fe3+; commonly given as Fe3+/ΣFe, where ΣFe = Fe2+ + Fe3+). Oxygen fugacity is related to iron oxidation state through equilibria between Fe-bearing minerals. The valence state of iron in minerals in a natural system, such as the Earth’s mantle, however, is the result of the complex interplay of mineral crystal chemistry and the oxidation state of the system. The oxidation state of iron may thus change as a result of crystal chemical effects, without a correlative change in oxygen fugacity.

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Table 1. Equations* for solid oxygen buffers. Buffer IW Fe – Fe1−xO

FMQ Fe2SiO4 – Fe3O4 – SiO2

NNO Ni – NiO

MH Fe3O4 – Fe2O3

C

T range (°C)

A

B

Data Source(s)

−27589.7

6.790

769-1371

Regression of data from O’Neill and Pownceby (1993)

−27654.0

6.849

769-1177

Regression of data from O’Neill (1988)

−26834.7

6.471

−27215

6.57

−24935.0

8.489

−25096.3

8.735

−24441.9

8.290

−25035

8.74

> 914

−25738

9.00

600-800

Wones and Gilbert (1969)

−24525.4

8.944

427-1065

Regression of data from O’Neill and Pownceby (1993)

−24569.5

8.960

527-1147

Regression of data from O’Neill (1987a)

−25025

9.46

771-1204

Schwab and Küstner (1981)

−24930

9.36

0.046

519-1319

Huebner and Sato (1970) with “C” term from Chou (1987)

14.55

0.019

682-1100

Frost (1991)

−25700.6

0.056

0.110

800-1260

Myers and Eugster (1983)

1050-1400

Eugster and Wones (1962) using the data of Darken and Gurry (1945)

627-1147

Regression of data from O’Neill (1987b)

573-1200

Frost (1991)

600-1140

Myers and Eugster (1983) Schwab and Küstner (1981)

−23847.6

13.48

1040-1270

−25632

14.62

25-1227

Haas and Robie (1973)

−24912

14.41

800-1500

Eugster and Wones (1962) using the data of Norton (1955)

0.019

Myers and Eugster (1983)

*where log fO2 = A/T + B + C(P−1)/T; T is in K and P is in bars Note: T range is the range of temperatures used in data calibration

The relationship of oxygen fugacity to iron oxidation state through mineral equilibria can only be determined under the assumption of equilibrium. The quantification of this relationship is dependent on our knowledge of the stability of mineral assemblages at pressures and temperatures appropriate for the Earth’s mantle. This knowledge is derived from studies of mantle xenoliths and complementary experimental data, as well as high-pressure experiments for deeper parts of the mantle, from which no direct samples exist. A brief overview of the oxidation state and oxygen fugacity of the lower and upper parts of the Earth’s mantle is provided here; for a more complete review, see McCammon (2005).

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The lower mantle High-pressure experiments indicate that the lower mantle is dominated by (Mg,Fe)(Si,Al)O3 perovskite, with minor ferropericlase (Mg,Fe)O and CaSiO3 perovskite (e.g., Kesson et al. 1998). Recent experimental work on (Mg,Fe2+)SiO3 perovskite has demonstrated that the substitution of Al in the structure has significant effects on the iron oxidation state in perovskite (McCammon 1997). The coupled substitution Mg2+ + Si4+ = Fe3+ + Al3+ is energetically favorable, and causes an increase in the Fe3+ content of perovskite while maintaining charge balance in the structure. The Al content is correlated with Fe3+/ΣFe, even at low oxygen fugacity (Lauterbach et al. 2000). Thus, the Al-Fe3+ substitution in Mg perovskite is an example of a crystal-chemical control on the iron oxidation state recorded by a mineral, independent of oxygen fugacity. The implication of these observations is that the lower mantle has a higher Fe3+ content than previously assumed. Frost et al. (2004) calculate that the Fe3+/ΣFe of the lower mantle is 0.60, on the basis of their experimental results in conjunction with a bulk silicate earth composition in which the lower mantle contains 70 wt% perovskite containing approximately 5 wt% Al2O3. This implies that the lower mantle is enriched in Fe3+ relative to the upper mantle, or that the formation of Al-substituted perovskite is accompanied by a complementary reaction to maintain the same overall oxidation state as the upper mantle. The latter could involve the reduction of volatile (C-H-O) species, or the disproportionation of iron through the reaction: 3 Fe2+ = Fe0 + 2 Fe3+

(1)

In experiments on Al-substituted perovskite, reaction (1) is manifested in the presence of discrete blebs of iron metal (Lauterbach et al. 2000; Frost et al. 2004). The volatile budget of the mantle (Wood et al. 1996) is inadequate to account for the amount of reduction required (Frost et al. 2004). Evidence for whole-mantle convection (e.g., van der Hilst et al. 1997) precludes a scenario in which the lower mantle has a higher oxidation state than the upper mantle. The disproportionation reaction has other implications for the redox evolution of the Earth’s mantle. If the experimental results are representative, they imply that the lower mantle contains metal blebs, which are dispersed throughout the silicate assemblage; the recombination of iron metal and ferric iron would occur as material moves out of the perovskite stability field, resulting in no net change in bulk oxygen content (Frost et al. 2004). In the lower mantle of the early Earth, however, the metal formed by disproportionation may have been transported to the core during core formation, resulting in a net increase in O relative to Fe in the mantle (Wood and Halliday 2005). Assuming whole-mantle convection, this process may explain why the upper mantle is apparently out of equilibrium with an iron-rich core. Furthermore, the formation of iron blebs in the lower mantle would likely have implications for the abundances of siderophile elements in the Earth’s mantle (Frost et al. 2004). Although ferric iron may be the dominant form of iron in the lower mantle, its bulk abundance does not constrain the oxygen fugacity of the lower mantle. Ultimately, the problem with determining the oxygen fugacity of the lower mantle is the lack of appropriate mineral equilibria. Unlike the upper mantle, the temperature dependency of the lower mantle mineral assemblage is poorly known. At relevant pressures and temperatures, experimental studies show that Ca-perovskite can contain Fe3+, although natural samples are chemically very pure (Harte et al. 1999; Stachel et al. 2000). Ferropericlase is the only phase in the mantle in which Fe3+/ΣFe reflects oxygen fugacity (McCammon et al. 1998; Frost et al. 2001). McCammon et al. (2004) used the composition and Fe3+/ΣFe of ferropericlase in diamond inclusions to estimate an oxygen fugacity for the lower mantle between the IW and Re-ReO2 (e.g., Pownceby and O’Neill 1994) buffers. However, the estimate is largely qualitative and poorly constrained. Other, indirect approaches show promise, such as geophysical measurements of the lower mantle calibrated using experimental results (e.g., Wood and Nell 1991).

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The upper mantle The upper mantle is dominated by olivine, orthopyroxene, clinopyroxene, and spinel or garnet. All of these phases are iron-bearing; mineral equilibria involving them can be used to calculate oxygen fugacity in the spinel (lower pressure) and garnet (higher pressure) facies of the upper mantle. In the spinel facies, at depths less than ~60 km (~2 GPa), the dominant mineral equilibrium is 6 Fe2SiO4 + O2 = 2 Fe2+Fe3+2O4 + 6 FeSiO3 olivine

spinel

(2)

opx

which is commonly referred to as the spinel peridotite reaction. Calculation of oxygen fugacity in spinel peridotites, as outlined in a later section, shows a range of over 4 orders of magnitude, and a relationship to tectonic environment, metasomatism, and partial melting (e.g., Mattioli et al. 1989; Ballhaus et al. 1990; Bryndzia and Wood 1990; Wood et al. 1990; Ballhaus 1993; Kadik 1997; Parkinson and Arculus 1999; McCammon et al. 2001). Generally, the more reduced samples have oxygen fugacities between 2 log units below the fayalite-magnetite-quartz buffer and the buffer (FMQ − 2 to FMQ), and are derived from suboceanic abyssal peridotites and undepleted, fertile subcontinental mantle xenoliths, whereas the more oxidized samples, up to ~ FMQ + 2, are peridotites that have been influenced by the effects of subduction and metasomatism. Summaries are provided by Wood et al. (1990), Wood (1991), Ballhaus et al. (1990), Ballhaus (1993), Ionov and Wood (1992), Woodland et al. (1992) and Amundsen and Neumann (1992). The effect of pressure on equilibrium (2) is to drive the reaction to the right, favoring smaller-volume phases and resulting in lower oxygen fugacity. The volume change for the solids in equilibrium (2) is about half that of the solids in the FMQ buffer (8.6 cm3 vs. 17.95 cm3; Wood et al. 1996). Therefore, all else being equal, the oxygen fugacity will decrease by about 0.25 log units per GPa pressure increase relative to the FMQ buffer (Ballhaus 1995; Wood et al. 1996). At greater depths, garnet becomes stable and the mineral equilibrium that dominates is 4 Fe2SiO4 + 2 FeSiO3 + O2 = 2 Fe2+3Fe3+2Si3O12 olivine

opx

(3)

garnet

Calibration of this equilibrium for calculation of oxygen fugacity is provided by Gudmundsson and Wood (1995). Woodland and Koch (2003) point out that an erroneous expression for skiagite (Fe2+3Fe3+2Si3O12) activity was given by Gudmundsson and Wood (1995), and refer to Woodland and Peltonen (1999) for the correct expression. Application to garnet peridotite xenoliths from the Kaapvaal (Southern Africa) and Slave (Canada) Cratons yields oxygen fugacities of FMQ − 3 or lower, decreasing to below FMQ − 4 at about 6 GPa (Gudmundsson and Wood 1995; Woodland and Koch 2003; McCammon and Kopylova 2004). The decrease in oxygen fugacity with depth observed in garnet peridotite xenoliths is consistent with thermodynamic arguments, specifically volume effects. The volume change of reaction (3) is greater than that for reaction (2), and increasing pressure favors the incorporation of Fe3+ into garnet. The expected decrease in oxygen fugacity is 0.9 log units/GPa (Wood et al. 1996). The volume effects on the oxygen fugacity-depth relationships within the garnet and spinel facies are examples of crystal chemical controls on iron oxidation state. The lower oxygen fugacity of the garnet facies relative to the spinel facies is attributable to the effect of the relative modal abundances of Fe3+-bearing mineral phases—whereas the Fe3+ content (and therefore the O content) of the upper mantle is quite low, Fe3+/ΣFe ~ 0.023±0.010 (O’Neill et al. 1993), the relative oxygen fugacity of the spinel facies is quite high (near FMQ) because the Fe3+ is concentrated into spinel, the least abundant phase, and virtually excluded from

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all other phases. In fact, the oxygen fugacity of the spinel facies would be 4 log units lower if Fe3+ were concentrated equally in all mineral phases (O’Neill et al. 1993). Assuming the same bulk chemistry for the garnet, including Fe3+/ΣFe (O’Neill et al. 1993), the higher modal abundance of garnet at these pressures will dilute the concentration of Fe3+ in the garnet, lowering the activity of the skiagite component and resulting in a lower relative oxygen fugacity. Furthermore, the increase in the modal abundance of garnet with depth contributes to a further decrease in relative oxygen fugacity with increasing pressure (Wood et al. 1990; Ballhaus 1995). Deeper into the upper mantle, the modal abundances of garnet and clinopyroxene increase at the expense of orthopyroxene. Oxygen fugacity may therefore be controlled by 4 Fe2SiO4 + 2 FeSiO3 + O2 = 2 Fe2+3Fe3+2Si3O12 olivine

cpx

(4)

garnet

where FeSiO3 is the clinoferrosilite component in clinopyroxene (McCammon 2005). The breakdown of pyroxene to majorite garnet dilutes the Fe3+ in garnet, further lowering the relative oxygen fugacity (Wood et al. 1996). This may be offset by the difference in volume change in equilibrium (4) relative to equilibrium (3). The decrease in oxygen fugacity relative to FMQ may therefore be muted at these depths (McCammon 2005). Uncertainties in the thermodynamic properties of the components currently prevent an accurate assessment. Oxybarometry of the lower part of the upper mantle is important for mapping the overall redox stratigraphy of the mantle, however, and for addressing specific redox-dependent questions, such as where the oxygen fugacity of the bottom of the upper mantle lies relative to the Ni precipitation curve (e.g., Ballhaus 1995). Carbon is considered to be another important element in the Earth’s interior that is involved in redox-dependent equilibria. Fluids in the Earth’s mantle are dominated by species involving C, H and O in oxidized (e.g., CO2, H2O) and reduced (e.g., CH4, CO) forms, in equilibrium with C (as graphite in the upper mantle). The fluid species may also be dissolved in a melt. Two such equilibria are CH4 + O2 = C + 2 H2O

(5)

C + O2 = CO2

(6)

As written, reaction (5) involves the oxidation of methane and reaction (6) the oxidation of C; reaction (6) is often referred to as the CCO buffer. Because these equilibria involve fluids and/ or melt and graphite instead of mostly solid phases, the effect of pressure is different than that for equilibria (2), (3) and (4). This is simply due to volume effects—solid phases have smaller volumes than fluid phases. As a result, equilibria (5) and (6) have opposite slopes on plots of oxygen fugacity (relative to FMQ) as a function of increasing pressure. The implication is that mantle material that is buffered by Fe-bearing equilibria will have associated fluids that gradually shift with depth from H2O–CO2 to H2O–CH4, without any change in the activity of the ferric iron components. Opinions diverge as to the relative importance of Fe-bearing mineral equilibria and CH-O equilibria in controlling the oxygen fugacity of the upper mantle. Wood et al. (1996) argue that volatile speciation in the upper mantle depends on the oxygen fugacity determined by Fe-bearing mineral equilibria, because the concentration of C (80 ppm) is much less than that of iron (FeO ~ 8 wt%, Fe2O3 ~ 0.2 wt%). Ballhaus (1995) estimates the extent of relative oxidation with depth, taking into account all the factors that would influence this property, including those that would force a decrease, such as: volume effects (−0.3 to −0.4 log units/ GPa, mainly for equilibrium (2)); solid solution (−0.1 log units/GPa); and the spinel-to-garnet transition (−1.5 to −2 log units/GPa). However, he further argues that volatile equilibria (C-HO and S) are a moderating influence on the oxygen fugacity-depth gradient, providing enough

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oxygen for one-third to one-half of the Fe2O3 in the upper mantle. The average gradient would therefore be about −0.6 log units/GPa. Thus, this perspective has volatiles playing a more important role, relative to ferrous-ferric equilibria, in influencing the oxygen fugacity of the upper mantle.

OXYBAROMETERS APPLICABLE TO BASALTIC ROCKS Basaltic volcanism is a fundamental process on differentiated planetary bodies; samples from the Moon, Mars, and differentiated asteroids (e.g., 4 Vesta) are dominated by basaltic samples. Like other physical and chemical factors involved in its petrogenesis, the oxygen fugacity of a basalt is the result of the complex history of partial melting, extraction, ascension, eruption and emplacement. Basalts can be used as probes of planetary interiors, provided that the effects of post-extraction processes can be adequately assessed. The effort is worthwhile— ultimately, insights into fundamental differences in the origin and evolution of the terrestrial planets can be gained, as exemplified by the comparative studies of the Basaltic Volcanism Study Project (BVSP 1981). The suite of planetary samples has expanded significantly since the completion of the BVSP, resulting in further insights (e.g., Wadhwa 2008). Igneous petrologists, it often seems, would be happy if magmas never crystallized, and instead quenched to glasses that are representative of the parent magmas from which the physical and chemical factors involved in their petrogeneses can be directly determined. The relationship between redox state and the concentrations of FeO and Fe2O3 in a multicomponent silicate liquid, i.e., 2 Fe2+O + ½ O2 = Fe3+2O3

(7)

has been calibrated experimentally over a wide range of oxygen fugacity and bulk composition (Sack et al. 1980; Kilinc et al. 1983; Kress and Carmichael 1988, 1991) and is given by the empirical equation ln[XFe2O3/XFeO] = a ln fO2 + b/T + c + Σ Xi di

(8)

where a, b, c and di are constants determined by regression of experimental data; details are provided by Carmichael and Ghiorso (1990). Equation (8) allows calculation of oxygen fugacity from determinations of FeO and Fe2O3 in glassy lavas. More importantly, it has been demonstrated that the change in oxygen fugacity with temperature of a silicate liquid is parallel to the change in oxygen fugacity of a solid oxygen buffer such as NNO or FMQ. Otherwise said, the relative oxygen fugacity of a silicate liquid is independent of temperature. The implication is that in the absence of crystallization, determination of the redox state of a magma is straightforward. Furthermore, the redox states of different magmas can be easily compared by calculating the oxygen fugacity relative to a solid oxygen buffer at an arbitrary temperature. This has been done for glassy basic lavas from a range of tectonic environments on the Earth, demonstrating that oxygen fugacity varies more widely than any other variable in petrology, by over 7 orders of magnitude (Carmichael and Ghiorso 1990; Carmichael 1991). With a couple of rare exceptions, glassy lavas are absent from the planetary sample suite and we must rely on oxybarometers to see through the effects of crystallization. The ultimate goal of an oxybarometer is to employ a redox-sensitive element, or suite of elements, to provide a record of oxygen fugacity at a particular stage in a sample’s petrogenesis. Given that the stage of interest is typically prior to crystallization, in order to assess the redox conditions of the parent magma and infer the nature of its source, a couple of caveats are worth noting. Oxybarometers are thermodynamic models, and as such, the inherent assumptions outlined by Ghiorso (1997) regarding the application of thermodynamic models to igneous systems are relevant: that the igneous system is in equilibrium everywhere along its evolutionary path;

