Nano and Micro Diamond Formation in Nature: Ultrafine Carbon Particles on Earth and Space (SpringerBriefs in Earth Sciences) 3031432770, 9783031432774

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SpringerBriefs in Earth Sciences Sergei Simakov · Vittorio Scribano · Nikolai Melnik · Victor Pechnikov · Irina Drozdova · Vladimir Vyalov · Mikhail Novikov

Nano and Micro Diamond Formation in Nature Ultrafine Carbon Particles on Earth and Space

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Sergei Simakov · Vittorio Scribano · Nikolai Melnik · Victor Pechnikov · Irina Drozdova · Vladimir Vyalov · Mikhail Novikov

Nano and Micro Diamond Formation in Nature Ultrafine Carbon Particles on Earth and Space

Sergei Simakov LLC Adamant St. Petersburg, Russia Nikolai Melnik Lebedev Physics Institute Russian Academy of Sciences Moscow, Russia Irina Drozdova Grebenshchikov Institute of Silicate Chemistry Russian Academy of Sciences St. Petersburg, Russia Mikhail Novikov Institute of Experimental Mineralogy Russian Academy of Sciences Chernogolovka, Russia

Vittorio Scribano Department of Biological Geological and Environmental Sciences University of Catania Catania, Italy Victor Pechnikov Central Research Institute of Geological Prospecting for Base and Precious Metals (TsNIGRI) Moscow, Russia Vladimir Vyalov A. P. Karpinsky Russian Geological Research Institute St. Petersburg, Russia

ISSN 2191-5369 ISSN 2191-5377 (electronic) SpringerBriefs in Earth Sciences ISBN 978-3-031-43277-4 ISBN 978-3-031-43278-1 (eBook) https://doi.org/10.1007/978-3-031-43278-1 © The Editor(s) (if applicable) and The Author(s), under exclusive license to Springer Nature Switzerland AG 2023 This work is subject to copyright. All rights are solely and exclusively licensed by the Publisher, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilms or in any other physical way, and transmission or information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed. The use of general descriptive names, registered names, trademarks, service marks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. The publisher, the authors, and the editors are safe to assume that the advice and information in this book are believed to be true and accurate at the date of publication. Neither the publisher nor the authors or the editors give a warranty, expressed or implied, with respect to the material contained herein or for any errors or omissions that may have been made. The publisher remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. This Springer imprint is published by the registered company Springer Nature Switzerland AG The registered company address is: Gewerbestrasse 11, 6330 Cham, Switzerland Paper in this product is recyclable.

Contents

1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1 5

2 Diamond Thermodynamic Stability: The Paradox of Crystal Size . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

9 16

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

19 27

4 Diamond Formation in the Oceanic Lithosphere . . . . . . . . . . . . . . . . . . 4.1 Nano- and Micron-Sized Diamond Formation in the Oceanic Lithosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.1 Nanodiamond and Organic Compound Formation in Oceanic Serpentinite Systems . . . . . . . . . . . . . . . . . . . . . . . 4.1.2 General Notes on the Serpentinization of the Oceanic Lithosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.3 Nanodiamonds in Serpentinites—A Case Study from Sicily and Other Occurrences . . . . . . . . . . . . . . . . . . . . . 4.1.4 Nanodiamonds in Serpentinites as Proxies for the Emergence of Life on Early Earth . . . . . . . . . . . . . . . . 4.1.5 Micron-Sized Diamonds in Ophiolites—Insights into Their Formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.6 Nano- and Micron-Sized Diamond Formation in Hawaiian Salt Lake Crater Xenoliths . . . . . . . . . . . . . . . . . 4.1.7 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2 Carbonado Genesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