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and that the processes being modeled are reversible. The example given by Ghiorso (1997) is one in which a magmatic system evolves from melting in the source region (state A) to final solidification in a shallow reservoir (state B); as noted, there is an infinite number of reversible paths that will take the system from state A to state B. We could equally imagine the evolution of the redox state of a magma where state X is pooling in a shallow magma body, and state Y is 20% crystallization of an assemblage of olivine, pyroxene and chromite phenocrysts. In the application of an oxybarometer, one would do well to remember the following: “Utilizing thermodynamic models of igneous processes, one cannot ever hope to attain a unique inversion to provide the unique history of magma evolution. That history is lost in the assumption of the applicability of the method.” (Ghiorso 1997) Instead of trying to uniquely determine the history of a magma, thermodynamic models ought to be used in forward modeling of magmatic evolution in order to assist in discrimination between competing hypotheses (Ghiorso 1997). Another potential pitfall worth noting is that the oxygen fugacity recorded by an oxybarometer depends on how readily the redox-sensitive elements can be reset by subsequent processes. Using ferrous-ferric iron equilibria as an example, it does not require much oxygen to affect the ferric iron in a system with initial Fe3+/ΣFe = 0.10. The effects of subsolidus reequilibration are a concern for any oxybarometer, as eloquently summarized by the following poem by Cin-Ty A. Lee (loosely in the haiku style): Fugacity has no memory It has no past Only what it sees last The memory of an oxybarometer has been compared to those of elephants and goldfish by John Delano: whereas elephants remember paths for yearly migration and the locations of burial grounds, goldfish cannot remember what happened in the previous few seconds such that every trip around the fishbowl is a new one. Whether a particular oxybarometer is an “elephant” or a “goldfish” will depend equally on the geochemical behavior of the redox-sensitive element involved and on the petrogenesis of the rock to which it is applied. Textural or compositional evidence for equilibrium among mineral phases will bolster application of the oxybarometer, and results need to be assessed in the context of petrologic studies of the sample. Forward modeling of changes in oxygen fugacity with crystallization are useful in assessing, for example, the effect of crystallization on the redox state of the melt. Ghiorso (1997) models the equilibrium and fractional crystallization of primitive MORB under closed system conditions using the MELTS program (Ghiorso and Sack 1995). In contrast to experiments in which oxygen fugacity is fixed relative to a buffer, MELTS allows modeling of a system in which the total oxygen content of the system (liquid + solids) is constant. The results show the expected trend of increasing ferric iron in the melt resulting in an increase in the relative oxygen fugacity of the melt due to crystallization of Fe2+-bearing olivine and pyroxene; in this example, the increase is a maximum (under equilibrium crystallization) of 0.7 log units. Subsequent crystallization of (Fe3+-bearing) spinel results in a decrease in relative oxygen fugacity of 1 log unit; the total excursion throughout the crystallization history is a maximum of 0.8 log units (Ghiorso 1997). This is consistent with Carmichael and Ghiorso (1990), who argue that in a closed system, the iron redox state of the liquid during crystallization will be regulated such that it may resemble a buffered path; specifically, that the increase in ferric iron in the melt as a result of crystallization of an early, Fe3+-poor phase will stabilize a later, Fe3+-rich phase such as spinel, and an immiscible (Fe-S-O) sulfide liquid, both of which will counteract the increase in relative oxygen fugacity of the melt. In a crystallizing liquid under open-system conditions, any addition or subtraction of oxygen should result (initially) in a

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change in the proportion of the solids; any change in oxygen fugacity should result in a change in the iron redox ratio in the solids and liquid (Carmichael and Ghiorso 1990). These insights represent a framework in which oxybarometry results can be interpreted; for example, although Fe-Ti oxides (titanomagnetite and ilmenite) typically appear later in the crystallization of a basalt, the oxygen fugacity that they record may be reflective (i.e., within a log unit) of magmatic redox conditions, if the system was closed with respect to oxygen. One approach to determination of the redox state of a basaltic sample is to apply an oxybarometer, or multiple oxybarometers, to a range of assemblages in the rock, to determine the fO2-T path of the rock as best as possible. The advantage of this approach is twofold: changes in oxygen fugacity during crystallization can be assessed, resulting in insights into the petrogenesis of the rock (e.g., open- vs. closed-system); and the results from the highesttemperature, presumably near-liquidus assemblages can be more confidently interpreted as representing the redox conditions of the magma. Several methods (oxybarometers) currently exist to determine oxygen fugacity in basalts, based on ferrous-ferric mineral equilibria, or multivalent trace elements (e.g., Eu, V). The user faces two challenges: choosing the appropriate method; and assessing the relevance of the results within the petrologic context. To assist the reader in determining the best method for his or her particular needs and applying it in an informed manner, a description of each method is provided below, and their respective strengths, assumptions and limitations are outlined.

Oxygen fugacity from mineral equilibria Fe-Ti oxide. The Fe-Ti oxide oxybarometer owes its existence to A.F. Buddington, who suggested that the TiO2 content of magnetite was largely a function of temperature, on the basis of over 200 analyses of magnetite compiled from a range of igneous rock types (Buddington et al. 1955; Buddington 1956), and to J. Verhoogen, who demonstrated, on theoretical grounds, that the compositions of Fe-Ti oxides are significantly affected by the partial pressure of oxygen (Verhoogen 1962). Buddington and Lindsley (1964) overcame the lack of experimental data on the compositions of oxides as a function of T and oxygen fugacity, and extrapolated their data to experimentally inaccessible but petrologically important conditions, enabling the first practical application of the oxybarometer, albeit limited to the Fe-Ti binary system (i.e., oxide pairs containing exclusively the Fe and Ti cations). More information on the development of this oxybarometer can be found in Ghiorso and Sack (1991a), Lindsley and Frost (1992), and Lattard et al. (2005). The strength of the Fe-Ti oxide oxybarometer rests in the nature of the two oxides involved; the cubic oxide in the magnetite (Fe2+Fe3+2O4) - ulvöspinel (Fe2+2TiO4) series and the rhombohedral oxide in the hematite (Fe3+2O3) - ilmenite (Fe2+TiO3) series are each solid solutions of end-members with different oxidation states. As such, the activities of magnetite in the cubic oxide and hematite in the rhombohedral oxide are used as the oxybarometer according to the reaction 4 Fe2+Fe3+2O4 + O2 = 6 Fe3+2O3

(9)

Note that this is equivalent to the Magnetite-Hematite (MH) solid oxygen buffer. Exchange of Fe and Ti between the oxide pairs is strongly dependent on temperature and only weakly pressure-dependent. This exchange is expressed in the following equilibrium Fe3+2O3 + Fe2+2TiO4 = Fe2+TiO3 + Fe2+Fe3+2O4

(10)

The main weakness of this oxybarometer is that Fe and Ti exchange between pairs continues during cooling. As such, the oxides do not retain their high-temperature compositions and are prone to resetting. Furthermore, in many basaltic rocks, they are among the last phases to crystallize. Thus, the oxides may record temperatures and oxygen fugacities that differ from

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those of the original magma. Fortunately, subsolidus equilibration between the oxides behaves in a somewhat predictable manner, and may be accounted for in many cases, as outlined by Lindsley and Frost (1992). Currently there exist two commonly-used formulations of the Fe-Ti oxide oxybarometer. Ghiorso and Sack (1991a) noted that the application of Buddington and Lindsley’s formulation to multicomponent oxides required algorithms to project the compositions found in nature into the relevant binary systems. Citing a need for a more robust thermodynamic treatment of the Fe-Ti oxides that includes the main substituting cations, Ghiorso and Sack (1991a) presented a thermodynamic model that accounts simultaneously for all of the complex peculiarities of the Fe-Ti oxides (i.e., phase equilibrium constraints, cation order-disorder, and mixing and end-member properties), while minimizing projection schemes. This is the strength of the formulation, because proportions of some substituting cations, such as Al and Cr in cubic oxides, can be significant. The formulation is available as a software package from the MSA website (Supplemental Material at http://www.minsocam.org), and is simple to use, requiring standard wt %-oxide compositional data for the two oxides. A new version of the oxybarometer, which is much improved for conditions of fO2 = NNO to NNO + 3 and T = 700 to 900 °C is expected in 2008 (Ghiorso, pers.comm.). The formulation is based on the quinary model for cubic oxides, using solution theory adopted from Sack and Ghiorso (1991a, 1991b) for cubic oxides in the system (Mg,Fe2+)(Al,Cr,Fe3+)2O4 – (Mg,Fe2+)2TiO4. Minor elements, specifically V, Mn, Ca, Zn and Ni, are included in the formulation; however, their inclusion is based on several assumptions. For example Mn, Zn and Ni are modeled as MnAl2O4, ZnAl2O4 and NiAl2O4 components, respectively, and are proxied by an equivalent amount of FeAl2O4. Non-zero concentrations of V2O3, MnO, CaO and ZnO are not rigorously accounted for in the calculation of equilibration temperatures; the authors note that the concentrations of these components in their sample dataset are less than 0.5 wt%, with the exception of MnO (up to 2.8 wt%). Therefore the user should be wary of using this oxybarometer when the concentrations of these components exceed the concentrations used in the oxybarometer formulation. The energetics of rhombohedral oxides are modeled in the quaternary system Fe2O3 – FeTiO3 – MgTiO3 – MnTiO3. Minor elements include Al, V, Cr, Ca, Zn and Ni. Non-zero concentrations of Al2O3, V2O3, Cr2O3, CaO and ZnO are not rigorously accounted for, but Al2O3 and V2O3 are proxied as hematite, and the concentrations of each of these components in the authors’ dataset is less than 0.2 wt%. Once again, the utility of the oxybarometer may be limited if the concentrations of these components exceed those used in the formulation. Since the calibration of the model does not account for the energetics of magnetic ordering in either oxide, it cannot be used for assemblages that equilibrated below 600 °C. Non-stoichiometry effects, which would be most significant above 1100 °C and at oxygen fugacity near the limits of cubic oxide stability, are excluded from the model; caution should be used in applying the formulation to assemblages suspected of equilibration under such conditions. The Buddington and Lindsley work enabled the estimation of oxygen fugacity in various oxide-bearing igneous rocks. Consequently, it was recognized that the oxygen fugacity was not only reflected in the compositions of Fe-Ti oxides, but also in the compositions of coexisting ferromagnesian silicates (Frost et al. 1988) and that much of the common subsolidus reequilibration of the Fe-Ti oxides could be circumvented through the use of equilibria such as SiO2 + 2 Fe2TiO4 = 2 FeTiO3 + Fe2SiO4 quartz

ulvöspinel

ilmenite

(11)

fayalite

commonly referred to as QUIlF (Frost et al. 1998). Lindsley and Frost (1992) and Andersen et al. (1993) presented an updated thermodynamic model for the Mg- and Ca-bearing system, referred to as Ca-QUIlF. This formulation includes

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equilibria between augite, pigeonite, orthopyroxene, olivine and quartz in addition to the Fe-Ti oxides. The main advantage of this formulation is that it can reduce the uncertainty in using the Fe-Ti oxide oxybarometer alone. For example, the four-component subsystem FeO-MgO-Fe2O3-TiO2, with two phases (the Fe-Ti oxides) has a formal variance of four, but because the partitioning of Fe and Ti are coupled, and the effect of Mg on temperature and oxygen fugacity is minor, two intensive parameters (T and fO2) are tightly constrained. Furthermore, the equilibria can be used to assess equilibrium among phases. For example, if the temperature calculated from Fe-Ti oxides is consistent with a temperature calculated from the same oxides with co-existing olivine and pyroxene, then it supports equilibrium amongst all of these phases. By the same token, the model can be used to “see through” subsolidus re-equilibration of oxides, including oxyexsolution of the cubic oxide (i.e., titanomagnetite). For example, if pyroxene temperatures and oxide temperatures do not agree (assuming these phases were initially in equilibrium) then the original oxygen fugacity can be estimated, given the relative abundances of the oxides and assuming a closed system during cooling. In fact, the rhombohedral oxide is expected to gain FeTiO3 and the cubic oxide is expected to gain Fe3O4 upon equilibrium cooling in a closed system (Lindsley and Frost 1992). Thus, magmatic (or at least super-solidus) oxygen fugacity estimates can be made, even when the Fe-Ti oxides have undergone re-equilibration. Ca-QUIlF is available as a software package (QUIlF95) from the MSA website (Supplemental Material at http://www.minsocam.org). Mineral compositions must be expressed in terms of mole fractions of end-members (e.g., XHematite, XFayalite, XWollastonite) with the exception of the cubic oxide phase, for which the user must calculate the numbers of Ti, Mg and Mn cations per formula unit (NTi, NMg, NMn; cations per four oxygen). The program can be used to calculate temperature, pressure, oxygen fugacity, equilibrium mineral compositions, and activities of SiO2, Fe and TiO2. As in the previous case, some limitations are useful to note. The calibration of the oxide models was done below FMQ + 2, and therefore may not be applicable to highly oxidized assemblages. With regard to cubic oxide compositions, Fe, Ti and Mg are the only cations accounted for, and the assumptions of the model provide for only two independent compositional parameters, NTi and NMg. Therefore the model is not appropriate for cubic oxides with significant Al2O3 and Cr2O3 contents. The formulation assumes that silicates have negligible Fe2O3 and TiO2 contents and that CaO is unimportant in the oxides; therefore, caution should be used where, for example, pyroxene contains significant TiO2. Lastly, the model cannot be used quantitatively for MgFe3+2O4-rich cubic oxides because of the limits of the solution models and because such oxides will tend to be nonstoichiometric; as with the Ghiorso-Sack model, nonstoichiometry is not included in the model. The reader is referred to Ghiorso and Sack (1991a,b), for comparisons of the GhiorsoSack formulation with Ca-QUIlF. It should be noted that the Ghiorso-Sack model uses thermodynamic data that are internally consistent with olivine and orthopyroxene solution theory from Sack and Ghiorso (1989) as well as the fayalite, ferrosilite, O2 (gas), and quartz data after Berman (1988); these data could be assembled in a separate formulation of QUIlF. The most obvious difference between the Ca-QUIlF and Ghiorso-Sack formulations of the FeTi oxide oxybarometer is the treatment of minor substituting cations, especially in the cubic oxides. In cases where the concentrations of minor cations are low, the two models agree well (e.g., Herd et al. 2001). Citing the lack of experimental calibration of the Fe-Ti oxide oxybarometer at high temperatures and over a wide range of oxygen fugacity, Lattard et al. (2005) carried out a series of experiments in the Fe-Ti-O system at temperatures of 1000 to 1300 °C and oxygen fugacity ranging from NNO – 5 to NNO + 5 (~IW to FMQ + 6). The results should lead to an improved thermodynamic model for rhombohedral oxide, thereby reducing the discrepancies

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between experimental results and those calculated using either the Ca-QUIlF or Ghiorso-Sack models, as well as enabling more accurate calculation of oxygen fugacity, especially under oxidizing conditions (> FMQ + 2). The data of Lattard et al. (2005) also include compositions of ilmenite and pseudobrookite ((Fe3+,Fe2+)2(Ti,Fe3+)O5), which is an additional oxybarometer applicable to some terrestrial and lunar igneous rocks. In deciding which oxybarometer to apply, the user should exercise caution in applying a particular formulation to oxide compositions that deviate significantly from the Fe-Ti endmembers upon which the thermodynamic model is based. The best application of any formulation is informed by detailed petrography and an assessment of the degree of equilibrium. Olivine-pyroxene-spinel. The olivine-pyroxene-spinel oxybarometer, also referred to as the spinel peridotite oxybarometer, was developed for application to mantle xenoliths from the spinel facies. It is governed by equilibrium (2), repeated here: 6 Fe2SiO4 + O2 = 2 Fe2+Fe3+2O4 + 6 FeSiO3 olivine

spinel

(2)

opx

This equilibrium has also been referred to as the fayalite-ferrosilite-magnetite (FFM) buffer (e.g., King et al. 2000). Oxygen fugacity is calculated using log (fO2) = − 6 log aolFe2SiO4 + 2 log aspFe3O4+ 6 log aopxFeSiO3

(2a)

involving the activities (a) of the respective iron end-members. In principle, this equilibrium is applicable to low-pressure assemblages in basaltic samples, assuming equilibrium between olivine, spinel, and low-Ca pyroxene. Spinel in this case is typically chromian spinel. Because the mineral phases in Equation (2) are near the liquidi of many basaltic rocks, the results from this oxybarometer are presumably good indicators of magmatic temperatures and oxygen fugacities. Subsolidus re-equilibration is limited to Fe-Mg exchange. The involvement of three phases requires determination of whether all three are cogenetic and remained in equilibrium. Detailed petrography can assist in overcoming this obstacle. An overview of the development of the olivine-pyroxene-spinel oxybarometer is given by Wood (1991), where he presents the equation for oxygen fugacity relative to the FMQ buffer, as originally derived by Wood (1990): log (fO2)P,T = log fO2 (FMQ)P,T + 220/T + 0.35 − 0.0369P/T − 12 log XolFe − (2620/T)(XolMg)2 + 3 log (XM1Fe ·XM2Fe)opx + 2 log aspFe3O4

(12)

where T is the temperature in K and P is the pressure in bars. Each part of this equation can be understood in terms of Equation (2a). The activity of Fe2SiO4 in olivine is modeled assuming random mixing over the two cation sites, and is represented by the term, “−12 log XolFe − (2620/ T)(XolMg)2”, in which XolFe and XolMg represent the mole fractions of fayalite and forsterite endmembers, respectively, in olivine. The activity of FeSiO3 in orthopyroxene is treated as an ideal two-site solution, represented by the term, “3 log (XM1Fe ·XM2Fe)opx”, in which XM1Fe and XM2Fe represent the atomic fractions of Fe on M1 and M2, respectively. These are calculated as described by Wood (1990): “Al was added to Si to fix tetrahedral occupancy at 2.0 per 6 oxygens. The remaining Al (VI), Cr and Ti were placed in M1, while Ca and Mn were placed in M2. Fe and Mg were then evenly distributed between M1 and M2 positions and atomic fraction of Fe on M1 (XM1Fe) and M2 calculated.” The last term involves the activity of the magnetite component in spinel (aspFe3O4); markedly, no expression is incorporated into the equation. The largest uncertainty in determining oxygen fugacity with this method derives from the uncertainty in aspFe3O4. For this reason there exist several equations for calculating aspFe3O4, outlined below. The remaining terms in Equation (12) reflect the FMQ buffer to which the results of the calculation are related. The term, “log fO2 (FMQ)P,T” is the oxygen fugacity of the FMQ buffer at the P and T of interest. Numerous formulations of the FMQ buffer exist (Table 1); therefore,

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it is especially important to note that the Myers and Eugster (1983) formulation was used in the derivation of Equation (12), which is: log (fO2) = −24441.9/T + 8.29

(13)

The term, “220/T + 0.35 − 0.0369P/T” represents the pressure and temperature dependency of the difference between the FMQ buffer (Eqn. 13; Myers and Eugster 1983) and “FFM” (Eqn. 2), after Mattioli and Wood (1988; their Eqn. 32). The derivation of the pressure term in Equation (12) is not explicit in Wood (1990) or Wood (1991). Significantly, oxygen fugacity calculated using Equation (12) is de facto relative to the FMQ buffer of Myers and Eugster (1983). Absolute log (fO2) can be calculated by also calculating log fO2 (FMQ) according to Equation (13). Temperature and pressure are required for the calculation. Pressures for planetary basalts are typically close to atmospheric (~1 bar). Temperature can be calculated using any of several formulations of the chromite-spinel geothermometer (e.g., Fabries 1979; Sack and Ghiorso 1991a). Wood (1991) estimates that oxygen fugacity can be calculated to within ± 0.5 log units using the olivine-pyroxene-spinel oxybarometer, assuming good analyses of all three phases are available. The largest contributor to the uncertainty is in the determination of aspFe3O4. Mattioli and Wood (1988) determined aspFe3O4 across the MgAl2O4-Fe3O4 join, between 900 and 1000 °C at 1 atm (1 bar) pressure. They did not, however, account for the effects of the chromite (FeCr2O4) component. Insights into order-disorder in spinel from models and electrochemical measurements by O’Neill and Wall (1987) and Nell and Wood (1991) provided updates of the model to account for Cr. The Nell-Wood equation for aspFe3O4 (see also Wood 1991) is applicable to XFe3O4 = 0.008 to 0.06 and T = 800 to 1400 °C. The quinary model for cubic oxides of Sack and Ghiorso (1991a,b) provides an alternative method of calculating aspFe3O4. This is accessible using the MELTS Supplemental Calculator (http://melts.ofm-research.org/CalcForms/index.html), and requires that the user first calculate mole fractions of chromite, hercynite, magnetite, spinel and ulvöspinel, which are then used to calculate thermodynamic properties at a chosen T and P. The Sack and Ghiorso (1991a,b) models are applicable to a range of spinel compositions, which is more desirable when using the olivine-pyroxene-spinel oxybarometer for the compositions typical of basaltic rocks. Ballhaus et al. (1991) provide an empirical calibration of the O’Neill and Wall (1987) olivine-pyroxene-spinel oxybarometer, using synthetic spinel harzburgite and lherzolite assemblages between 1040 and 1300 °C and 0.3 to 2.7 GPa. The advantage of the formulation is that it obviates the need for an explicit calculation of the activity of the magnetite component in spinel. However, the formulation is simplified by canceling orthopyroxene against the ideal part of the fayalite activity in olivine. This simplification cannot be expected to be valid at XFeol > 0.15. As such, its application is limited to Mg-rich upper mantle-derived rocks. Wood (1990) ran experiments equilibrating olivine, orthopyroxene and spinel at known oxygen fugacity (between FMQ and FMQ – 2) and temperature (1188 to 1205 °C), in order to test the Mattioli-Wood, O’Neill-Wall and Nell-Wood expressions for aspFe3O4 in the calculation of fO2 using Equation (12). He demonstrated that the Mattioli-Wood model was slightly dependent on Cr content, and that the Nell-Wood version more accurately reproduced the known fO2 compared to the O’Neill-Wall version. For comparison, the calculation was reproduced for this work by using the MELTS calculator for magnetite activity. It more closely reproduced the known fO2 than did the Nell-Wood version, as shown in Figure 1a. Whereas the Nell-Wood version underestimated the fO2 by 0.3 log units, the present calculation underestimates fO2 by only 0.07 log units, on average. The calculation was repeated using the Ballhaus et al. (1991) formulation (Fig. 1b); this model underestimates fO2 by an average of 0.45 log units.

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Figure 1. (a) Comparison of oxygen fugacity calculated according to Equation (12), using the MELTS Supplemental Calculator for the activity of magnetite in spinel, applied to the experimental data of Wood (1990). The average of the calculations is offset from the known oxygen fugacity by 0.07 log units. No dependence on the Cr/(Cr + Al) ratio is observed. (b) Comparison of oxygen fugacity calculated according to the expression of Ballhaus et al. (1991). The average of the calculations is offset from the known oxygen fugacity by 0.45 log units. No dependence on the Cr/(Cr + Al) ratio is observed.

Therefore, the use of the MELTS Supplemental Calculator to calculate the activity of magnetite in spinel, in combination with Equation (12), is as good as other formulations for spinel peridotite compositions. The diversity of solid solutions used in the MELTS Supplemental Calculator allows for wider applicability. As an example, this method has been applied to olivine-phyric martian basalts (Herd et al. 2002; Goodrich et al. 2003; Herd 2003, 2006). In many cases, the results agree with other oxybarometers applied to the same rock (e.g., Goodrich et al. 2003; Herd 2006). The olivine-pyroxene-spinel oxybarometer has seen little use for low-pressure assemblages in terrestrial basaltic samples. Ballhaus et al. (1991) used their model to calculate fO2 for midocean ridge basalts (MORB), island arc basalts (IAB) and ocean island basalts (OIB), but due to the inherent assumptions, the model is limited to mantle-derived primitive melts, and is not

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appropriate for more evolved basalts. The Wood – MELTS version of the oxybarometer is readily applicable to olivine-phyric terrestrial basalts. O’Neill and Wall (1987) used a slightly different approach to calculation of oxygen fugacity from olivine, pyroxene and spinel in mantle assemblages. Instead of “FFM”, they used FMQ, i.e., 3 Fe2+2SiO4 + O2 = 2 Fe2+Fe3+2O4 + 3 SiO2 olivine

spinel

(14)

quartz

Oxygen fugacity is then given by log fO2 = log fO2 (FMQ) − 3log aolFe2SiO4 + 3log aSiO2 + 2log aspFe3O4

(15)

The activity of SiO2 is calculated based on the equilibrium Mg2SiO4 + SiO2 = Mg2Si2O6 olivine

quartz

(16)

opx

and given by log aSiO2 = −350/T + 0.016 − 0.020P/T + log aopxMg2Si2O6 – log aolMg2SiO4

(17)

This method has not been tested for use with basaltic samples from planetary surfaces, although it is potentially applicable, with use of the MELTS Supplemental Calculator (or some other method) to calculate the activities of fayalite, forsterite, magnetite and enstatite.

Multivalent trace elements The previous oxybarometers are based on the equilibria between iron-bearing minerals in planetary basalts. In essence, they depend on the partitioning of ferrous and ferric iron among phases. Similarly, there exist a number of other multivalent elements whose valence states can be used to determine oxygen fugacity. Those that have a range of valence states under the redox conditions of planetary basalts include some of the transition elements (Ti, V, and Cr) and Eu, a particularly useful rare earth element (REE). The main difference between these elements and iron is that they are present in trace concentrations in major minerals. Their concentrations are such that they can be analyzed by in situ microbeam methods, including Electron Microprobe (EMP) and Secondary Ion Mass Spectrometry (SIMS). Determining their valence states, however, is a challenge, and methods have been developed, or are presently in development, to overcome this hurdle. The multivalent trace elements of interest for planetary basalts include Eu2+,3+, V , Cr2+, 3+, and Ti3+, 4+. These are shown schematically (along with iron) in Figure 2, after Papike et al. (2005). This diagram shows the range of valence states of the multivalent trace elements, for comparison with the range of fO2 of planetary basalts. Each point represents the oxygen fugacity at which the oxidized and reduced species are present in approximately equal proportions in a basaltic melt. The diagram is a useful “roadmap” for selecting appropriate oxybarometers—it is readily seen that mineral equilibria oxybarometers involving Fe2+ and Fe3+ are most applicable to the range of fO2 experienced by terrestrial and martian basalts. Of these elements, only vanadium exists in 4 valence states, covering the range of redox conditions of planetary basalts. For this reason, much recent effort has focused on the development of vanadium oxybarometers. 2+, 3+, 4+, 5+

Oxygen fugacity from multivalent trace elements Europium. Europium is the only REE whose geochemical behavior is significantly different from the rest of the REE in planetary magmas, due to its stability as Eu2+ or Eu3+ at fO2 ≤ FMQ. The “Eu anomaly” that is often observed in chondrite-normalized REE patterns in minerals and bulk rocks is due to this effect, coupled with crystal chemistry. The valence state

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Figure 2. Schematic diagram showing the dominant valence states of multivalent elements in planetary basalts, and their relationship to oxygen fugacity relative to the IW buffer, after Papike et al. (2005). See text for discussion.

of Eu in a silicate melt is related to oxygen fugacity by EuO (melt) + ¼O2 (g) = EuO1.5 (melt)

(18)

which is analogous to the relationship between ferrous and ferric iron and fO2. However, europium’s status as a trace element requires a different approach to for determination of the relationship of Eu behavior to oxygen fugacity, which involves the partitioning of Eu between minerals or between mineral and melt. Equilibrium (18) can be rearranged to solve for oxygen fugacity as log fO2 = − 4 log [aEuO(melt)/aEuO1.5(melt)] − 4 log K

(19) 2+

3+

which demonstrates that at constant T and P, the ratio of the activities of Eu and Eu is a function of oxygen fugacity. As soon as minerals become involved, crystal chemical effects, especially crystal/liquid partitioning, must be taken into account. Thus, it is expected that the Eu3+/Eu2+ ratio in a given mineral will be a function of the composition of the melt, the crystal chemistry of the mineral, and the oxygen fugacity at the time of crystallization. Recognizing the potential of Eu as an oxybarometer, J. A. Philpotts developed a method for calculating Eu2+ and Eu3+ concentrations in igneous phases (Philpotts 1970). His formulation addresses the fact that Eu2+ and Eu3+ cannot be directly measured. Instead, it uses partition coefficients, i.e., DEuX+β/α = EuX+β/EuX+α

(20)

X+

where Eu y is the concentration of Eu of some valence X in phase y. In each phase, the concentration of Eu can be expressed as Euα = Eu2+α + Eu3+α Euβ =

Eu2+β

+

Eu3+β

(21a) (21b)

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Assuming two phases in equilibrium, for example, where α is the matrix (as a proxy for the melt) and β is a plagioclase phenocryst, Equations (20), (21a) and (21b) can be combined to solve for the concentration of one of the valence states of Eu in the matrix (α), i.e., Eu3+α = [Eu β − DEu2+β/α· Euα]/[DEu3+β/α − DEu2+β/α]

(22a)

2+

with the concentration of Eu in the matrix determined by Eu2+α = Euα − Eu3+α

(22b)

The same could be repeated if other phenocrysts are in equilibrium with the matrix, to check for internal consistency (Philpotts 1970). Note that, in order to apply Equation (22a), the concentrations of Eu in the matrix and phenocryst (Euα and Euβ, respectively), and the partition coefficients for Eu2+ and Eu3+ for the phenocryst mineral must be known. Philpotts (1970) used analyses determined by mineral separation and stable isotope dilution mass spectrometry for the former (e.g., Schnetzler and Philpotts 1970); today, in situ microbeam techniques (e.g., SIMS, LA-ICPMS) would be used. The partition coefficients were estimated by Philpotts (1970), with DSr2+β/α used as a proxy for DEu2+ β/α, and DEu3+ β/α interpolated from a plot of DREE3+ β/α for the same phases. Since that time, experimental studies have addressed the partitioning of Eu in different phases under different redox conditions (e.g., Grutzeck et al. 1974; Sun et al. 1974; Drake 1975; Weill and McKay 1975; McCanta et al. 2004), and parameterized it using DEu/DGd or DEu/DSm, as described below. Philpotts (1970) applied Equations (22a) and (22b) to a suite of terrestrial and lunar samples, and to the eucrite meteorites Moore County and Juvinas (thought to be from the differentiated asteroid 4 Vesta; Drake 1979). He employed the Eu2+/Eu3+ ratios of the sample matrices to calculate oxygen fugacity according to Equation (19). In spite of certain variations in temperature, pressure and bulk composition among the samples, he determined that lunar basalt crystallized under redox conditions 4 to 5 log units below that of terrestrial basalts, and that the Juvinas eucrite oxygen fugacity is an additional 2 or 3 log units more reduced than lunar basalts. Drake (1975) applied his experimental Eu partitioning data for plagioclase to the Philpotts (1970) Eu2+/Eu3+ ratios and obtained fO2 results consistent with those of Philpotts (1970). These results are broadly consistent with observations of lunar basalts, which are essentially at metal saturation; however, similarities in mineral compositions and experimental petrology results for lunar and eucrite basalts suggest virtually identical oxygen fugacities (e.g., Kesson and Lindsley 1976; Stolper 1977; Longhi 1992). Regardless, subsequent studies of planetary basalts (e.g., BVSP 1981; Wadhwa 2008) have corroborated this range of oxygen fugacity among the terrestrial planets. The approach of McKay (1989) and McKay et al. (1994) is worth highlighting, since it uses the Eu oxybarometer to constrain the oxygen fugacity of a planetary basalt, specifically the LEW 86010 angrite meteorite, to within one log unit. Using appropriate experimental partitioning data for Eu among Al-, Ti-rich (fassaitic) pyroxene, anorthite and melt at 1175 to 1210 °C and atmospheric pressure, McKay (1989) determined the relationship between DEu/DGd and oxygen fugacity for plagioclase and pyroxene. At high oxygen fugacity (~ FMQ), Eu3+ is the dominant species and, having a smaller ionic radius ([VI]Eu3+ radius = 1.09 Å; Shannon 1976) than the divalent cation ([VI]Eu2+ radius = 1.31 Å; Shannon, 1976), it is more readily incorporated into the pyroxene M sites. Presumably the substitution of Eu3+ occurs via a coupled substitution with Na+ in the M2 site or Al3+ in the tetrahedral (T) site. At low fO2 (~ IW), Eu2+ is favored in the large feldspar site, as a substitution for Ca2+ or Na+ ([VIII]Eu2+ radius = 1.39 Å; [VIII]Eu3+ radius = 1.21 Å; Shannon 1976). The predicted relationship of DEu as a function of fO2 is an S-shaped curve that asymptotically approaches the values of DEu2+ at low fO2 and DEu3+ at high fO2. The observations follow the predicted relationship between DEu and fO2 on the basis of theory, as outlined by McKay et al. (1994). The advantage of the application of Eu in pyroxene and plagioclase for planetary basalts

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is that the greatest variation in DEu in each of these phases occurs over the range of fO2 between IW-2 and FMQ, which covers the majority of the range of oxygen fugacity of planetary basalts. The contrasting behavior of DEupl and DEupx, with the former increasing and the latter decreasing with fO2, provides an additional advantage: assuming that the concentrations of Eu and Gd in pyroxene and plagioclase are known, the two phases are in equilibrium, and the experimental phase compositions and temperatures are similar to those of the rock, then the two curves can be combined to give a calibration curve for the Eu oxybarometer. This is done by determining the ratio of (DEupl/DGdpl)/(DEupx/DGdpx); because these are mineral-melt partition coefficients, the melt concentrations cancel out, and only the Eu and Gd concentrations in the two phases are required. Thus, a plot of log (Eu/Gd)pl/(Eu/Gd)px vs. log fO2 will yield a linear relationship from which oxygen fugacity can be calculated (McKay et al. 1994). In the application of McKay et al. (1994), the Eu oxybarometer was used to determine that the LEW 86010 meteorite crystallized at an oxygen fugacity between IW and IW + 1. The primary advantage of the Eu oxybarometer is that the REE are relatively immobile elements (e.g., Van Orman et al. 2001) that are taken up by igneous phases and are present in concentrations that are measurable by available in situ methods. Thus, the magmatic Eu/ Gd (or Eu/Sm) ratio is likely to be preserved through subsolidus re-equilibration; assuming an appropriate calibration, the magmatic oxygen fugacity can be determined. As with other oxybarometers, the user should be aware of potential pitfalls. Ideally, the Eu oxybarometer is applied to samples in which pyroxene and plagioclase are in equilibrium on the liquidus. In reality, this does not often occur. The crystal chemistry of the REE in pyroxene and plagioclase is such that compositional effects on Eu partitioning may be significant. For example, the SiO2/Al2O3 activity ratio in the melt may influence Eu partitioning in plagioclase (Morris and Haskin 1974; Drake 1975). Likewise, the uptake of Eu3+ into pyroxene M sites requires a coupled substitution, and so is presumably influenced by the content of univalent cations (especially Na+); hence the need for calibration curves using appropriate melt compositions. Recently, a variation on the Eu oxybarometer has been developed for martian meteorites involving Eu in pyroxene (Wadhwa 2001; Musselwhite and Jones 2003; McCanta et al. 2004). Pyroxene, especially low-Ca pyroxene, is a liquidus phase in these basalts, with plagioclase crystallizing later. This work has provided an important means of estimating the magmatic oxygen fugacity in martian basalts (Wadhwa 2008). The fractionation of Eu into pyroxene becomes small at higher oxygen fugacity, however, where Eu3+ becomes dominant, resulting in greater uncertainty, up to ± 1 log unit at FMQ (McCanta et al. 2004). Therefore, this version of the oxybarometer has the greatest resolution at lower oxygen fugacity (< IW + 1). Vanadium. As shown in Figure 2, vanadium exists in four valence states. In theory, if the relative proportions of V species can be calibrated for oxygen fugacity, then it would provide an oxybarometer appropriate for all planetary basalts. This step has been accomplished for V in glass by Sutton et al. (2005) using vanadium K-edge X-ray Absorption Near-Edge Structure (XANES) spectroscopy, done with synchrotron light at the Advanced Photon Source (APS), Argonne National Laboratory in Argonne, IL. Vanadium K edge XANES spectra have a pronounced pre-edge feature whose intensity and energy increase systematically with valence state. Sutton et al. (2005) used a suite of synthetic basalt (forsterite-anorthite-silica and forsterite-anorthite-diopside) glasses in which vanadium valence state had been determined by titration. They then determined the relationship between the pre-edge peak intensity, I, and the effective vanadium valence, V*, I = − 153 + 199(V*) − 106(V*)2 + 22.4(V*)3

(23)

V * = 3f(V3+) + 4f(V4+) + 5f(V5+)

(24)

*

where V is given by and f is the fractional content of each species. It is assumed that V2+ has zero pre-edge peak