29 30 30 30 33 39 40 49 50 50 51

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Contents

5 Nanocarbon and Microdiamond Formation in the Lithogenesis and Metamorphic Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1 Nanocarbon Formation in the Lithogenesis and Contact Metamorphism Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1.1 Nanodiamond Formation in the Lithogenesis and Contact Metamorphism Processes . . . . . . . . . . . . . . . . . . 5.1.2 Fullerene and Fullerene-Like Phase Formation During Metamorphic Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.1.3 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2 Micron-Sized Diamond Formation During Metamorphic Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.1 Worldwide Metamorphic Diamonds . . . . . . . . . . . . . . . . . . . . 5.2.2 Kokchetav Diamond Deposit . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.3 Kokchetav Diamond-Bearing Crustal Rocks—An Overview . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.4 Mineralogical Features of Kokchetav Metamorphic Diamonds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2.5 Processes of Kokchetav Microdiamond Formation . . . . . . . . 5.2.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 Diamonds in Kimberlites and Their Xenoliths: A Reappraisal . . . . . . 6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.2 Diamond Inclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.3 Diamond Formation from Fluids in the Upper Mantle . . . . . . . . . . . . 6.4 Diamond Formation in the Postmagmatic Processes of Kimberlites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5 The Genesis of the Extralarge Type IIa Diamonds . . . . . . . . . . . . . . . 6.6 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

61 61 61 67 69 70 70 71 78 79 83 87 88 95 95 96 99 104 109 118 119

7 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 129 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132

Abbreviations

Aq Antg Coe CVD D DND Dol En Fa Fo FT-t FTIR G Gs hkl HPHT HRTEM ICDD IW KFS Liq Mgt MH Msn Muas NASA ND NNO OCCs Ol OM

Water Solution Antigirite Coesite SiO2 Chemical Vapor Deposition Diamond, C Detonation Nanodiamond Syntheses Dolomite CaMg(CO3 )2 Enstatite MgSiO3 Fayalite Fe2 SiO4 Forsterite Mg2 SiO4 Fischer–Tropsch-type Fourier Transform Infrared Spectra analysis Graphite, C Gaseous Diffraction Indexes High Pressure, High Temperature High-Resolution Transmission Electron Microscope International Center for Diffraction Data Iron–Wüstite Buffer K Feldspar Liquid, Melt Magnetite Fe3 O4 Magnetite-Hematite Buffer Magnesium Silicon Nitride MgSiN2 Muassonite SiC National Aeronautics and Space Administration Nanodiamond Nickel–Bunsenite Buffer Oceanic Core Complexes Olivine(Mg, Fe)2 SiO4 Organic Matter vii

viii

Opx Per Pyr QFM Ro S SIMS Srp SSZs TEM SuR U–Pb UHP UHPT UV WM

Abbreviations

Orthopyroxene (Mg, Fe)SiO3 Periclase MgO Pyrite FeS2 Quartz–Fayalite–Magnetite Buffer vitrinite reflectance Solid Secondary Ion Mass Spectrometry analysis Serpentine Mg2 Si2 O5 (OH)4 Suprasubduction Zones Transmission Electron Microscope Super Reduced Uranium-Plumbum Ultra High Pressure Ultra High Pressure Temperature Ultraviolet Wustite–Magnetite Buffer

Physical Symbols Å bar erg f G h3 kJ kV nm μm P ppm R T U θ

Angstrom Unit of Pressure Unit of Energy Fugacity Gibbs energy Volume of carbon atom in diamond kilojoules kilovolts Nanometer Micron Pressure Parts per million Gaseous Constant Temperature Work of Nuclei Forma Diffraction angle