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intensity. Equation (23) was used to determine V* for 5 other suites of synthetic glasses; observations were consistent with predictions of the dominant species at different oxygen fugacity. The effects of temperature and melt structure are accounted for, but introduce little additional uncertainty (± 0.2 and 0.5 log units, respectively). A calibration curve of vanadium pre-edge peak intensity vs. log fO2 was constructed from the data from all synthetic glasses. Due to uncertainties at very low fO2, and the dominance of V5+ at the highest fO2 (and a concomitant reduction in resolution) the oxybarometer can be effectively used for glasses between IW − 2 to IW + 6 (corrected to a temperature of 1400 °C); a range that is unparalleled in oxybarometry. Application of the V XANES oxybarometer to natural glasses from the Earth, Moon and Mars yields results that are broadly consistent with previous studies, over a range from IW – 2 (lunar) to IW + 4 (terrestrial) with an uncertainty of ± 0.2 log units (Karner et al. 2006). The method is non-destructive and can be used on traditional polished thin sections at micron-scale resolution; furthermore, it is sensitive to V concentrations at the ~100 ppm level (Sutton et al. 2005). The main disadvantage is the lack of preserved glasses in natural basaltic samples. For this reason, the application of K edge XANES spectrometry to V in minerals is under development. This application faces the dual challenges of orientation effects and crystal chemical controls on V species partitioning. Papike et al. (2005) developed a semi-quantitative oxybarometer for V in chromite. The basis of this oxybarometer is the crystal chemistry of spinel, which has an affinity for V3+ over V4+. Canil (2002) recognized the dependence of DVsp on oxygen fugacity, noting that DVsp in high Cr/Al spinels (chromite) decreases by about an order of magnitude (from 32 to 5) between IW – 1 and IW + 4. Papike et al. (2005) note that V behaves differently in spinel from different planetary basalts, as evidenced by core-to-rim traverses across grains using the EMP. As illustrated in Figure 2, V4+ dominates under terrestrial redox conditions, V3+ dominates under lunar redox conditions, and martian basalts have subequal proportions of both. This is reflected in the core-to-rim patterns in spinel, which show compositional zonation from chromite cores to ulvöspinel rims: in lunar basalts, V as predominantly V3+ follows Cr (as Cr3+), decreasing from core to rim; in terrestrial basalts, V as predominantly V4+ follows Ti (as Ti4+), increasing from core to rim; in martian basalts, the trends vary, consistent with differences in oxygen fugacity between different samples (Papike et al. 2005). A comparison of the V contents of chromite cores among planetary basalts also reflects differences in oxygen fugacity; a plot of 100V/(Cr+Al) atomic with distance in chromite cores from lunar, martian and terrestrial basalts qualitatively differentiates the ranges of oxygen fugacity for the Moon, Mars and Earth (Papike et al. 2005). The method can be made quantitative if the V content of the parental melt is known, assuming equilibrium between the melt and chromite cores, by calculating oxygen fugacity from the relationship of DVsp to fO2 of Canil (2002), log[d(Vmelt/Vsp)−1] = b log fO2 + c

(25)

Where d, b and c are fit parameters dependent on melt composition and temperature. Papike et al. (2005) use the fit parameters for komatiite from Canil (1999), which are d = 32.8, b = 0.41, and c = 0.77. Other fit parameters for different compositions, including spinel with low Cr/Al, are provided in the Background Data Set that accompanies Canil (2002), http://www.elsevier. com/locate/epsl. In the absence of parental melt data, Papike et al. (2005) used whole-rock V concentrations to approximate the V contents of parental melts, and obtained results that are broadly consistent with those of other studies. As it currently stands, application of the V-in-chromite oxybarometer is limited to fO2’s between IW – 1 and IW + 4, which is the range over which DVsp for chromite has been determined (Canil 2002). In addition, V4+ becomes dominant at ~ FMQ (IW + 3.5), and the oxybarometer loses resolution due to the limits of the crystal chemistry of spinel. The relationship of DVsp

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to fO2 is dependent on melt composition, P and T, and the user should take care to choose appropriate fit parameters. Furthermore, application of the oxybarometer to spinels in which Cr and Al vary significantly, or to ulvöspinel-rich compositions, could yield spurious results. Regardless, this oxybarometer relies only on EMP data, and so is widely accessible. Perhaps the most powerful use of vanadium to determine oxygen fugacity relies on its geochemical behavior in combination with its existence in multiple valence states. The V/Sc ratio in terrestrial basalts and mantle xenoliths has been used to infer the oxygen fugacity of the primary magma or mantle source (Lee et al. 2003, 2005; Li and Lee 2004). Vanadium and scandium behave so similarly in magmatic systems that the V/Sc ratio will remain largely unaffected by olivine fractionation, cryptic metasomatism, crustal contamination, or degassing (Canil 2004; Lee et al. 2005). Oxygen fugacity will have the most significant effect on V/Sc; whereas V has variable valence, Sc exists only as Sc3+. Therefore, the V/Sc ratio will “see through” post-extraction processes and reflect the oxygen fugacity of magmagenesis. Lee et al. (2005) implemented the V/Sc oxybarometer by modeling the dependence of V/Sc on oxygen fugacity, under assumptions of isothermal (1410 °C) and isobaric (1.5 GPa) partial melting within the spinel stability field, and a fertile convecting mantle with a constant V/Sc ratio. Analyses of whole-rock V and Sc concentrations are all that are required; these values are compared to the modeled values to determine the primary oxygen fugacity. The results of Lee et al. (2005) have significant implications for the question of whether basalt oxygen fugacity reflects that of its mantle source.

THE BASALT-MANTLE SOURCE REDOX RELATIONSHIP Is basalt oxygen fugacity reflective of the redox state of its mantle source? Whether the oxygen fugacity of a planetary basalt reflects the redox state of the mantle source from which it was derived is of fundamental importance in determining the redox states and histories of planetary interiors. In the absence of mantle samples from the other terrestrial planets, insights into the basalt-mantle source relationship can be gained from attempts to explain the diversity of basalt oxygen fugacity on the Earth. The relationship between oxygen fugacity and FeO and Fe2O3 in silicate liquids (Eqn. 8) is extended by Kress and Carmichael (1991) to quantify the effect of pressure. The Fe3+/ ΣFe ratio of a melt closed to oxygen during its ascent will change such that the oxygen fugacity will be broadly parallel to FMQ; i.e., the relative oxygen fugacity is nearly independent of pressure. Pressure affects the solid oxygen buffers differently, with the FMQ buffer changing by −0.17 log units/GPa and the NNO buffer by −0.51 log units/GPa (Kress and Carmichael 1991). The implication is that a silicate liquid that is closed to oxygen during ascent will retain a record of its oxygen fugacity relative to FMQ to within a fraction of a log unit; therefore, glassy lavas can be used as probes of planetary interior redox state. Given that there is a 7-log unit range in oxygen fugacity in basic lavas on the Earth, it follows that there must exist mantle sources with an equivalent range of oxygen fugacity (Carmichael 1991). Other workers hold a different view, in which the C-H-O volatile species play a more significant role in buffering basalt mantle sources, or ascending magmas, and in determining the oxygen fugacity of the erupted basalt (Mathez 1984; Blundy et al. 1991; Ballhaus and Frost 1994). In the Ballhaus and Frost (1994) model, the range of oxygen fugacity of basalts is explained by variation in mantle source redox state combined with pressure effects, and the redox states of mantle sources are constrained to a much narrower range than proposed by Carmichael (1991). Ballhaus and Frost (1994) envision decompression melting, along an adiabatic ascent path, and assume that the melt remains in major element equilibrium with the crystalline phases of the mantle material until fairly low pressure. They choose an arbitrary

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initial oxygen fugacity of FMQ – 4 at a depth of 4 – 5 GPa; at these conditions, the system is buffered by ferrous-ferric equilibria. Although this is within graphite-water-methane stability, the oxygen fugacity, and hence CO2 activity, are too low to be controlled by C-H-O equilibria. Given that the relative change in oxygen fugacity with increasing pressure is approximately −0.6 log units/GPa (Ballhaus 1995), upwelling asthenosphere will experience oxidation (relative to FMQ, for example) with decreasing pressure. This results in oxidation of graphite, an increase in the activity of CO2, and a shift in buffering from ferrous-ferric to C-H-O; that is, buffering by the presence of graphite and the activity of CO2. At some point, the adiabatic ascent path intersects the dry solidus of the asthenosphere, causing a rapid advance in the degree of partial melting. Eventually, the melt reaches a critical pressure and oxygen fugacity interval where the solubility of carbon in the melt as CO2 exceeds the amount of elemental carbon present in the residue; graphite is eliminated as a phase, and the buffering switches to ferrous-ferric equilibria, with an increase in relative oxygen fugacity with further decompression. Mathez (1984) outlines a model in which degassing of C-rich vapor species play a significant role in buffering fO2. An example is provided in which a C-supersaturated, reduced magma degasses in a shallow magma chamber (depths of < 0.3 GPa), exsolving a CO-rich gas and liberating O2, which oxidizes iron in the melt according to equation (7). Assuming slow, continuous and infinitesimal exsolution of the vapor, an initially reduced (~IW) magma can be oxidized such that the erupted melt approaches FMQ (Mathez 1984). However, this model assumes a magma that is C-supersaturated and reduced; contrast this with the Ballhaus and Frost (1994) model, in which the magma is no longer buffered by graphite by the time it has reached shallow depths. The Ballhaus and Frost (1994) model requires much less variation of redox state among basalt mantle sources. Instead of an intrinsic variation in redox state of mantle sources, the oxygen fugacity of the basalt at the surface will depend on the depth at which buffering switches from C-H-O to ferrous-ferric: the greater the depth at which graphite becomes eliminated in the residue, the more oxidized the melt will be at the surface. The model is used to explain the higher oxygen fugacity of OIB relative to MORB; because MORBs are typically derived from shallow mantle sources, their recorded oxygen fugacity is only slightly higher than that at which major melting (and separation from graphite buffering) occurred. Ocean island basalts, on the other hand, are more oxidized because melting occurs at greater depths and the melt experiences more relative oxidation during ascent. The much higher oxygen fugacity of island arc basalt relative to MORB and OIB can be explained by a source that is intrinsically more oxidized than graphite stability, due to the oxidation of the mantle wedge by slab-derived fluids. As such, the IAB source may be around FMQ, and these basalts have high oxygen fugacities as a result of relatively deep major melting. In contrast to Carmichael (1991), this model implies that the oxygen fugacities of MORB and OIB can be produced from the same mantle, by simply changing the depth of first major melting; the Mathez (1984) model implies that oxidation can occur by stalling and degassing the magma in a crustal reservoir. Such models highlight a potential incongruity between the oxygen fugacity of the basalt collected at the surface, and the redox state of the mantle from which the basaltic melt was derived. Insight into which of these models is more accurate may be derived from the use of V/ Sc ratios to infer the oxygen fugacity of melting for terrestrial basalts and mantle xenoliths (Lee et al. 2005). Using new V and Sc measurements, and data from the literature, Lee et al. (2005) obtain oxygen fugacity results of FMQ − 2 to FMQ for peridotites. This is in good agreement with the results from ferrous-ferric mineral equilibria (e.g., olivine-pyroxene spinel oxybarometry) for suboceanic abyssal peridotites and undepleted, fertile subcontinental mantle xenoliths. V/Sc results for MORB are self-consistent with the results for peridotites, falling within FMQ − 1.25 and FMQ + 0.25 (Lee et al. 2005), which agree with previous studies of MORB oxygen fugacity (Christie et al. 1986; Wood et al. 1990). Remarkably, V/Sc ratios for IAB overlap those from MORB, implying that the IAB and MORB mantle sources have the same oxygen fugacity, between FMQ − 1.25 and FMQ + 0.25. This is in contrast to

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results from the application of equation (8) to glassy arc lavas, which yield FMQ to FMQ + 6 (Carmichael 1991). The implication of the V/Sc results is that the oxygen fugacity of the Earth’s asthenosphere is buffered, perhaps in a manner suggested by Blundy et al. (1991) or Ballhaus and Frost (1994). It is notable that in the range of 1 to 3 GPa, the CCO buffer curve falls between FMQ − 2 and FMQ − 1 (Ballhaus and Frost 1994). Studies of V and Cr in basalts and mantle xenoliths with ages ranging from Archean to the present demonstrate that the oxygen fugacity of the Earth’s asthenosphere has remained buffered to within a log unit of FMQ throughout its history (Delano 2001; Canil 2002; Lee et al. 2003; Li and Lee 2004). The implications of the V/Sc results for the interpretation of oxygen fugacity results from planetary basalts are profound. The observation that the oxygen fugacity of IAB (as derived from mineral equilibria and ferrous-ferric ratios in glass) is several log units higher than that of its source implies that post-extraction processes, such as fractional crystallization, dissociation of volatiles and degassing, auto-oxidation (e.g., Holloway 2004), and hydrothermal alteration may significantly affect the oxygen fugacities recorded by most oxybarometers. Furthermore, it calls into question whether models of basaltic magma evolution in which the system is closed to oxygen (e.g., Kress and Carmichael 1991; Ghiorso 1997) are applicable to most natural systems.

Implications for understanding the redox states of planetary interiors There are several methods now available for determining the oxygen fugacities of planetary basalts; many of these methods have been applied to samples from the Moon, Mars and asteroids, as summarized by Wadhwa (2008). The debate and uncertainty regarding the cause of the variation in oxygen fugacity of terrestrial basaltic samples is instructive – in the absence of mantle xenoliths, what can be said about the redox state of the interiors of the other terrestrial planets? The observation from V/Sc ratios that the mantle source of arc basalts is reduced, while the corresponding eruptives are up to several log units more oxidized, indicates that basalts from arc environments are open systems whose oxygen fugacity reflects post-extraction processes such as differentiation, degassing and assimilation; the same might be true for basalts from other tectonic environments. Although plate tectonics is removed as a complicating factor when discussing other terrestrial planets, the question of whether planetary basalts are open to oxygen during ascension and eruption is equivocal. Variations in oxygen fugacity, such as those observed for Mars (Wadhwa 2001; Herd 2003), need to be interpreted in the context of other indicators of mantle source characteristics. In spite of the lack of mantle xenoliths among our sample suites from the other terrestrial planets, insights into redox states of planetary interiors can be gained through the consideration of the relative roles of volatiles and Fe-bearing mineral equilibria, by analogy with the Earth. High-pressure experiments on mantle or primitive basalt compositions can assist in elucidating the relative change in oxygen fugacity with depth, and determine whether processes such as iron disproportionation have influenced the redox states of the lower mantles of the larger bodies (e.g., Venus, Mars). Comparative studies are particularly relevant and lead to new avenues of research. For example, given the large differences in the C-H-O budgets of the Earth, Moon, Mars and asteroids, what factors in the formation and geologic evolution of the planetary body have the greatest influence on whether mantle source redox state is dominated by volatiles or Febearing mineral equilibria? Given the similarities in C-H-O budgets of the Earth and Mars, what is the role of volatiles in controlling basalt oxygen fugacity on Mars? How can the observation of methane in the martian atmosphere (Formisano et al. 2004) be interpreted in this context? The relatively high (~FMQ) and constant redox state of the Earth’s asthenosphere since the Archean has implications for the speciation of volatiles that were available during the development of life (Delano 2001; Canil 2002). Likewise, insights into the redox evolution of Mars will influence our understanding of the conditions on early Mars and contribute to the assessment of whether the environment of early Mars was conducive to life.

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ACKNOWLEDGMENTS The author thanks all of the participants of the Oxygen in the Terrestrial Planets Workshop for contributing to the discussion of redox conditions in planetary samples. This paper benefited from discussion with Bob Luth, Karlis Muehlenbachs, Tom Chacko and Thomas Stachel. Thanks to Cin-Ty Lee for the fugacity poetry. A thorough review by John Longhi and the editorship of Steve Simon are gratefully acknowledged.

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O’Neill HS, Pownceby MI (1993) Thermodynamic data from redox reactions at high temperatures 1. An experimental and theoretical assessment of the electrochemical method using stabilized zirconia electrolytes, with revised values for the Fe- FeO, Co-CoO, Ni-NiO and Cu-Cu2O oxygen buffers, and new data for the W-WO2 buffer. Contrib Mineral Petrol 114:296-314 O’Neill HS, Wall VJ (1987) The olivine-orthopyroxene-spinel oxygen geobarometer, the nickel precipitation curve, and the oxygen fugacity of the Earth’s upper mantle. J Petrol 28:1169-1191 O’Neill HSC, Rubie DC, Canil D, Geiger CA, Ross CR, Seifert F, Woodland AB (1993) Ferric iron in the upper mantle and in transition zone assemblages: Implications for relative oxygen fugacities in the mantle. In: Evolution of the Earth and Planets. Takahashi T, Jeanloz R, Rubie DC, (eds). American Geophysical Union, Washington D.C., p 73-88 Papike JJ, Karner JM, Shearer CK (2005) Comparative planetary mineralogy: Valence state partitioning of Cr, Fe, Ti, and V among crystallographic sites in olivine, pyroxene, and spinel from planetary basalts. Am Mineral 90:277-290 Parkinson IJ, Arculus RJ (1999) The redox state of subduction zones: insights from arc-peridotites. Chem Geol 160:409-423 Philpotts J (1970) Redox estimation from a calculation of Eu2+ and Eu3+ concentrations in natural phases. Earth Planet Sci Lett 9:257-268 Pownceby MI, O’Neill HSC (1994) Thermodynamic data from redox reactions at high temperatures. IV. Calibration of the Re-ReO2 oxygen buffer from EMF and NiO+Ni-Pd redox sensor measurements. Contrib Mineral Petrol 118:130-137 Sack RO, Carmichael ISE, Rivers M, Ghiorso MS (1980) Ferric-ferrous equilibria in natural silicate liquids at 1 bar. Contrib Mineral Petrol 75:369-376 Sack RO, Ghiorso MS (1989) Importance of considerations of mixing properties in establishing an internally consistent thermodynamic database - Thermochemistry of minerals in the system Mg2SiO4-Fe2SiO4-SiO2. Contrib Mineral Petrol 102:41-68 Sack RO, Ghiorso MS (1991a) Chromian spinels as petrogenetic indicators - Thermodynamics and petrological applications. Am Mineral 76:827-847 Sack RO, Ghiorso MS (1991b) An internally consistent model for the thermodynamic properties of Fe-Mgtitanomagnetite-aluminate spinels. Contrib Mineral Petrol 106:474-505 Schnetzler CC, Philpotts JA (1970) Partition coefficients of rare earth elements between igneous matrix material and rock-forming mineral phenocrysts-II. Geochim Cosmochim Acta 34:331-340 Schwab RG, Küstner D (1981) The equilibrium fugacities of important oxygen buffers in technology and petrology. Neues Jahrbuch für Mineralogie-Abhandlungen 140:111-142 Shannon RD (1976) Revised effective ionic radii and systematic studies of interatomic distances in halides and chalcogenides. Acta Crystallogr 32:751-767 Stachel T, Harris JW, Brey GP, Joswig W (2000) Kankan diamonds (Guinea) II: lower mantle inclusion parageneses. Contrib Mineral Petrol 140:16-27 Stolper E (1977) Experimental petrology of eucritic meteorites. Geochim Cosmochim Acta 41:587-611 Sun CO, Williams RJ, Sun SS (1974) Distribution coefficients of Eu and Sr for plagioclase-liquid and clinopyroxene-liquid equilibria in oceanic ridge basalt - Experimental study. Geochim Cosmochim Acta 38:1415-1433 Sutton SR, Karner J, Papike J, Delaney JS, Shearer C, Newville M, Eng P, Rivers M, Dyar MD (2005) Vanadium K edge XANES of synthetic and natural basaltic glasses and application to microscale oxygen barometry. Geochim Cosmochim Acta 69:2333-2348 van der Hilst RD, Widiyantoro S, Engdahl ER (1997) Evidence for deep mantle circulation from global tomography. Nature 386:578-584 Van Orman JA, Grove TL, Shimizu N (2001) Rare earth element diffusion in diopside: influence of temperature, pressure, and ionic radius, and an elastic model for diffusion in silicates. Contrib Mineral Petrol 141:687703 Verhoogen J (1962) Oxidation of iron-titanium oxides in igneous rocks. J Geol 70:168-181 Wadhwa M (2001) Redox state of Mars’ upper mantle and crust from Eu anomalies in shergottite pyroxenes. Science 291:1527-1530 Wadhwa M (2008) Redox conditions on small bodies, the Moon and Mars. Rev Mineral Geochem 68:493-510 Weill DF, McKay GA (1975) The partitioning of Mg, Fe, Sr, Ce, Sm, Eu, and Yb in lunar igneous systems and a possible origin of KREEP by equilibrium partial melting. Proc Lunar Planet Sci Conf 6:1143-1158 Wones DR, Gilbert MC (1969) The fayalite-magnetite-quartz assemblage between 600 ° and 800 °C. Am J Sci 267A:480-488 Wood BJ (1990) An experimental test of the spinel peridotite oxygen barometer. JGR-Solid Earth Planets 95:15845-15851 Wood BJ (1991) Oxygen barometry of spinel peridotites. Rev Mineral Geochem 25:417-431

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 555-569, 2008 Copyright © Mineralogical Society of America

Rheological Consequences of Redox State Stephen Mackwell Lunar and Planetary Institute 3600 Bay Area Blvd Houston, Texas 77058, U.S.A. [email protected]

ABSTRACT Within the upper mantle of Earth, there is a gradient from a relatively more oxidizing near-surface region, with oxygen fugacities near those of the fayalite-magnetite-quartz buffer (FMQ), to more reducing conditions at depth near the iron-wüstite buffer (IW). Oxygen fugacity appears to vary laterally, as well as vertically, by as much as a factor of 104. As flow within the interior of the Earth and other terrestrial planets occurs due to the (mostly) solidstate deformation of rocks, an understanding of the effect of oxygen fugacity on creep is critical in modeling planetary interior dynamical behavior. This is especially important for the asthenosphere, that anomalously weak region in the uppermost mantle that accommodates isostasy and largely decouples mantle convection from plate motions. Experimental studies of the rheological behavior of iron-bearing minerals have demonstrated that oxygen fugacity can play an important role in deformation. We have shown that olivine rich rocks deformed near FMQ deform in the dislocation creep regime about a factor of 6 faster when buffered near FMQ than at IW. Experiments on olivine single crystals and aggregates indicate that this difference in behavior results from an increase in the concentration of silicon vacancies under more oxidizing conditions, as dislocation creep is rate-limited by the climb of dislocations, which is controlled by diffusion of silicon defects. Although fewer data are available for the effects of oxygen fugacity on pyroxene deformation, clinopyroxene appears to be stronger under more oxidizing conditions, while the data on orthopyroxene deformation show no dependence on oxygen fugacity. These results indicate that vertical and lateral variations in oxygen fugacity may result in, at most, an order of magnitude difference in viscosity, while other factors, such as water fugacity and lithology may be more significant.

INTRODUCTION Experimental studies of the rheological behavior of iron-bearing minerals have demonstrated that oxygen fugacity can play an important role in deformation. As flow within the interior of the Earth and other terrestrial planets occurs due to the (mostly) solid-state deformation of rocks composed of such minerals, an understanding of the effect of oxygen fugacity on creep is critical in modeling planetary interior dynamical behavior. Within the Earth, there is a gradient from a relatively more oxidizing near-surface region, with oxygen fugacities near those of the fayalite-magnetite-quartz buffer (FMQ), to more reducing conditions at depth (for recent reviews, see McCammon 2005, Herd 2008). It is generally accepted that the oxygen fugacity in the lower mantle is buffered near the iron-wüstite buffer (IW). Measurements of oxygen fugacity from rocks derived from the upper mantle suggest that oxygen fugacity varies significantly both vertically and laterally (McCammon 2005; Herd 2008). Such variability may correspond to as much as a factor of 104 decrease in oxygen fugacity at asthenospheric depths. As the asthenosphere is the anomalously weak region in the uppermost mantle that accommodates isostasy and largely decouples mantle convection from plate motions, it is 1529-6466/08/0068-0020$05.00

DOI: 10.2138/rmg.2008.68.20

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important to quantify the difference in mechanical behavior of mantle rocks as a function of oxygen fugacity. The following sections describe the current understanding of the effect of oxygen fugacity on deformation of mantle minerals of the terrestrial planets, including some new data from our current research.

DEFORMATION OF OLIVINE Olivine (MgxFe1−x)2SiO4 is the most abundant mineral in the upper mantle of Earth (with x~0.9) and, likely, those of the other terrestrial planets (x~0.7-0.9). Based on experimental measurements and observations from ophiolites and other olivine-bearing rocks exposed at the Earth’s surface, it is also believed to be the weakest mineral in the upper mantle. Thus, the mechanical behavior of the upper mantle is approximated well by the rheological behavior of olivine aggregates (dunites). There have been numerous studies of the deformation behavior of olivine, dating back to the earliest rheological experiments on olivine aggregates (e.g., Raleigh 1968; Carter and Avé Lallemant 1970; Raleigh and Kirby 1970; Kirby and Raleigh 1973; Post 1977) and single crystals (e.g., Blacic 1972; Phakey et al. 1972). These early studies made no attempt to control, least of all vary, oxygen fugacity, however; only in the last three decades has there been a systematic attempt to constrain the role of oxygen fugacity in deformation experiments on olivine single crystals and aggregates.

Olivine single crystal studies Although the earliest work on deformation of olivine single crystals dates to the pioneering work of Phakey et al. (1972) and Blacic (1972) using solid-medium deformation apparatus, rigorous control of oxygen fugacity was first attempted by Kohlstedt and Goetze (1974) on unoriented samples of single-crystal olivine, followed by the work of Durham and Goetze (1977) on single crystals oriented to favor slip on the easiest slip systems. Both of these studies used a fixed mixture of CO2 and H2 to control oxygen fugacity. While neither study attempted to vary the oxygen fugacity during the experiment, they both acknowledged that iron-bearing natural olivine is only stable under conditions significantly more reducing than in air. In addition, the study reported in Durham and Goetze (1977) and Durham et al. (1977) demonstrated that, at high temperatures in the range 1080 to 1575 °C, significant differences in dislocation microstructure accompany the different slip systems. Interestingly, these authors observed the dislocation microstructures optically in thin sections, using an oxidation technique (Kohlstedt et al. 1976) to decorate dislocations by briefly heating the thin sections in air. The first studies to investigate the dependence of creep of olivine single crystals on oxygen fugacity were performed by Hornack (1978), Jaoul et al. (1980) and Kohlstedt and Hornack (1981). Working with natural olivine single crystals that were unbuffered in silica activity and using mixed CO-CO2 gases to control oxygen fugacity, they demonstrated a positive dependence of creep rate on oxygen fugacity, with the strain rate approximately proportional to fO21/6, consistent with point defect models (e.g., Stocker 1978a, 1978b; Stocker and Smyth 1978). In a subsequent study, Ricoult and Kohlstedt (1985) deformed olivine single crystals buffered against either orthopyroxene (Mg,Fe)SiO3 or magnesiowüstite (Mg,Fe)O to control the silica activity. Although the oxygen fugacity was nominally varied over a range of conditions using mixtures of CO and CO2 gases, a subsequent investigation by Bai et al. (1991) demonstrated that the tungsten/molybdenum furnace and foils used in the Ricoult and Kohlstedt (1985) study effectively buffered the oxygen fugacity at a fixed, much more reducing condition than suggested by the gas mixtures, so that no range in oxygen fugacity was actually investigated. Bai et al. (1991) used a new room-pressure apparatus with internal components that did not react with the mixed CO-CO2 gases, and with a downline zirconia oxygen sensor that monitored the oxygen fugacity of the gas mixture. They performed a detailed study of the

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creep behavior of natural olivine single crystals, investigating three orientations that favor the four easiest slip systems, varying temperature from 1200 to 1525 °C, oxygen fugacity from 10−12 to 10−3 atm, and compressive stresses between 14 and 180 MPa to yield creep rates between 10-4 and 10−6.7 s−1. Samples were buffered in silica activity using either orthopyroxene or magnesiowüstite buffers. While the dependence of creep rate on applied stress (σ) was remarkably constant, with creep rate proportional to σ3.5, they noted significant complexity in creep behavior with variation in oxygen fugacity, silica buffer and temperature. Figure 1 (after Fig. 5 of Bai et al. 1991) illustrates this complexity for samples deformed favoring the easiest high-temperature slip system in olivine (010)[100], involving motion of [100] Burgers vector dislocations on the (010) planes. The dependence of creep rate on oxygen fugacity changes as oxygen fugacity is increased at lower temperatures, and a third dependence on oxygen fugacity becomes apparent at higher temperatures. Thus, creep of olivine single crystals favoring this slip system shows changes in dependence on oxygen fugacity and temperature as oxygen fugacity and temperature are varied, requiring complex formalisms (see equations in Table 4 of Bai et al. 1991) to describe the high-temperature deformation for each slip system. In a subsequent analysis of the sample microstructures using optical and electron microscope techniques (Bai and Kohlstedt 1992), the dislocation substructures were shown to be distinctly different between many of the deformation fields that possess unique fugacity dependencies and activation energies. These observations suggest that extrapolation of experimentally determined rheologies to model the behavior of the Earth or other planets must be done cautiously in order to ensure that the constitutive parameters are appropriate to the deformation mechanism operative in that body (Mackwell et al. 1990). Recent unpublished work by S. Mackwell, G. Hirth and D. Kohlstedt on deformation of olivine single crystals in high-pressure gas-medium deformation apparatus demonstrates a

-4 (010)[100] opx buffer -5

30 MPa Figure 1. Plot of strain rate versus oxygen fugacity from experiments by Bai et al. (1991) on single crystals of olivine buffered by orthopyroxene and oriented to favor slip on (010)[100], the easiest high-temperature slip system in olivine. The symbols represent strain rate measurements at a variety of temperatures for an applied stress of 30 MPa. The solid lines are fits to the experimental data, which have been decomposed into individual creep laws illustrated by the dashed lines. The dotted lines show the oxygen fugacities for iron-wüstite (IW) and nickel-nickel oxide (NNO) for each temperature.

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similar dependence of creep rate on oxygen fugacity as the room pressure studies. In this work, the oxygen fugacity was controlled using either a Ni jacket with a NiO powder coating the samples (NNO buffer), or an Fe jacket with an FeO coating (IW buffer). It is worth noting that the NNO buffer is ~1 order of magnitude in oxygen fugacity more oxidizing than the fayalitequartz-magnetite (FMQ) buffer (Fig. 2). Figure 3 shows that, in general, samples buffered at IW have lower strain rates, and are therefore stronger than, samples buffered at NNO under both wet and dry deformation conditions.

Olivine aggregate studies As noted above, dunites (olivine aggregates) have been the focus of experimental deformation studies for more than forty years. Early pioneering studies by Raleigh (1968), Carter and Avé Lallemant (1970), Raleigh and Kirby (1970), Kirby and Raleigh (1973), and Post (1977) mapped out the deformation behavior of olivine aggregates using solid-medium, high-pressure deformation apparatus. These studies established that the behavior followed a power-law behavior, where the strain rate is proportional to σn, where n lies in the range 3.5 to 5.5, consistent with deformation by dislocation creep. While these studies were pivotal in establishing rheologies applicable to lithospheric and asthenospheric processes, they were not well constrained in terms of the chemical state of the samples during deformation, notably in terms of the oxygen and water fugacity. While it is often assumed that the talc solid-medium assemblies approximately buffer the samples near FMQ, oxygen fugacity was not rigorously constrained. Also, the dehydration of hydrous minerals that are ubiquitous in natural dunites likely resulted in some water-weakening of the samples. The presence of some orthopyroxene in dunite probably buffered the silica activity. In later experiments on dunites in the solid-medium apparatus, Zeuch and Green (1984) attempted to minimize a number of these problems. They used samples synthesized from olivine sands and made improvements in the solid-medium apparatus that allowed an increase in stress resolution and decrease in o

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Figure 2. Plot of oxygen fugacity versus inverse temperature illustrating the phase boundaries for the iron-wüstite (IW), wüstite-magnetite (WM) and magnetite-hematite (MH), NNO, and fayalite (Fe2SiO4) systems at room pressure. The subvertical solid line represents the melting of fayalite at room pressure. IQF represents the phase boundary between fayalite and quartz/iron, and FMQ represents that between fayalite and quartz/magnetite.

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σ (MPa) 50

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Figure 3. Plots of strain rate versus stress for olivine single crystals deformed under wet (bottom) and dry (top) conditions at 1300 °C and buffered against orthopyroxene, for slip on the easiest high-temperature slip system. The oxygen fugacity was controlled using IW (solid symbols, solid lines) or NNO (open symbols, dashed lines) buffers. The different symbols represent distinct experiments, with the lines fit to the data using least squares regressions. The dotted lines represent the flow laws of Bai et al. (1991) for single crystals of olivine deformed at room pressure interpolated to the oxygen fugacities of the iron-wüstite and nickel-nickel oxide buffers.

2.2

log σ (MPa) thermal gradients. However, it was only when experiments were performed in the gas-medium deformation apparatus that rigorous control of oxygen fugacity was attained. Chopra and Paterson (1981, 1984) deformed dunite samples in a gas-medium apparatus with an internal load cell and minimal thermal gradients. Their dry experiments were performed on samples that had been pre-annealed under controlled oxygen fugacities at room pressure to remove any hydrous minerals without oxidizing the samples. As the samples were jacketed in metal sleeves during the deformation, it was believed that oxidation of the inside surface of the jacket resulted in effective buffering of the oxygen fugacity within the sample. This expectation is certainly true for iron jackets, even though the oxide layer does partially react with the sample. It is not clear, however, whether nickel or copper jackets oxidized during the experiments to provide an effective buffer, or if the dunite sample itself buffered the oxygen fugacity at conditions more reducing than NNO or Cu/CuO. Unfortunately, the authors do not report unambiguously the presence of an oxide phase after the experiments. There is also a possibility that any oxide phase that did form reacted with the olivine sample to form (Fe,Ni)O or (Fe,Cu)O that might have buffered the assemblage at much more reducing conditions than the end-member oxides. Based on their experiments, Chopra and Paterson (1981, 1984) determined flow laws for dunite deforming by dislocation creep under both wet and dry conditions with a stress exponent of around 3.5 and an activation energy of around 540 kJ/mol. They could discern no difference in mechanical behavior due to a change in the jacketing material, and reported that dunite deformation shows no dependence on oxygen fugacity. Subsequent to their work, Karato et al. (1986) used the same apparatus to deform dunite analogs that had been synthesized from ground powders of natural olivine single crystals. They

560

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only investigated a single temperature (1300 °C) but were able to confirm, in general terms, the results of Chopra and Paterson (1981, 1984), and assumed the same activation energy as reported in the previous studies. Due to the finer grain size of the aggregates, Karato et al. (1986) were also able to investigate the deformation field dominated by grain boundary diffusional creep. No attempt was made to investigate the role of oxygen fugacity, and the authors assumed that their samples were buffered by the iron jacket at IW. It is interesting to note that the synthesized dunite samples of Karato et al. (1986) that were deformed dry within the dislocation creep field were generally weaker than the natural samples of Chopra and Paterson (1984). Hirth and Kohlstedt (1995) interpreted this disparity as resulting from the greater component of grain boundary sliding in the finer-grained material of the synthesized samples. More recent investigations in the dislocation creep field by Hirth and Kohlstedt (1995) on synthetic dunite with a component of partial melt, and by Mei and Kohlstedt (2000), looking at the role of water in weakening of dunite, used nickel sleeves around their samples on the assumption that oxidation of the sleeve would buffer the oxygen fugacity. In neither case were experiments performed using other oxygen buffers. As their experiments give results similar to previous studies that were buffered by IW, it is easy to interpret their results as indicating that there is no significant effect of oxygen fugacity on creep of dunite. As noted by Hirth and Kohlstedt (1995), however, longer anneal times for their samples prior to deformation resulted in increased grain size and an overall strengthening of their samples, consistent with a reduction in the role of grain boundary sliding with increasing grain size. Thus, direct comparison between different studies using synthesized samples, particularly those annealed for only short times, may not be terribly robust. Nonetheless, when the long-anneal (21 h) results of Hirth and Kohlstedt (1995) are compared with those of Chopra and Paterson (1984), the NNO-buffered samples do appear to be weaker than those buffered at IW. It may seem surprising that deformation of single crystals of olivine, such as in the study by Bai et al. (1991), shows a clear dependence on oxygen fugacity but deformation of dunite appears to show only a weak dependence. Such an apparent contradiction can potentially be explained through an understanding of how single crystal flow laws can be combined to describe aggregate behavior. In order for an aggregate of randomly oriented grains of a single mineral to deform homogeneously by dislocation motion on slip systems within individual grains, geometric arguments require a minimum of 5 independent slip systems to be operative (von Mises 1928). While most deformation will be accommodated with the easiest slip systems, some slip is required on all. This criterion can be relaxed somewhat by allowing some component of dislocation climb or grain boundary sliding, so that, in general, three independent slip systems seem to be sufficient for aggregate deformation. Despite this relaxation, it is not clear whether the aggregate strength would be defined by the weaker or stronger of the required slip systems, and no simple formalism allows a prediction of aggregate behavior based on single crystal flow laws. Using the flow laws for olivine single crystals buffered by orthopyroxene from the study of Bai et al. (1991), Figure 4 shows the strength of the major slip systems for cases where the oxygen fugacity of the deforming samples is buffered by either NNO (heavy lines) or IW (lighter lines). As is evident, the contrast in strength for NNO- and IW-buffered olivine differs for each slip system and varies significantly with temperature, with more than an order of magnitude contrast in strain rate for the (010)[100] slip system at 1100 to 1200 °C, and less than a factor of 2 contrast in strain rate for the (010)[001] slip system over the same temperature range. Thus, if one assumes that the strongest of the three slip systems controls the creep of an aggregate, little difference in mechanical behavior might be anticipated as a function of oxygen fugacity. However, if either of the weaker slip systems controls aggregate behavior, then oxygen fugacity will have an important effect on creep of dunite.

Rheological Consequences of Redox State

561

T (oC) -3

1400

1300

1200

1100

σ = 100 MPa -4

(010)[100]

IW

(010)[001]

-5 NNO

.