Chapter 1

Introduction

Carbon is a many-sided chemical element, as it forms millions of compounds related to both biochemical and geochemical processes. Moreover, elemental carbon occurs in various polymorphs (or allotropes), such as graphite, diamond, amorphous carbon, lonsdaleite, and fullerenes. The physical properties of the different carbon polymorphs vary widely due to the different ways in which the atoms in each are bonded. Diamond is the most compact, Sp3 -bonded, polymorph of carbon, having nearly twice the density of graphite. The study of diamond has seen a recent burst of activity in geochemistry and astrophysics, in novel methods of synthesis, and in the development of useful applications. Diamond is recognized as an extraordinary recorder of astrophysical and geodynamic events that extend from the most remote regions of space to Earth’s deep interior. As will be specified later, different types of diamonds have been recognized based on their size, geological occurrence, morphological characteristics, types of solid or fluid inclusions, etc. The processes of formation of some diamond types still raise many contentious questions. The formation of macroscopic diamonds is mainly connected with deep-seated igneous rocks, such as kimberlites and lamproites, and therein mantle xenoliths. To date, in mineralogy, there is no widely accepted definition of a size boundary between kimberlitic microdiamond and macrodiamond. It varies from 0.5 to 1 mm (Chapman and Boxer 2004; Pattison and Levinson 1995). We have assumed 1 mm as the size limit between the microdiamond and macrodiamond. Interestingly, microsized diamonds were also found in metamorphic rocks and ophiolite complexes originating in the graphite stability field. The formation mechanism of these metamorphic rock-hosted diamonds still represents a hotly debated topic. Currently, the origin of natural nanocarbon particles is still poorly constrained. The family of “nanocarbons” includes the following main components: nanodiamonds, nanosized amorphous carbon, fullerenes, diamondoids, graphene, tubes, onions, horns, rods, cones, peapods, bells, whiskers, platelets, and foam (Shenderova et al. 2002). Diamond structures at the nanoscale (length ~1–100 nm) include purephase diamond films, diamond particles, recently fabricated 1-D diamond nanorods,

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2023 S. Simakov et al., Nano and Micro Diamond Formation in Nature, SpringerBriefs in Earth Sciences, https://doi.org/10.1007/978-3-031-43278-1_1

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2

1 Introduction

and 2-D diamond nanoplatelets. There is a special class of nanodiamond material often called “ultrananocrystalline” diamond with the characteristic size of the basic diamond constituents encompassing the range of just a few nanometers that distinguishes it from other diamond-based nanostructures with characteristic sizes below~10 nm (Schrand et al. 2009). The smallest nanometer-sized diamond particles are known mainly in space objects (Tielens et al. 1987; Allamandola et al. 1993). Nanodiamonds in terrestrial rocks are rarely found. They are known in serpentinites, Younger Dryas Boundary layers, coals, and spinel lherzolites (Dubinchuk et al. 1976; Wirth and Rocholl 2003; Tian et al. 2011; Kinzie et al. 2014; Simakov et al. 2015, 2018). Currently, the creation of new methods of nanocarbon synthesis is one of the urgent tasks of modern technologies. There are three main methods of diamond synthesis: 1. HPHT synthesis of macroscopic diamonds from metal and melts of solvent catalysts at P–T corresponding to the thermodynamic field of diamond stability. 2. DND syntheses at ultrahigh temperatures and pressures. 3. CVD synthesis of diamond films, composites, and coatings at high temperatures and low pressures, corresponding to the field of graphite stability. In addition, there are rare nano- and micron-sized diamond syntheses at low P– T parameters. Nakano et al. (2002) synthesized nanodiamond particles of 1 nm in size from an interstellar-like organic mixture with water at 150–400 °C at relatively low pressure. Ultrananocrystalline diamonds were synthesized at 220 ° C and the saturated vapor pressure of water using a simple and available hydrocarbon (glucose) (Alzahrani and Alkahtani 2023). Nano- and micron-sized diamonds (up to 1 μm in size) were synthesized by heating a mixture of Li2 CO3 and nanocarbon particles at 420–550 °C and ambient pressure (Kamali and Fray 2015). Micron-sized diamonds (up to 100 μm) were also synthesized via the interaction between molten aluminum and carbide-containing halide melt at 700–750 °C and ambient pressure (Yolshina et al. 2015). Ishimaru et al. (2001) detected nanodiamond structures in wood charcoal carbonized at 700 °C. In this respect, the book presents the results of theoretical and experimental studies in Chaps. 2 and 3 for a new approach that could be developed for the synthesis of nanodiamonds at low pressures and temperatures. The first detailed studies on synthetic nanodiamonds were carried out in the 1960s in Russia. The particles of synthetic nanodiamonds contain many impurities and defects; therefore, their density is lower than that of diamond and corresponds to the range of 2.8–3.1 g/cm3 . Nitrogen and ketone, hydrocarbon, carboxyl, and alcohol groups with Sp3 and Sp2 hybridizations are present in the impurities of nanodiamonds (Fig. 1.1). It is known that synthesized nanodiamonds display surface bonds terminated with hydrogen and oxygen atoms (Costa et al. 2014; Schrand et al. 2009). Badziag et al. (1990) suggested that nanodiamonds could be synthesized at low P–T from hydrocarbons with H/C < 0.25. Later, Dahl et al. (2003) showed that it is possible to distinguish diamond crystal cages in diamondoids (Sp3 -bonded hydrocarbons), and in this respect, diamond could be defined as the archetypal “macroscopic molecule.” Diamondoids occur in crude oils and in gas condensates in remarkable