-1

log ε (s )

(001)[100]

-6

IW NNO

-7

-8

6.0

6.5 4

7.0

-1

Figure 4. Plot of strain rate versus temperature for deformation of olivine single crystals at an applied stress of 100 MPa, buffered against orthopyroxene, and deforming on the three easiest high-temperature slip systems from the work of Bai et al. (1991). The thick lines show the curves for samples buffered by NNO, while the thinner lines are for samples buffered by IW. The vertical separation between the NNO and IW lines for each slip system is an indication of the dependence of creep on oxygen fugacity for that slip system at that temperature.

7.5

10 /T (K ) Due to concerns about the effectiveness of past attempts at buffering oxygen fugacity in high-pressure apparatus, we (unpublished work by J. Keefner, S. Mackwell, F. Heidelbach and D. Kohlstedt) recently revisited the question of the effect of oxygen buffering on dunite behavior (see also Keefner et al. 2005). As aggregate deformation generally requires high pressures to limit microcrack formation and a contribution of brittle processes to sample deformation, we used a similar high-pressure gas-medium apparatus to that previously used by Chopra and Paterson (1981, 1984), Karato et al. (1986), Hirth and Kohlstedt (1995) and Mei and Kohlstedt (2000). As in previous studies, we used metal sleeves (iron or nickel) around the samples. However, we coated the surfaces of the samples and interiors of the sleeves with metal oxide prior to deformation. After the experiments, we carefully removed the sleeve from around each sample and verified that the oxide phase was still present. As NiO and Ni-olivine have similar colors but very different buffering characteristics, we used microprobe analyses of the surface coatings in addition to visual inspection. Figure 5 summarizes the results of our study. All data from all experiments have been normalized to a single temperature (1250 °C) using an activation energy of 540 kJ/mol for both NNO- and IW-buffered experiments. This activation energy was determined as the mean of the calculations from all experiments, as the energies determined for the two buffers are essentially indistinguishable. The solid lines represent fits to the results for NNO and IW experiments; the dashed line shows the results from Chopra and Paterson (1984) for dunite samples from the same source deformed at IW; and the dash-dot line represents the results from Hirth and Kohlstedt (1995) for their 21-h annealed synthesized dunite samples deformed nominally at NNO. Direct comparison with the results of Karato et al. (1986) and Mei and Kohlstedt (2000) is not

562

Mackwell

σ (MPa) 100

.

-1

log ε (s )

-4.0

-5.0

200

Aheim Dunite

300

400

o

Normalized to 1250 C NNO IW IW: C&P NNO: H&K

meaningful due to the increased role of grain-boundary sliding in their experiments (Hirth and Kohlstedt 1995). Our data clearly demonstrate an effect of oxygen fugacity on creep of dunite under these experimental conditions, with a factor of ~6 increase in strain rate resulting from ~4 orders of magnitude increase in oxygen fugacity.

When we normalize the dunite creep data to an applied stress of 100 MPa and overlay them on the single crystal -6.0 deformation curves of Bai et al. (1991), we find that the dunite data lie at strain rates 2.0 2.2 2.4 2.6 intermediate between the strongest and weakest of the log σ (MPa) three studied high-temperature slip systems (Fig. 6). The Hirth Figure 5. Plot of strain rate vs. stress for samples of dunite deformed and Kohlstedt (1995) data for at 300 MPa confining pressure and buffered by orthopyroxene. While the data were collected at a range of temperatures, they have their 21-h annealed samples all been extrapolated to 1250 °C in this figure. The solid symbols are in reasonable agreement and thick solid line represent samples buffered at oxygen fugacities with our results. Interestingly, defined by NNO, while the plus symbols and thinner solid line the NNO-buffered samples show data for samples buffered by IW. C&P represents the flow appear to require less activity law for IW-buffered samples deformed by Chopra and Paterson (1984). H&K represents the 21-h annealed, NNO-buffered samples on the strongest slip system by Hirth and Kohlstedt (1995). relative to the IW-buffered samples. Thus, a simple model where aggregate behavior can be defined as being dominated by a single slip system does not work, and oxygen fugacity appears to affect individual slip systems, as well as the contribution of grain-boundary sliding to overall flow behavior.

How does oxygen fugacity affect creep of olivine? Oxygen fugacity within a single mineral grain is expressed in the form of point defects. More oxidizing conditions are reflected in lower concentrations of oxygen vacancies and/ or higher concentrations of cation vacancies. As most minerals of importance in Earth’s interior are either insulators or semi-conductors, excess charge associated with point defects must be locally compensated in the crystal lattice by oppositely charged point defects. Thus, octahedrally coordinated cation (metal) vacancies in olivine are generally charge-compensated by two local ferric iron ions occupying metal sites, and increased oxygen fugacity correlates with higher concentrations of ferric iron and metal vacancies. As diffusion of divalent cations in olivine is directly coupled to the concentration of metal vacancies, it is easy to demonstrate that higher oxygen fugacities lead to higher rates of diffusion for iron and magnesium. While a direct correlation between diffusion and oxygen fugacity is fairly straightforward, the basis for an effect of oxygen fugacity on creep in the dislocation creep field is not so clear. The correlation between dunite deformation data and the single crystal results from Bai et al. (1991) demonstrates, to first order, that the mechanical behavior of dunite is controlled by the

Rheological Consequences of Redox State o

T ( C) -3

1400

1300

1200

1100

σ = 100 MPa -4

(010)[100]

IW

(010)[001]

-5 NNO

.

log ε (s-1)

(001)[100]

-6

IW

NNO

-7

-8

6.0

6.5 4

7.0

-1

7.5

10 /T (K ) Figure 6. Plot of strain rate vs. inverse temperature for deformation of olivine single crystals at an applied stress of 100 MPa, buffered against orthopyroxene, and deforming on the easiest 3 high-temperature slip systems from the work of Bai et al. (1991). The various lines are as described in the caption for Figure 4. Superimposed on this diagram are the data from the experimental deformation of dunite, extrapolated to an applied stress of 100 MPa. The open circles are the results from Hirth and Kohlstedt (1995) for their 21-h anneal, NNObuffered samples.

563

motion of dislocations on well-defined slip systems within individual olivine grains. High-temperature dislocation creep in olivine involves both glide of dislocations along crystallographically defined slip planes and climb of dislocations around obstacles to dislocation motion. Weertman (1968), for example, has modeled dislocation motion along parallel crystallographic planes, where glide occurs until electrostatic interactions between dislocations gliding in adjacent slip planes cause the dislocations to lock in place. Rather than remain locked, the edge segments of the dislocations climb toward each other and annihilate. In general, bulk strain represents glide of dislocations while strain rate is defined by dislocation climb. The climb of dislocations is a diffusion-controlled process that is ratelimited by the slowest diffusing intrinsic species in the material. In olivine under high-temperature deformation conditions, the slowest diffusing species is silicon, which is believed to diffuse by a vacancy mechanism (Dohmen et al. 2002). Thus, for oxygen fugacity to affect creep of olivine single crystals, or dunite deforming by dislocation creep, the concentration of silicon vacancies must be dependent on oxygen fugacity.

When diffusion of an ionic species is vacancy-controlled, it is possible to write the ionic self-diffusion coefficient in terms of the vacancy diffusivity and concentration. Thus for silicon diffusion in olivine, DSi = [VSi//// ] DVSi

where DSi is the silicon self diffusivity, DVSi is the diffusivity for silicon vacancies, and [ VSi//// ] is the concentration of silicon vacancies, expressed as a fraction of total silicon sites, and using the point defect notation of Kröger and Vink (1956). In this notation, the species (in this case a vacancy, V) is shown to the left, with the subscript showing the crystallographic site, and the superscript showing the net charge on the defect, which is equal to the charge on the site minus the charge on the species. Thus, a vacant silicon site has a net charge of −4, expressed as ////. Equally, a ferric iron ion in an octahedrally coordinated cation (metal) site is expressed as Fe •Me, where the net single positive charge on the defect is expressed by the • symbol. In olivine, it is generally considered that the predominant defects that define charge neutrality are

564

Mackwell

ferric iron ions on metal sites and vacant metal sites; thus, because ferric iron occupying metal sites is twice as abundant as metal vacancies, // 2[ VMe ] = [Fe •Me ]

The formation of silicon vacancies can be written in the form of a chemical equation Me 2SiO 4 + O 2 + Si×Si + 4Fe ×Me = 4Fe •Me + VSi//// + 2MeSiO 3

where olivine (Me2SiO4) reacts with oxygen to form enstatite (MeSiO3), and × indicates zero charge on the defect. Following the law of mass action and the charge neutrality equation above, we obtain −10 / 3 [VSi//// ] ∝ [ Fe×Me ]4 / 3 fO12/ 3 aen

where aen is the activity of enstatite, fO2 is the oxygen fugacity, the activity of olivine is fixed at aol = 1, and [Si×Si ] ≈1 (see Table III, Stocker and Smyth 1978). More detailed presentation of such point defect arguments pertaining to olivine can be found in Stocker (1978a,b) and Stocker and Smyth (1978). Based on these arguments, and on the assumption that defect diffusivities are largely independent of the chemical environment (see, e.g., Schmalzried 1981) −10 / 3 DSi ∝[VSi//// ] ∝ [ Fe×Me ]4 / 3 fO12/ 3 aen

Thus, thermodynamic arguments would suggest positive dependencies of creep on iron content and oxygen fugacity and a negative dependence on enstatite activity when silicon is the slowest diffusing species in olivine and it diffuses by a vacancy mechanism. Although Houlier et al. (1990) have argued, based on their diffusion experiments, for an interstitial control of silicon diffusion, work by Dohmen et al. (2002) raises concerns about the earlier work. The Dohmen et al. (2002) study clearly demonstrates that silicon diffusion is far slower than oxygen diffusion in olivine, and reports an activation energy for silicon diffusion that is in reasonable agreement with activation energies for creep. It is worth noting, however, that while the activation energy for silicon diffusion appears to match that for creep, the rate of diffusion is still too slow to account for the magnitude of the strain rate. While a dependence of creep rate on oxygen fugacity is apparent in point defect models, control of dislocation motion based solely on diffusion rates for silicon does not fully account for the complex deformation behavior observed by Bai et al. (1991) and Bai and Kohlstedt (1992), suggesting that other factors, such as kink or jog nucleation and/or propagation, may have to also be considered. However, to a first order, this model is consistent with the dunite deformation data, both in terms of the increase in creep rate as a function of oxygen fugacity and in terms of the observation of a decreased requirement for activity on the stronger slip systems at more oxidizing conditions where increased rates of diffusion may further weaken the requirements of the von Mises criterion discussed earlier.

DEFORMATION OF OTHER SILICATES While the deformation behavior of olivine single crystals and aggregates has been the focus of many more studies than that of the other major minerals of the interiors of the terrestrial planets, a number of studies have investigated the mechanical behavior of mantle pyroxenes. Raleigh et al. (1971), Ave Lallemant (1978), Ross and Nielsen (1978) and Kollé and Blacic (1982, 1983) used high-pressure deformation apparatus to study the mechanical behavior of pyroxene aggregates. As in the early work on dunites, no attempt was made to buffer the oxygen fugacity in these experiments. More recent deformation experiments have been performed in room-pressure apparatus by Raterron and Jaoul (1991) on single-crystal

Rheological Consequences of Redox State

565

samples of diopside under controlled oxygen fugacity conditions using mixed Ar/H2O/H2 gases. While they did not test directly for an oxygen fugacity dependence of creep, Jaoul and Raterron (1994), in a subsequent study using the same apparatus, did vary the oxygen fugacity for diopside and demonstrated a dependence of creep rate on oxygen fugacity with a powerlaw exponent of approximately −0.18 at 1100 °C (Fig. 7; see also Jaoul and Raterron 1994, Fig. 1). At 1200 °C, they found no dependence of creep rate on oxygen fugacity. They also noted, however, that silica-rich precipitates in the diopside had a dramatic effect on mechanical behavior at the higher temperature. As these precipitates are only observed after deformation at low confining pressures, the higher-temperature flow law is unlikely to be appropriate to upper mantle deformation. Bystricky and Mackwell (2001) deformed clinopyroxene aggregates in a gas-medium deformation apparatus at 1100 to 1250 °C with the oxygen fugacity buffered by either IW, using an iron jacket, or NNO using a nickel jacket and NiO powder coating the samples. They found little effect of oxygen fugacity on creep behavior in either the diffusional or dislocation creep regimes. They also found no silica-rich precipitates and did not observe the drop in activation energy reported by Jaoul and Raterron (1994). Of note, the iron content of the clinopyroxene in the Bystricky and Mackwell (2001) study is approximately a factor of 2 higher than for the single crystal samples of Jaoul and Raterron (1994). Mackwell (1991) investigated the dependence of creep rate on oxygen fugacity for a range of enstatite (orthopyroxene) compositions using a room-pressure deformation apparatus similar to that of Jaoul and Raterron (1994), but using CO/CO2 gas mixtures to control oxygen fugacity (Fig. 8). He found no dependence of creep rate on oxygen fugacity over the range of conditions and compositions investigated.

-6.0 Diopside {110}1/2 o

1100 C Figure 7. Plot of strain rate vs. oxygen fugacity for room-pressure deformation of single crystal samples of diopside (clinopyroxene) from the study of Jaoul and Raterron (1994). The samples were oriented for creep on the {110}½ slip system, and were deformed at 1100 °C and applied stresses of either 110 or 142 MPa.

.

-1

log ε (s )

-6.5

-7.0 110 MPa

-7.5

-8.0

142 MPa

-14

-12

-10

log fO2 (atm)

-8

566

Mackwell

OPX 12#1 o

1400 C

OPX 5#1

En99.2

1400 C

o

40 MPa En96.2

-1

log ε (s )

-5

35 MPa

.

-6 m = -0.01 ± 0.03 m = 0.00 ± 0.04

-7

-8

-10

-8

-6

-4

-2 -10

log fO2 (atm)

-8

-6

-4

-2

log fO2 (atm)

Figure 8. Plots of strain rate vs. oxygen fugacity for single crystal samples of enstatite (orthopyroxene) deformed at 1400 °C and applied stresses of 35 MPa (left) and 40 MPa (right). The left figure shows data for a sample with a low FeO content (En99.2 or Mg0.992Fe0.008SiO3), while the right figure shows data for a higher FeO content (En96.2 or Mg0.962Fe0.038SiO3). The solid lines represent least squares regression fits to the data, yielding oxygen fugacity exponents for creep indistinguishable from 0.

No work to date has attempted to quantify the effects of oxygen fugacity on the deformation of garnet (the other major upper mantle mineral), wadsleyite, ringwoodite or majorite garnet (the predominant transition zone minerals), or silicate perovskite or magnesiowüstite (the major lower mantle minerals). Of these minerals, those from the transition zone and silicate perovskite cannot be studied using conventional deformation apparatus due to their highpressure stability fields, and garnet requires solid-medium deformation apparatus. Only magnesiowüstite is stable at sufficiently low pressures to allow deformation in gas-medium apparatus or at room pressure, but no studies have yet been performed to characterize the role of oxygen fugacity.

How does oxygen fugacity affect creep of other silicates? As in olivine, it is anticipated that point defects control the effect of oxygen fugacity on the pyroxenes. Jaoul and Raterron (1994) interpret their measurements on diopside single crystals, which show an inverse dependence of creep rate on oxygen fugacity, as indicating rate control of deformation by diffusion of either metal (magnesium/iron) or silicon interstitial defects. On the assumption that charge neutrality is established by // 2[ VMe ] = [Fe •Me ]

then, in analogy to the deformation of olivine discussed above, 7/3 DSi ∝[Si •••• ] ∝ [ Fe×Me ]−4 / 3 fO−21 / 3 aSiO i 2

where Me can be Mg, Fe or Ca. Assuming that silicon diffuses more slowly than the other major elements in diopside, the diffusion rate of the rate-limiting species decreases with increasing oxygen fugacity, consistent with experimental observations. The decrease in diffusion rate, and hence creep, with iron content may also help explain the change from a

Rheological Consequences of Redox State

567

clear dependence of creep on oxygen fugacity in the experiments of Jaoul and Raterron (1994) to nearly no dependence on oxygen fugacity in the experiments of Bystricky and Mackwell (2001). As the latter study was performed on diopside with significantly higher iron content, the consequent reduction in ionic diffusivities may have resulted in full or partial transition to deformation control by another species, such as silicon vacancies. A full characterization of the role of defects in controlling deformation in diopside awaits a more systematic study. In his study of the deformation behavior of enstatite (orthopyroxene), Mackwell (1991) argues that if creep is controlled by diffusion of silicon defects (vacancies or interstitials), then either the charge neutrality for the samples was not established by ferric iron (as all ferric iron was locally charge- compensated by aluminum defects) or there was too little time for the samples to equilibrate to new oxygen fugacities during the experiments. This latter concern was not an issue in the deformation of olivine (see also Mackwell et al. 1988) or diopside (Jaoul and Raterron 1994) in gas-mixing furnaces but remains an open question for enstatite.

SUMMARY Under high-temperature conditions appropriate to the lower lithospheric and sublithospheric regions in the mantles of Earth and the other terrestrial planets, deformation generally occurs by either diffusional creep or dislocation creep processes, both of which are rate-controlled by ionic diffusion. As such, any role for oxygen fugacity in affecting mechanical behavior results from an influence on the concentrations of the point defects that control atomic transport within the minerals or at grain boundaries. As illustrated above, such a model, while perhaps a little simplistic, can be applied in general terms to understand flow in the interior of the terrestrial planets. Oxygen fugacity has a clear effect on the high-temperature mechanical behavior of most terrestrial materials for which a dependency has been measured. Although oxygen fugacity affects the mechanical behavior of various mantle minerals in different ways, as described above, the volumetric predominance of olivine in planetary mantles implies that deformation of mantle rocks more likely conforms to the oxygen fugacity dependence for that mineral. In the Earth’s mantle, the oxygen fugacity variation is likely to span no more than 6 orders of magnitude, defined by the stabilities of the phases present there. Thus, for rocks where an oxygen fugacity dependence of creep has been demonstrated, variations in viscosity of less than a factor of ~10 likely result from differences in oxygen fugacity alone. By comparison, water-weakening has been demonstrated to have a potentially much greater role. Even at relatively modest water fugacities of around 0.3 GPa, the drop in viscosity relative to dry deformation of olivine aggregates is comparable to that resulting from changing the oxygen fugacity by 4-5 orders of magnitude. Thus, while the effect of oxygen fugacity on creep cannot be discounted, the roles of other components of the thermochemical environment are at least equally important to constrain. In addition, definition of the mechanical behavior of the asthenosphere is limited due to uncertainty in the in situ oxygen fugacity, which may result in order of magnitude errors in viscosity. Within the mantle, lateral variations in oxygen fugacity between, for example, cratonic and sub-oceanic upper mantle may influence convection patterns. Of course, other differences between regions of the mantle, such as modal abundance of minerals and spatial variation in water fugacity, are likely to have significant effects that may be comparable to or exceed the effects of oxygen fugacity. Thus, while higher oxygen fugacities in cratonic regions might weaken olivine, reduced water fugacities and lithologies richer in pyroxene will also have important effects on mechanical behavior.

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Mackwell REFERENCES

Ave Lallemant HG (1978) Experimental deformation of diopside and websterite. Tectonophys 48:1-27 Bai Q, Kohlstedt DL (1992) High-temperature creep of olivine single crystals. 2. Dislocation microstructures. Tectonophys 206:1-29 Bai Q, Mackwell SJ, Kohlstedt DL (1991) High-temperature creep of olivine single crystals. 1. Mechanical results for buffered samples. J Geophys Res 96:2441-2463 Blacic JD (1972) Effect of water on the experimental deformation of olivine. In: Flow and Fracture of Rocks: the Griggs Volume, Geophysical Monograph 16. Heard HC, Borg IY, Carter NL, Raleigh CB (eds) AGU, Washington, DC pp 109-115 Bystricky M, Mackwell S (2001) Creep of dry clinopyroxene aggregates. J Geophys Res 106:13,443-13,454 Carter NL, Ave Lallemant HG (1970) High-temperature flow of dunite and peridotite. Geol Soc Amer Bull 81:2182-2202 Chopra PN, Paterson MS (1981) The experimental deformation of dunite. Tectonophys 78:453-473 Chopra PN, Paterson MS (1984) The role of water in the deformation of dunite. J Geophys Res 89:78617876 Dohmen R, Chakraborty S, Becker H-W (2002) Si and O diffusion in olivine and implications for characterizing plastic flow in the mantle. Geophys Res Lett 29:2030 doi:10.1029/2002GL015480 Durham WB, Goetze C (1977) Plastic flow of oriented single crystals of olivine 1. Mechanical data. J Geophys Res 82:5737-5753 Durham WB, Goetze C, Blake B (1977) Plastic flow of oriented single crystals of olivine 2. Observations and interpretations of the dislocation structures. J Geophys Res 82:5755-5770 Herd CDK (200X) Basalts as probes of planetary interior redox state. Rev Mineral Geochem XX:xxx Hirth G, Kohlstedt DL (1995) Experimental constraints on the dynamics of the partially molten upper mantle, 2. Deformation in the dislocation creep regime. J Geophys Res 100:15,441-15,449 Hornack PG (1978) The effect of oxygen fugacity on the creep of olivine. MS Thesis, Cornell University Houlier B, Cheraghmakani M, Jaoul O (1990) Silicon diffusion in San Carlos olivine. Phys Earth Planet Int 62:329-340 Jaoul O, Froidevaux C, Durham WB, Michaut M (1980) Oxygen self-diffusion in forsterite: Implications for the high temperature creep mechanism. Earth Plan Sci Lett 47:391-397 Jaoul O, Raterron P (1994) High-temperature deformation of diopside crystal 3. Effect of pO2 and SiO2 precipitation. J Geophys Res 99:9423-9439 Karato S, Paterson MS, Fitz Gerald JD (1986) Rheology of synthetic olivine aggregates: influence of grain size and water. J Geophys Res 91:8151-8179 Keefner JW, Mackwell SJ, Kohlstedt DL (2005) Dunite viscosity dependence on oxygen fugacity. Lunar Planet Sci Conf XXXVI:1415 Kirby SH, Raleigh CB (1973) Mechanisms of high-temperature, solid-state flow in minerals and ceramics and their bearing on the creep behavior of the mantle. Tectonophys 19:165-194 Kohlstedt DL, Goetze C (1974) Low-stress high-temperature creep in olivine single crystals. J Geophys Res 79:2045-2051 Kohlstedt DL, Hornack P (1981) Effect of oxygen partial pressure on the creep of olivine. In: Anelasticity in the Earth, Geodynamic Series 4. Stacey FD, Paterson MS, Nicolas A (eds), AGU, Washington DC p 101-107 Kohlstedt DL, Goetz C, Durham WB, VanderSande J (1976) New technique for decorating dislocations in olivine. Science 191:1045-1046 Kollé JJ, Blacic JD (1982) Deformation of single-crystal clinopyroxenes: 1. Mechanical twinning in diopside and hedenbergite. J Geophys Res 87:4019-4034 Kollé JJ, Blacic JD (1983) Deformation of single-crystal clinopyroxenes: 2. Dislocation-controlled flow processes in hedenbergite. J Geophys Res 88:2381-2393 Kröger FA, Vink HJ (1956) Relations between the concentrations of imperfections in crystalline solids. In: Solid State Physics, vol 3. Seitz F, Turnball D (eds) Academic Press, NY p 307-435 Mackwell SJ, Bai Q, Kohlstedt DL (1990) Rheology of olivine and strength of the lithosphere. Geophys Res Lett 17:9-12 Mackwell SJ (1991) High-temperature rheology of enstatite: Implications for creep in the mantle. Geophys Res Lett 18:2027-2030 Mackwell SJ, Dimos D, Kohlstedt DL (1988) Transient creep of olivine: Point defect relaxation times. Phil Mag 57:779-789 McCammon C (2005) Mantle oxidation state and oxygen fugacity: Constraints on mantle chemistry, structure, and dynamics. In: Earth’s Deep Mantle: Structure, Composition, and Evolution. Van der Hilst RD, Bass J, Matas J, Trampert J (eds) AGU, Washington DC, p 221-242 Mei S, Kohlstedt DL (2000) Influence of water on plastic deformation of olivine aggregates. 2. Dislocation creep regime. J Geophys Res 105:21,471-21,481

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Phakey P, Dollinger G, Christie J (1972) Transmission electron microscopy of experimentally deformed olivine crystals. In: Flow and Fracture of Rocks: the Griggs Volume, Geophysical Monograph 16. Heard HC, Borg IY, Carter NL, Raleigh CB (eds) AGU, Washington, DC, p 117-138 Poirier J-P (1985) Creep of Crystals. Cambridge University Press, Cambridge UK Post RL (1977) High-temperature creep of Mt. Burnet dunite. Tectonophys 42:75-110 Raleigh CB (1968) Mechanisms of plastic flow in olivine. J Geophys Res 73:5391-5406 Raleigh CB, Kirby SH (1970) Creep in the upper mantle. Mineral Soc Am Spec Paper 3:113-121 Raleigh CB, Kirby SH, Carter NL, Ave Lallemant HG (1971) Slip and the clinoenstatite transformation as competing rate processes in enstatite. J Geophys Res 76:4011-4022 Raterron P, Jaoul O (1991) High-temperature deformation of diopside single crystal, 1. Mechanical data. J Geophys Res 96:14,277-14,286 Ricoult DL, Kohlstedt DL (1985) Experimental evidence for the effect of chemical environment upon the creep rate of olivine. In: Point Defects in Minerals, Geophysical Monograph 31. AGU, Washington, DC, pp 171-184 Ross JV, Nielsen KC (1978) High-temperature flow of wet polycrystalline enstatite. Tectonophys 44:233-261 Schmalzried H (1981) Solid State Reactions. Verlag Chemie, Weinheim Stocker RL (1978a) Point-defect formation parameters in olivine. Phys Earth Plan Int 17:108-117 Stocker RL (1978b) Influence of oxygen pressure on defect concentrations in olivine with a fixed cationic ratio. Phys Earth Plan Int 17:118-129 Stocker RL, Smyth DM (1978) Effect of enstatite activity and oxygen partial pressure on the point-defect chemistry of olivine. Phys Earth Planet Int 16:145-156 von Mises R (1928) Mechanik der plastischen Formänderung von Kristallen. Z Ang Math Mech 8:161-185 Weertman J (1968) Dislocation creep theory of steady-state creep. Trans Am Soc Metals 61:681-694 Zeuch DH, Green HW II (1984) Experimental deformation of a synthetic dunite high temperature and pressure. I. Mechanical behavior, optical microstructure and deformation mechanism. Tectonophys 110:233-262

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SUBJECT INDEX 1 Ceres ........................................302, 305-306 2 Pallas ............................................... 302, 306 4 Vesta ................................ 277, 302, 313, 381 11 Parthenope..................................... 279, 302

—A— acapulcoites .................................376, 577-579 achondrites ..........................................576-578 basaltic ............................................380-381 bulk......................................................... 348 primitive ................................................. 375 AGB stars ............................................... 40, 45 alteration products...................................... 370 amoeboid olivine aggregates ............. 126, 144, .................................. 156-160, 362, 572 angrites ........................381, 385, 496, 578-579 anorthite ..................................................... 166 aqueous alteration, O-isotopes and ...... 11, 436 aqueous fluid .......................................369-370 assimilation-fractional crystallization (AFC) .. ...................................................515-516 asteroid belt dynamical evolution ............................... 329 resonances .............................................. 275 structure.................................................. 275 asteroid types ............................................. 430 asteroidal water .......................................... 454 asteroids A-type .....................................301, 304-305 C-complex .............................................. 304 D-type .............................................306-307 dynamical evolution ............................... 327 E-type ..................................................... 307 heliocentric distribution ..................314-321 hydrated...........................................321-323 K-type .................................................... 308 L-type ..................................................... 308 M-type .............................................308-309 near-Earth ........................................324-326 O-type .................................................... 310 P-type ..............................................306-307 Q-type .............................................310-311 reflectance ...................................... 277, 280 1529-6466/08/0068-0ind$05.00

R-type..............................................310-311 S-complex ...............................310-312, 318 taxonomy.........................................297-303 T-type ..................................................... 312 V-type (Vestoids).............................312-314 Xe-type................................................... 307 astronomical techniques brightness ............................................... 276 spectroscopy ....................................276-277 mid-IR .............................................278-279 near-IR ................................................... 278 corrections .......................................279-280 atomic isotope shifts .................................... 60 aubrites ................................360, 496, 578-580

—B— banded iron formations .............. 477, 480, 482 basalt lunar ................................................498-499 oxygen fugacity ...............................546-547 big delta........................19, 147, 189, 347-348, ....................................407-408, 415-416 Bowen’s series............................................ 515 brachinites .................................. 377, 578, 580 redox conditions ..................................... 495 buffers .................................................528-529 iron-Ti oxide oxybarometer ............535-537 iron-wüstite (IW) ...................494, 496-499, ................... 501-503, 582-529, 557-563 magnetite-hematite (MH)........528-529, 558 nickel-nickel oxide (NNO).............528-529, ...........................................557-563, 565 quartz-fayalite-magnetite (QFM or FMQ) ... ..................................498, 500-501, 503, ...........................528-529, 531, 537, 547 wüstite-magnetite (WM) ........................ 558

—C— C/O ratio, solar ....................................106-107 Ca-QUIlF model .................................536-537 capital delta .................................................. 20 capital delta prime ........................................ 20 carbon burning ............................................. 35 DOI: 10.2138/rmg.2008.68.ind

592

Subject Index — Oxygen in the Solar System

carbonaceous chondrites bulk................................................. 414, 417 groups..............................................573-575 olivine..................................................... 407 carbonaceous planetesimals ....................... 238 carbonates ...........................147-148, 370-372, ....................................431-432, 436-438 Ca-Tschermak’s molecule .................... 98, 101 CB chondrites..................................... 155, 374 chondrules in ...................................170-171 CCAM (carbonaceous chondrite anhydrous minerals) ................... 7, 348, 361 CH chondrites ............................ 127, 155, 373 chondrules in .......................................... 170 chemical isotope effects ................................. 9 chondrites bulk.................. 146-147, 347-348, 401-404, ....................................409-414, 573-575 groups..............................................572-573 Kakangari-like........................................ 576 moderately-volatile elements ..........415-417 petrologic types ...................................... 574 refractory lithophiles in .......................... 413 refractory siderophiles in ................415-416 secondary phases .............................145-147 separates ..........................................354-355 chondrules ............... 93-94, 124-125, 141-144, ............................168-172, 446-447, 574 Al-rich .............................................355-356 in CH chondrites .................................... 373 in CV3 chondrites ...........................363-364 in ordinary chondrites .....................351-353 CI chondrites ...............................371, 430-434 circumstellar disks ......201, 204, 206, 209-210 chemical evolution ................................. 205 CK chondrites ............................................ 366 clinopyroxene activity coefficients ................................ 102 synthetic ..............................................97-99 CM chondrites CM chondrites.........................368, 432-435 CM2 chondrites...............148, 154, 159-160 refractory inclusions in .................... 109 CNO burning .....................................33-34, 39 CO chondrites ......151, 153, 159-160, 366-368 alteration ................................................ 448 chondrules in ...................................171-173

CO, isotopologues ...................................60-61 comets ...................................63, 237, 247-248 oxygen in.........................................252-258 oxygen isotopes in...................261-262, 264 volatiles in .............................................. 255 continents, oxygen isotopic composition .......... ...................................................523-524 CR chondrites.......153, 159-160, 372-374, 431 alteration .........................................438-439 chondrules in ...................................170-171 creep oxygen fugacity ...............................566-565 olivine aggregates (dunites) ...........558-560, .................................................. 562-563 olivine single crystals ......................556-563 pyroxene aggregates ........................564-565 pyroxene single crystals ..................565-566 cristobalite, in ordinary chondrites ............ 354 CV chondrites .....................154, 158-159, 361 alteration .......... 439-440, 442-444, 446-448 bulk O isotopes....................................... 441 chondrules in .................................. 172, 173

—D— dark inclusions, in CV3 chondrites ... 365, 439, ...........................................441, 445-447 deep ocean .................................................. 483 delta notation.......................... 16, 19, 403, 464 delta-delta plot ....................................517-518 dex units ..................................................74-76 diogenites ............................................380-381 disk evolution ............................................. 239 dust enrichment.......................................120-122 formation .................................................. 67 in the early solar nebula ......................... 117 dusty olivine ............................................... 124

—E— electron spin resonance .............................. 109 enstatite chondrites...............94, 127-135, 151, ...........................................358-360, 576 bulk......................................................... 417 chondrules in .................................. 172, 174 fragments in Kaidun............................... 453

Subject Index — Oxygen in the Solar System Eu oxybarometer ..................500-503, 541-544 Eucrites .......................................380-381, 496 extinction...................................................... 58

—F— fassaite..........................98, 100, 103-104, 109, .......................................... 158, 162, 165 fayalite........................................ 110, 148, 446 condensation of ...............115, 117, 121-122 diffusion into forsterite...........118, 121-122, ...................................................157-158 O isotopic composition .......................... 447 feldspar, reflectance spectra ................288-289 FeNi metal in meteorites ........................................... 286 reflectance spectrum............................... 287 forsterite ............................................. 110, 356 fractional crystallization, and oxygen isotopes...................................... 519 Fremdlinge ................................................. 107 FUN CAIs ...................................................... 8

—G— Galileo Galileo NIMS ..................................226-227 Galileo Probe Mass Spectrometer (GPMS) .. ...........................................223-225, 230 gas-phase molecules, oxygen in ..............59-60 GEMS .................................................259-260 Genesis mission ......................................87-88 glass, lunar ..........................................498-499 granulation ................................................... 75

—H— halides .................................................451-452 HED meteorites...........................578, 580-581 helium burning ..................................32, 34-35 hibonite .............................................. 167, 175 howardites ...........................................380-381 hydrogen burning ....................................32-34 hydrothermal alteration .......................519-521

—I— ice in the early solar nebula ......................... 116

593

in clouds ..................................250-251, 254 in comets ................................................ 253 icy migrators ...............................112, 119-120 icy planetesimals ........................................ 237 interplanetary dust particles (IDPs) .. 247, 251, ...........................................258, 263-264 anhydrous ....................................... 248, 252 chondritic porous ............................258-259 hydrous chondritic ..........................448-450 interstellar dust, oxygen in ..................... 63, 68 interstellar medium (ISM) phases in ........................................56, 64-68 oxygen in.............56, 64, 248-249, 252, 261 iron meteorites ............................581-583, 585 IAB..................................................383-384 IAB, silicate inclusions in ...................... 579 IIAB ....................................................... 384 IIE ...................................................384-385 IIIAB ...................................................... 385 IVA ..................................................386-387 magmatic .........................................583-584 source ..................................................... 330 ungrouped .............................................. 388 iron-Ti oxide oxybarometer ................535-537 iron-wüstite (IW) buffer .............494, 496-499, ................... 501-503, 582-529, 557-563 irradiation, UV ....................................250-251 isotope ratios, notation ................................. 18 isotopic fractionation .....86, 514-516, 520-522 kinetic processes .................................26-27 subsolidus........................................517-518 temperature dependence........................... 25 isotopic fractionation factor ..............19, 25-26 isotopologues ....................17, 59, 61, 199-200 CO .......................................................60-61 isotopomer.................................................... 59

—J— Jovian atmosphere ........................................ 84 Jovian O/H ................................................. 229 Jupiter 5-μm hot spot ..........................223, 225-227 cloud base........................................222-223 cloud structure.................................222-227 lightning ..........................................227-228 oxygen isotopes...................................... 228

594

Subject Index — Oxygen in the Solar System

satellites.................................................. 234 volatile and noble gases ..................230-231 water abundance ..................... 226, 229, 240

—K— Kuiper Belt objects ............................. 235-236

—L— life, origin of .............................................. 257 lodranites ............................................ 376, 579 low-mass stars, evolution of ....................38-39 lunar metals grains, oxygen in ........81-83, 188 lunar samples oxygen isotopes in...........................512-513 redox conditions ..............................497-499

—M — magnetite .............................148, 431-432, 451 magnetite-hematite (MH) buffer ...........................................528-529, 558 mantle, oxidation state ........................528-532 Mars, redox conditions........................499-500 mass-dependent fractionation ...... 24, 142, 209 mass-independent fractionation .142, 187-188, .......................... 192-194, 200, 214, 347 and CAIs .........................................195-196 chemical ..........................................195-197 photochemical ........................................ 198 matrix ..........................145, 175-179, 371, 440 melilite ........................................150, 163-168 synthetic ..............................................97-98 MELTS program .........................534, 539-540 Mercury, redox conditions ..................504-505 mesosiderites ...................................... 387, 586 metallicity .................................................... 61 metal-silicate fractionation................. 402, 403 meteorites bulk oxygen isotopic compositions .............. ...................................................417-418 delivery................................................... 328 mid-ocean ridge basalt (MORB) ........ 512, 547 mineral equilibria, as oxybaroeters .....535-537 mineralogic alteration index....................... 437 molecular clouds ............................... 200, 207, ....................................210-212, 249-250 Moon, redox conditions ............................. 497