1 Introduction

3

Fig. 1.1 Structures in nanodiamonds in accordance with Schrand et al. (2009) and Costa et al. (2014)

amounts (from 35 to 2,075 ppm, respectively; Nekhaev et al. 2010). They are also known in coals and sediments in the middle and final stages of sediment lithogenesis (Tissot and Welte 1978) with vitrinite reflectances from 1 to 4 R° (Wei et al. 2006, 2007). Dahl et al. (2003) extracted diamond molecules from oil diamondoids containing diamond lattices within their tetrahedral structures and suggested that hydrogen-terminated diamonds and nanometer-sized diamondoid hydrocarbons form a continuous structural series that includes small diamondoids ( Pid (when rg > rd ). In accordance with Nuth (1987a, b), the surface energy of diamond with a diameter of 5 nm lies in the range from 3700 to 9800 erg/cm2 , while the surface energy of graphite of a similar size corresponds to the interval from 2000 to 4400 erg/cm2 . If σg = 2000 erg/cm2 , then σd = 4000 erg/cm2 , and if σg = 4000 erg/cm2 , then σd = 5500 erg/cm2 . Using these values and the corresponding atomic volume of carbon (Vg = 7.28 and Vd = 5.67), we obtained Vgσ g/Vdσ d ratios close to 1 (0.93 and 1.07). It follows that for nanoparticles, Vg σg /Vd σd ≈ 1. Thus, the rg / rd ratio can be expressed as: rg /rd = Δμd /Δμg

(2.9)

The difference between Δμg and Δμd depends upon the difference in Pi/Pig and Pi/Pid , and at lower Picar , the difference tends to zero, which corresponds to the optimal condition of diamond formation from a gaseous mixture within the range of graphite stability. In a fluid hydrocarbon–hydrogen mixture (H–C gaseous system), the gas–solid reaction of hydrocarbon shedding can be proposed for carbon formation: CH4 → C + 2H2

(I)

In the cases where hydrogen is much more abundant than methane, the equilibrium values of Pig and Pid increase. In that case, the difference between PCH4 /PCH4(g) and PCH4 /PCH4(d) will be decreased due to small values in the numerator and large values in the denominator. This gives rise to the stabilization of diamond nuclei, hence conforming to the well-known fact that diamond is more stable in a hydrogen medium than graphite (Fedoseev et al. 1984). Gebbiea et al. (2018) showed that the formation of metastable, hydrogen-terminated diamondoid clusters during CVD synthesis leads to a nucleation barrier in 26 carbon atoms instead of several thousand atoms. On the C–H–O diagram, this region of diamond stability is close to hydrogen (Fig. 2.3). On the other hand, Simakov (1995, 2010, 2011) has shown that the addition of oxygen to hydrocarbon gases can stabilize diamond nuclei at the P–T of graphite stability. This coincides with the yearly established circumstance that the oxygen reduction of graphite is greater than that of diamond, which is why diamond is more stable in the oxygen environment than graphite (Kawato and Kondo 1987). Later, Costa et al. (2014) confirmed that nanodiamonds terminated by oxygen functional groups are much more stable than graphite. In a fluid hydrocarbon–hydrogen–oxygen mixture (C–H–O gaseous system), the following gas–solid reaction can be proposed in addition to reaction (I):

2 Diamond Thermodynamic Stability: The Paradox of Crystal Size

13

Fig. 2.3 Atomic C–O–H diagram with a half-parabolic nanodiamond domain. Solid lines: (I) fields of the CVD synthesis, (II) field of the hydrothermal syntheses; pointed lines—boundary of Bachman et al. (1991) experimental point field. A composition of ice mixture extracted from carbonaceous chondrites by Nakano et al. (2002)

CO2 → C + O2

(II)

CO → C + O.5O2

(III)

H2 O → H2 + O.5O2

(IV)

The upper limit of carbon stability in terms of oxygen fugacity corresponds to reactions II and III (CCO buffer). Above CCO, free carbon transfers to CO2 and CO (Jakobsson and Oskarsson 1994). In the range of the CCO buffer, PCH4(car) is low, and PCH4,g ≈PCH4,d , which corresponds to the similarity of the diamond and graphite critical radii (rg ≈ rd , see Eqs. 2.7 and 2.8) and diamond nucleus stabilization (Fig. 2.4). At lower oxygen fugacity, the diamond nucleus is unstable (rg < rd ). From the P–T- f O2 calculations (Fig. 2.4), it follows that there are three types of

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2 Diamond Thermodynamic Stability: The Paradox of Crystal Size

Fig. 2.4 Variation in the composition of the C–O–H system: a at 1000 °C and 10–3 bar, b at 250 °C and 10–3 bar, c at 500 °C and 1 kbar, and d at 1000 °C and 20 kbar. CCO denotes the upper limit of carbon by oxygen. The dotted line denotes the boundary of the range of diamond preferable growth

gaseous mixtures of diamond nucleus stabilization: CO2 –H2 O–CH4 , CO2 –H2 O–H2 , and CO–H2 (Fig. 2.5). CO–H2 compositions correspond to CVD syntheses, and CO2 –H2 O–CH4 (H2 ) compositions correspond to hydrothermal syntheses (Fig. 2.5). On the C–O–H diagram, they form a half-parabolic domain that extends between the CO and CO2 compositions and the acetone-ethanol line of the diagram (Fig. 2.3). Increasing temperature and decreasing pressure lead to the transition of the CO2 –H2 O composition to CO–H2 (Figs. 2.4 and 2.5), which coincides with the relationships established previously for the C–H–O system by Eaton and Sunkara (2000) and Gogotsi et al. (1998). These relationships also explain the phenomenon that the addition of hydrocarbons to pure CO2 first initiates diamond growth and then graphite growth Marinelli et al. (1994) (see Fig. 2.4a, b).

2 Diamond Thermodynamic Stability: The Paradox of Crystal Size

15

Fig. 2.5 Variation in the composition of the C-O–H system with pressure and temperature (I − CO + H2 ; IIa − CO + H2 O + H2 ; IIb − CO + H2 O + CH4 )

Sommer et al. (1989) and Wang et at. (1992) developed a thermodynamic analysis of diamond CVD based on a quasiequilibrium model, where nonequilibrium steady-state depositions of diamond and graphite were analyzed using equilibrium thermodynamics. It follows that our thermodynamic treatment could provide some guidelines for CVD syntheses. On the C–O–H diagram, the CVD region lies on the CO line and below and is limited by the CO composition and the line XC/(C-O) ≈ 0.41 on the C-O side of the diagram (Fig. 2.3). It covers most of the Bachman et al. (1991) and Marinelli et al. (1994) experimental points lying on the CO line and below it. The hydrothermal part lies beyond the CVD region and is limited by the CO2 composition on the C–O side of the C–O–H diagram (Fig. 2.3). This field corresponds to the hydrothermal CO2 –H2 O–H2 (CH4 ) compositions over a wide range of P–T conditions (Fig. 2.5). This explains the relationship of interstellar diamonds with water. The fluid calculations performed at 250 °C and 10–3 bar show that the diamond stability range here corresponds to the CO2 and H2 O compositions of the fluid (Fig. 2.4b). The composition of the ice mixture extracted from carbonaceous chondrites during the laboratory experiments of Nakano et al. (2002) on the C–H– O diagram corresponds to the hydrothermal field of metastable diamond formation (Fig. 2.3). The compositions of glucose-water solutions from experiment of Alzahrani and Alkahtani (2023) also correspond to this field. The aforementioned data show that the presented thermodynamic models of diamond formation from C–H and C–H–O fluid systems provide a common basis for diamond CVD and hydrothermal syntheses. It follows that nanodiamond particles could be formed from carbon-bearing fluids at low P–T parameters without seeds in the range of the upper limits of carbon stability by oxygen and hydrogen. It is also known that supercritical aqueous fluids are favorable media for diamond growth (Capelli 1995; Roy et al. 1996). Supercritical conditions (e.g., T > 375 °C and P > 0.22 kbar) favor the formation of intermediate reaction products, which are necessary for the conversion of carbon to diamond. They are CH3 , CH2 , CH3 OH, CH2 OH, HO2 , H2 O2 , H, and OH (Antal et al. 1987). In accordance with Wehley and Tester (1989), CH3 is believed to be the dominant growth precursor in the synthesis of diamond. The CO2 entrance into supercritical water results in CH3 formation.