—N— nanophase iron ....................................291-292 neon burning ................................................ 35 Neptune ...................................................... 232 nickel-nickel oxide (NNO) buffer ......528-529, ...........................................557-563, 565 nitrogen ................................11-12, 86-87, 213 in lunar soils ............................................. 80 isotopes in SiC ......................................... 11 solar .....................................................83-85 non-mass-dependent fractionation ...........9-10, ................27-28, 400, 405-406, 411, 469 novae ............................................................ 41

—O— ocean island basalt (OIB) ........................... 547 oldhamite.....................................130-131, 134 olivine..................................................109-110 deformation ......................556-559, 561-563 diffusion in ......................................562-564 reflectance spectra ...........................282-284 olivine-pyroxene-spinel oxybarometer ............ ...................................................538-540 ordinary chondrites ............ 109, 151, 575, 576 alteration ................................................ 451 basaltic inclusions in .............................. 382 bulk..................................................349-351 chondrules in .......................... 172, 174, 408 matrix ..................................................... 352 olivine..................................................... 408 outer planet satellites bulk porosity .......................................... 234 Galilean ...........................................233-234 oxygen .................................................... 232 Titan ....................................................... 233 Triton ...................................................... 233 outer planets atmospheres......................220-221, 231-232 formation .........................................236-238 oxidizing environments, nebular .........122-123 oxybarometers .....................................533-535 oxygen abundance of .................. 31, 62, 66, 68, 142 atomic (O I), in ISM............................56-57

Subject Index — Oxygen in the Solar System atomic (O I), in Sun ............................75-76 burning ..................................................... 36 cycles.......................465-466, 472, 482, 484 early Earth .......................................476-477 evolution of isotopes .......................189-191 fugacity .................................................. 528 in organics .......................................251-252 ionization states........................................ 58 isotopes, chemical evolution of .... 41, 42, 44 isotopic anomalies...........................468-469 isotopic composition .............................. 463 mixing .................................................... 470 molecular................................................ 471 nucleosynthesis ...................... 189, 190, 248 reservoirs, terrestrial............................... 467 planetary processing........................475-476 solar abundance ....................................... 66 terrestrial surface environments ......474-475 tropospheric............................................ 472 oxygenation, of terrestrial atmosphere ............. .................................................. 481, 484 ozone ................9, 192-194, 405, 467, 469-473 isotopomers .................................... 193, 194

—P— palisade bodies ........................................... 165 pallasites......................385, 387-389, 585, 586 petrologic buffers ................................528-529 photodissociation ........198-202, 205-206, 208, ...........................210-212, 405, 409, 411 photolysis ........................................... 123, 208 photons ....................................................... 281 photosynthesis ............................................ 480 phyllosilicates .................................... 433, 436 reflectance spectra .................................. 290 planetary carbon ......................................86-87 planetary interiors, redox state ................... 548 polycyclic aromatic hydrocarbons (PAHs) ...... .......................................................... 251 presolar grains analysis of ................................................ 46 oxygen isotopes in.....................9, 45-50, 84 presolar oxides ........................................47-48 presolar silicates ......................................47-48 protoplanetary disk............................. 221, 239 pyroxene, reflectance spectra ......282, 285-286

595

—Q— quartz-fayalite-magnetite (QFM or FMQ) buffer ..................................498, 500-501, 503, ...........................528-529, 531, 537, 547 QUIlF equilibria ......................................... 536

—R— R chondrites ............................................... 576 alteration ................................................ 452 bulk................................................. 356, 357 components ............................................ 357 radial transport processes, nebular ................... .......................................... 111, 112, 119 radial transport, of dust .............................. 202 reflectance spectra feldspars ..........................................288-289 FeNi metal.............................................. 287 Gaussian modeling ................................. 297 HEDs .............................................. 287, 290 mineral chemistry................................... 296 olivine.............................................282- 284 ordinary chondrites ................................ 295 pyroxene ..................................282, 285-287 spinels .............................................288-289 sulfides ............................................288-289 refractory inclusions (CAIs) .............. 362, 572 ages ........................................................ 180 chondrules in .................................. 172, 175 coarse-grained ..........103-106, 144, 161-169 experiments .........................................95-98 fine-grained .....................................149-156 in CH chondrites .................................... 374 in ordinary chondrites ............................ 355 oxygen fugacity of formation....93, 105-106 oxygen isotopes...................................... 405 oxygen isotopic compositions ................ 7, 8 siderophile elements in .......................... 107 silicon isotopes ................................404-405 refractory organics ..................................... 233 resonances, asteroid belt ............................ 275

—S— Saturn ......................................................... 231 satellites...........................................234-235

596

Subject Index — Oxygen in the Solar System

self-shielding ........................10, 187, 198-200, ............................203-204, 207, 209-214 shergottites ..........................................500-503 shielding effects ........................................... 28 silicates, amorphous ............................175-176 silicon burning.............................................. 36 sinoite ..................................................129-134 SNC meteorites .................................. 499, 500 oxygen isotopes...................................... 468 solar C/O ratio .................... 106, 107, 221, 222 solar nebula, oxygen fugacity ...... 93, 105, 106 solar wind ............................................... 73, 87 in lunar soils ......................... 79, 80, 81, 409 space weathering ........................................ 291 and asteroids........................................... 294 and reflectance spectra ....................291-293 experiments .....................................293-294 spinel .................................. 158, 162, 165, 167 reflectance spectra ...........................288-289 Stardust mission ................. 248, 252, 258, 260 O isotopes............................................... 264 stars ...................................................... 65, 67 stellar ejecta, oxygen isotopes in ................. 37 stony iron meteorites .......................... 582, 585 sulfides, reflectance spectra .................288-289 sulfur isotopes .............................................. 23 in terrestrial sediments ....................477-480 Sun light stable elements in........................85-86 oxygen content ..............................73-74, 77 oxygen isotopic composition .............77-78, .............................................. 84, 88, 191 photoshere .....................................74-75, 87 supernovae Type Ia...................................................... 41 Type II ...................................................... 43

oxygen cycles ......................................... 466 oxygen isotopes in...........513, 515, 522-523 redox conditions ..................................... 493 thermochemistry ...................................99-103 titanium, trivalent ............... 105, 106, 108, 109

—U— Uranus ........................................................ 232 ureilites................................378-379, 578, 581 bulk big delta.................................. 417, 419 redox conditions ..................................... 495

—V— vanadium, as oxybarometer ...............544,-546 Venus, redox conditions ............................. 505 Veritas family ..................................... 448, 449 vertical transport processes, nebular ................ .......................................... 112, 113, 120 vestoids ...................................................... 381 vibrational absorption features ................... 288 visual extinction ......................................... 201

—W— Wark-Lovering rims .....108, 161-163, 166-167 water asteroidal ........................................ 290, 291 in the early solar nebula ........ 112, 113, 116, ...................191, 196, 201-202, 207, 212 winonaites .................................. 377, 578, 579 wüstite-magnetite (WM) buffer ................. 558

—X— X-wind model .................................... 108, 203

—T—

—Y—

terrestrial atmosphere, oxygen ...480-481, 483 terrestrial fractionation line ........................ 347

young stellar objects (YSO) ......................... 64 Young-Russell line ..................................... 362

terrestrial planets big delta.................................................. 420 chromium isotopic compositions ....421-422 isotopic compositions............................. 420 metal in................................................... 420

METEORITE INDEX Acfer 094 .....159-160, 171, 173, 175-179, 191 Acfer 214 ................................................... 170 Acfer 287 ................................................... 358 Adelaide ..............................................175-176 Alais .......................................................... 148 ALH 84001 ................................................ 503 ALH 81258 ................................................ 440 ALH 84028 ................................................ 444 ALHA77003 .............................................. 367 ALHA77005 .............................................. 499 ALHA77299 .......................................352-353 ALHA77307 ................175-176, 367, 407-408 Allegan ....................................................... 295 Allende .........................5, 6, 98, 103-104, 148, ..................................154, 157-158, 163, ..................................173, 363-364, 407, ....................................440-441, 444-447 Angra dos Reis ........................................... 285 Barwell ....................................................... 382 Belgica 7904 .......................................368-369 Bencubbin ...........................................374-375 Bholghati .................................................... 381 Bishop Canyon ........................................... 386 Bishunpur ................................................... 126 Bo Xian ...................................................... 351 Bouvante .....................................287, 313-314 Bovedy ....................................................... 354 Bremervorde............................................... 351 Burnwell..............................................349-350 Caldera ....................................................... 381 Carlisle Lakes............................................. 357 Chainpur..............................................352-353 Colony ........................................................ 149 Cumberland Falls ....................................... 358 DaG 476 ..................................................... 501 Dar al Gani 431 .......................................... 366 Deep Springs .............................................. 388 Dhajala ................................................352-353 Dimmitt ...................................................... 382 Divnoe ........................................................ 377 D’Orbigny .................................................. 285 Eagle Station .............................................. 389 EET 87503 ......................................... 287, 313

EET 99402 ................................................. 284 EET 80047 ................................................. 435 EET 83334 ................................................. 435 EETA 79001 ............................................... 501 Efremovka ...................154, 157-158, 163, 445 El Djouf 001............................................... 153 Essebi ......................................................... 148 Farmington ................................................. 354 Fountain Hills............................................. 374 Galim (b) .............................................358-359 Gibeon ........................................................ 386 Girgenti ...................................................... 295 Greenwell Springs...................................... 295 GRO 95577 ................................................ 373 Gujba .......................................................... 374 HH 237 ....................................................... 374 Ibitira .......................................................... 381 Isheyevo ..................................................... 374 Ivuna ...................................147-148, 371, 434 Kaba ...........................148, 440-441, 443, 447 Kaidun .........................................449, 452-453 Kainsaz....................................................... 149 Karoonda .................................................... 366 LAP 91900 ................................................. 287 Leoville ...................................................... 158 LEW 85332 ................................................ 408 LEW 87223 .........................................358-359 LEW 90500 ................................................ 290 LEW 85311 ........................................ 432, 434 LEW 86010 ................................................ 449 Lodran ........................................................ 296 Los Angeles................................................ 501 MAC88107......................... 432, 434, 436, 449 Maltahöhe .................................................. 383 MET00430 ................................................. 440 Mezö-Madaras ....................................352-353 Milton..................................................388-389 Mokoia ....................... 148, 363, 441, 443, 447 Murchison ............ 11, 154, 160, 251, 433, 436 Ngawi ..........................................353-354, 451 Ningqiang................................................... 148 NWA 011 ................................................... 381 NWA 468 ................................................... 384

598

Meteorite Index — Oxygen in the Solar System

NWA 739 ................................................... 373 Odessa ........................................................ 287 Orgeil ..................................................147-148 Orgueil .........................371-372, 407, 431-434 Ornans ........................................................ 149 Parnallee ............................................. 352, 354 Pasamonte .................................................. 381 PAT91456 ................................................... 127 PCA 82500 ................................................. 366 PCA 91241 ................................................. 357 PCA 91461 ..........................................358-359 PCA 9102 ................................................... 452 Pena Blanca Spring .................................... 287 QUE 94201 .........................................500-501 QUE 94411 ................................................ 374 Quinyambie .........................................355-356 Rennazzo .................................................... 148 Rose City.................................................... 350 Rumuruti ............................................ 284, 357 San Joao Nepomuceno ............................... 386 Semarkona...........148, 151, 176, 353-355, 451 Shallowater................................. 358, 360, 449 Shergotty .............................................499-501 Sikhote-Alin ........................................383-384 Sombrerete ................................................. 384 South Oman ........................................358-359 St. Mary’s County ...................................... 351 Steinbach .................................................... 386 Tagish Lake ..................147, 432-433, 437-439 Tieschitz ..................................................... 351 Tucson ........................................................ 388 Vermillion .................................................. 384 Vigarano .......................154, 166-167, 177-178 Warrenton ................................................... 367 Weatherford.........................................374-375 Weston.................................................352-353 Weston 002................................................. 366 Y 74659...................................................... 296 Y 75035...................................................... 377 Y 75097...................................................... 382 Y 75274...................................................... 376 Y 7903........................................................ 366 Y 791198.................................................... 436

Y 791717.................................................... 367 Y 793241.................................................... 382 Y 81020...............149, 152, 156, 367, 407-408 Y 82050...................................................... 367 Y 8451........................................................ 384 Y 86029...................................................... 371 Zagami ................................................499-501

Reviews in Mineralogy & Geochemistry Vol. 68, pp. 571-590, 2008 Copyright © Mineralogical Society of America

Appendix: Meteorites – A Brief Tutorial David W. Mittlefehldt KR/NASA Johnson Space Center Houston, Texas 77058, U.S.A. [email protected]

ABSTRACT There are four broad categories of meteorites—chondrites, achondrites, irons and stony irons. These are subdivided into meteorite groups, the basic unit of meteorite classification. Although no formal guideline is in place for the minimum number of meteorites needed to define a group, common practice is that there should be five or more members in a group. A defined meteorite group is thought to be derived from a single asteroid. However, some groups are genetically related and are derived from a common parent asteroid. Chondrites are primitive stony meteorites; rocks whose compositions are little changed since their formation in the solar nebula. There are fourteen defined groups of chondrites, and they make up the vast majority of meteorites falling to Earth in the current epoch. Achondrites are stony meteorites of two broad types. Some are primitive materials like chondrites, but most are the products of igneous differentiation. There are ten defined groups of achondrites, of which seven are differentiated types. Irons are also the products of asteroidal differentiation, having crystallized from metallic melts separated from chondritic precursors. There are thirteen defined groups of iron meteorites. Stony-irons are also differentiated materials, and both the rocky and metallic phases were formed by igneous processes on asteroids. There are two defined groups of stonyiron meteorites. In addition to those that fit into groups, there are many meteorites that are unique, or for which there are less than five examples. These ungrouped meteorites make up a substantial fraction of meteorites recovered to date.

INTRODUCTION The purpose of this chapter is to present the basic nomenclature and classification scheme most commonly used in the meteoritical literature as a framework for understanding meteoriterelated chapters in this book. It is not the purpose of this chapter to discuss meteorite groups and their origins in detail. The vast majority of meteorites are debris from the smaller bodies in the Solar System—the asteroids. To date, no meteorites have properties that strongly indicate cometary origin (Campins and Swindle 1998). Some meteorites are derived from the Moon and Mars, but discussion of these more properly belongs in the chapters of this book discussing the terrestrial planets. Herein only the classification of asteroidal materials is discussed. One key component of the scientific enterprise is development of a useful and robust classification system for the objects of study. Meteorites are not an exception. A well developed classification system based on chemical composition, mineralogy and texture is currently in use. A brief overview of meteorite classification is given here; detailed discussion of the topic is found in Krot et al. (2003). Meteorites can be divided into four broad categories: chondrites; achondrites; irons; and stony irons. Chondrites are primitive stony meteorites; that is, they are rocks whose compositions have remained virtually unchanged over the history of the Solar System. They have suffered differing degrees of textural and mineralogical modification by 1529-6466/08/0068-0app$05.00

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thermal metamorphism, aqueous alteration and/or shock metamorphism. Achondrites are also stony meteorites, and some are also primitive materials. However, most achondrites are differentiated materials—rocks that have been transmogrified by asteroidal igneous processes. Irons are perhaps the most familiar type of meteorite because large specimens are often prominent exhibits in museums and planetariums. They are also differentiated materials; all of the major iron meteorite groups crystallized from metallic melts. Finally, the stony-irons too are differentiated materials; both the silicate and metallic phases in them were formed by igneous processes. Within each major meteorite type, such as the chondrites, the basic unit of classification is the “group.” Many groups are defined based mostly on bulk chemical characteristics, but especially among the achondrites and stony-irons, texture, mineralogy and mineral compositions are important criteria. It may surprise the reader that, despite the importance of having a robust classification system in science, there is no formal set of definitions for all the different meteorite groups. Similarly, there is no definition for the minimum number of meteorites that constitutes a group. Wasson (1974) recommended that ≥5 meteorites should be used, and many researchers follow this. Not all use this rule, however, and indeed, the angrite achondrite “group” was originally defined based on a single meteorite. The general presumption is that a meteorite group is derived from a single asteroid. In some cases among achondrites and stony-irons, two or more groups seem closely related, and may have been derived from a common parent asteroid.

CHONDRITES Chondrites are the most numerous of the meteorites falling on Earth in the current epoch. Roughly 82% of observed falls are chondrites (Grady 2000). Most chondrites are placed in one of three broad classes: the carbonaceous; ordinary; and enstatite classes. Two chondrite groups are unassociated with any of the classes. Table 1 lists the different groups of chondrites. For a listing of ungrouped chondrites, see Krot et al. (2003). Chondrites are composed of four distinct petrographic components that are found in most or all chondrite types: chondrules; refractory inclusions; matrix; and metal plus sulfide (e.g., Brearley and Jones 1998; Krot et al. 2003; Scott and Krot 2003). Chondrules are spherical, submillimeter-sized bodies dominantly composed of silicates that were partially or completely molten as free objects in space. Two distinct types of materials compose the refractory component: calcium-, aluminum-rich inclusions, and amoeboid olivine aggregates. Calcium-, aluminum-rich inclusions (CAIs) are bodies up to ~1 cm in size composed of the oxide and silicate minerals predicted to condense from a cooling gas of solar composition at the highest temperatures. Amoeboid olivine aggregates (AOAs) are irregularly shaped agglomerations dominated by fine-grained magnesian olivine and containing Fe,Ni metal and refractory silicates. Matrix is a fine-grained mixture of silicates, sulfides, metal, oxides and/or carbonates. The mineralogy of matrix varies with chondrite type (e.g., Scott and Krot 2003), but matrix is always enriched in the volatile components. Table 1 gives the rough proportions of different textural units in chondrite groups. Chondrules and CAIs have been the subjects of major research efforts, and extensive bodies of literature are devoted to them. The CAIs were early recognized to be composed of minerals predicted to condense at high temperatures from the solar nebula (Grossman 1972). As such, their compositions reflect only a small fraction of the condensable matter of the solar nebula. Subsequent work demonstrated that they carry oxygen with anomalous isotopic compositions, have isotopic anomalies in other elements, and are the oldest objects of the Solar System (see MacPherson 2003). There are distinct types of CAIs that reflect different types of processing in the solar nebula. Detailed reviews of them are in Brearley and Jones (1998) and MacPherson (2003). Chondrules are igneous-textured objects that were formed

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