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2 Diamond Thermodynamic Stability: The Paradox of Crystal Size

There is a specific distribution of carbon isotopes between graphite and diamond during metastable diamond growth from the gaseous phase. Diamond is more enriched in heavy and larger 13 C, while graphite is more enriched in lighter and smaller 12 C (Galimov et al. 1973; Fedoseev et al. 1971). The distribution of carbon isotopes leads to changes in carbon densities and molar volumes and, as a result, to changes in Pid and Pig connected with diamond stabilization. 13 C is larger than 12 C, which also leads to dislocation and defect formation, which affects the diamond surface energy (σ). From the low-pressure experiments of Galimov et al. (1973), it follows that the introduction of larger 13 C could decrease σ because the value of ΔGP at vacuum pressure is minor and cannot influence diamond stabilization. Nitrogen is the main admixture in synthetic and natural diamonds. In synthetic diamonds, the nitrogen content is variable from 0.3% in HPHT macrodiamonds (Borzdov et al. 2002) to 3% for nanosized detonation and CVD diamonds (Turner et al., 2009; Vlasov et al., 2010; Zhang et al., 1999). From the aforementioned result, it also follows that the introduction of nitrogen could stabilize diamond nuclei formation because nitrogen has a larger size than 13 C, which should lead to diamond nuclei stabilization at low pressures.

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Deryagin BV, Fedoseev DV (1977) Growth of diamond and graphite from the gas phase. Nauka, Moscow Eaton SC, Sunkara MK (2000) Construction of a new C–H–O ternary diagram for diamond deposition from the vapour phase. Diam Relat Mater 9:1320–1326 Fedoseev DV, Galimov EM, Varnin VP et al (1971) Carbon isotope fractionation in the process of physic-chemical synthesis of diamond. Dokl Akad Nauk SSSR 201:1149–1151 Fedoseev DV, Deryagin BV, Varshavskaya IG et al (1984) Diamond crystallization. Nauka, Moscow Galimov EM, Prohorov VS, Fedoseev DV et al (1973) Heterogeneous carbon isotope effects at the synthesis of graphite and diamond from gas. Geokhimiya 17:416–425 Gamarnik MY (1996) Energetical preference of diamond nanoparticles. Phys Rev B: Condens Matter 54:2150–2156 Gebbiea MA, Ishiwatab H, McQuadea PJ et al (2018) Experimental measurement of the diamond nucleation landscape reveals classical and nonclassical features. PNAS 115:8284–8289 Gogotsi Y, Kraft T, Nickel KG et al (1998) Hydrothermal behavior of diamond. Diam Relat Mater 7:1459–1465 Jakobsson S, Oskarsson N (1994) The system C–O in equilibrium with graphite at high pressure and temperature: an experimental study. Geoch Cosm Acta 58:9–17 Jiang Q, Li JC, Wilde G (2000) The size dependence of the diamond-graphite transition. J Phys: Cond Matt 12:5623–5627 Hirano S, Shimono K, Naka S (1982) Diamond formation from glassy carbon under high pressure and temperature conditions. J Mater Sci 17:1856–1862 Hong SM, Akashi M, Yamaoka S (1999) Nucleation of diamond in the system of carbon and water under very high pressureand temperature. J Cryst Growth 200:326–328 Ishimaru K, Vystavel T, Bronsveld P et al (2001) Diamond and pore structure observed in wood charcoal. J Wood Scie 47(5):414–416 Kamali AR, Fray DJ (2015) Preparation of nanodiamonds from carbon nanoparticles at atmospheric pressure. Chem Commun 51:5594–5597 Kawato T, Kondo K (1987) Effects of oxygen on CVD diamond synthesis. Jpn J Appl Phys 26:1429– 1432 Kumar MDS, Akashi M, Yamaoka S (2000) Formation of diamond from supercritical H2 O + CO2 fluid. at high pressure and temperature. J Cryst Growth 213:326–328 Lewis RS, Ming T, Wacker JF et al (1987) Interstella diamonds in meteorites. Nature 326:160–162 Manuella FC (2013) Can nanodiamonds grow in serpentinite-hosted hydrothermal systems? a theoretical modelling study? Mineral Mag 77(8):3163–3174 Marinelli M, Milani E, Montuori M et al (1994) Compositional and spectroscopic study of the growth of diamond films from several gaseous mixtures. J Appl Phys 76:5702–5705 Nakano H, Kouchi A, Arakawa M (2002) Alteration of interstellar organic materials in meteorites’ parent bodies: a novel route in diamond formation. Proc Jp Acad 78B:277–281 Nuth JA (1987a) Small-particle physics and interstellar diamonds. Nature 329:589 Nuth JA (1987b) Are small diamonds thermodynamically stable in interstellar medium? Astroph Space Scie 139:103–109 Onodera A, Irie Y, Higachi K et al (1991) Graphitization of amoprphous carbon at high pressures to 15 GPA. J Appl Phys 69:2611–2614 Pal’yanovYuN, Sokol AG, et al (1999) Diamond formation from mantle carbonate fluids. Nature 400(6743):417–418 Roy R, Ravichandran D, Ravindranathan P et al (1996) Evidence for hydrothermal growth of diamond in the C–H–O and C–H–O halogen system. Mater Res 11:1164–1168 Sato K, Katsura T (2001) Sulfur: a new solvent-catalyst for diamond synthesis under high-pressure and high-temperature conditions. J Cryst Growth 223:189–194 Shul’zenko AA, Getman AF (1971) Verhafrenzursynthese von diamanten. Germ Pat 2032103 Simakov SK (1995) Thermodynamic estimation of oxygen-hydrogen conditions influence on diamond and graphite critical nucleus formation at processes of methane destruction at low pressures. Russ J Phys Chem 69:346–247

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2 Diamond Thermodynamic Stability: The Paradox of Crystal Size

Simakov SK (2010) Metastable nanosized diamond formation from C–H–O fluid system. J Mater Res 25(12):2336–2340 Simakov SK (2011) Nanodiamond formation in natural processes from fluid systems at low P-T parameters. Dokl Earth Scie 436(1):148–151 Sommer M, Mui K, Smith FW (1989) Thermodynamic analysis of the chemical vapour deposition of diamond films. Solid State Commun 69:775–778 Turner S, Lebedev OI, Shenderova O et al (2009) Determination of size, morphology, and nitrogen impurity location in treated detonation nanodiamond by transmission electron microscopy. Adv Funct Mater 19:2116–2124 Vlasov II, Shenderova O, Turner S et al (2010) Nitrogen and luminescent nitrogen-vacancy defects in detonation nanodiamond. Small 6:687–694 Wang RB, Sommer M, Smith FM (1992) The deposition of diamond films via the oxyacetelene torch: Experimental results and thermodynamic predictions. J Cryst Growth 119:271–280 Wehley PA, Tester JW (1989) Fundamental kinetics of methanol oxidation in supercritical water. Amer Chem Soc Symp Ser 406:259–275 Yolshina LA, Muradymov RV, Vovkotrub EG et al (2015) Diamond synthesis in aluminium matrix in molten alkali-halide at ambient pressure. DiamRel Mat 55(1):1–11 Zhang GF, Geng DS, Yang ZJ (1999) High nitrogen amounts incorporated diamond films deposited by the addition of nitrogen in a hot-filament CVD system. Surf Coat Technol 122:268–272

Chapter 3

Experimental Data on Nanocarbon Formation Under Low P–T Conditions

The synthesis of nanosized diamond and diamond-like phases from organics at hydrothermal temperatures and pressures has been experimentally proven (Simakov et al. 2008, 2010). Synthetic nanodiamonds contain many impurities, such as ketones, hydrocarbons, and alcohols, with Sp3 hybridization (see Chap.1, Fig. 1.1. Introd). The aforementioned impurities could be potential sources for ND syntheses because their carbon–hydrogen bond energies are much lower than the energy of the carbon– carbon bond in graphite. The compositions of water-alcohol and hydrocarbon solutions, as shown in the C-H–O diagram (Fig. 3.1), closely approach the hydrothermal nanodiamond domain. Therefore, ethyl alcohol was used as the carbon source for ND synthesis. Nanodiamond synthesis was obtained at 500 °C and a pressure of nearly 1 kbar in a mixture of water–alcohol solution and nitrogen-bearing organic matter. The mixture was heated in a high-pressure reactor (BT-8 alloy) with a total volume of 50 cm3 . The run duration was 5–7 days. After treatment in concentrated HCl to remove possible metallic phases and iron oxides, experimental products were studied by X-ray diffraction on a D/max-RC diffractometer (Rigaku, Japan). Reflections were obtained with CuKα radiation 2θ = 0–100°, step 0.01°, and scan rate 1°/min), and hence, we identified graphite. Thereafter, the products were treated in concentrated HClO4 to remove graphite. The product left over consisted of white and opaque particles of unknown nature. Raman observations were then performed to identify the aforementioned white and opaque particles. Raman and photoluminescence (PL) spectra were recorded on a U-1000 spectrometer with a microscope for the micro-Raman study setup. Argon laser radiation (488 nm) was used for excitation. The laser exciting radiation was focused on a spot ~20–40 μm in diameter. This allowed us to study the homogeneous and transition areas of the sample surface and monitor any laser heating effect on the sample. The Raman spectra were recorded with a resolution of 1–5 cm−1 . The obtained spectra are compared with the “standard” spectra of carbon materials. Some samples were nanodispersed carbon structures in which sp2 (G) and sp3 (D) bonds were present. Other samples (labeled 2 and 4) contained both nanocrystalline © The Author(s), under exclusive license to Springer Nature Switzerland AG 2023 S. Simakov et al., Nano and Micro Diamond Formation in Nature, SpringerBriefs in Earth Sciences, https://doi.org/10.1007/978-3-031-43278-1_3

19

20

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions

Fig. 3.1 Compositions of organic-water solutions on the C-H–O diagram relative to the field of hydrothermal syntheses in accordance with Fig. 2.3 of Chap. 2. Alcohols: 1—cetyl, 2—Amyl, 3— ethyl, 4—methyl, 5—allyl, 6—glycerol, 7—xylitol (a); Hydrocarbons: 1—pentane, 2—hexane, 3—hectane, 4—toluene, C2 H2 , —average composition of crude oil (b) Fig. 3.2 Raman spectra of experimental samples (1–4); foam (5), diamond (6), disordered graphite (7)

graphite and nanodiamonds (as indicated by the broadening of the 1332 cm−1 band corresponding to the fundamental vibration of the diamond crystal lattice). Another confirmation of the presence of nanodiamonds in those samples is the presence of photoluminescence. This is clearly seen in Fig. 3.2, where samples 2 and 4 are shifted upwards due to luminescence. In fact, this luminescence can indicate the presence of nanodiamonds (sp3 bonds) (Karavanski et al. 2001). The remaining powder after washing was loaded in a colloidal solution in amyl acetate and then applied onto a water surface. The thin film obtained was transferred onto a grid for electron microscopic investigation. Samples were studied with a Tecnai-12 transmission electron microscope (TEM) at an accelerating voltage of

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions

21

120 kV equipped with a “Gotana” camera system. The carrier film typically contained 70–80 nm particles of different forms. Their microdiffraction patterns correspond to particles of cubic (elementary cell a = 3.56 Å) and hexagonal (a = 2.52 Å, b = 3.54 Å) carbine. Photomicrographs of particles and their microdiffraction patterns are shown in Fig. 3.3a, b. The first microdiffraction fits the reciprocal lattice zone [2–15]. Inclination of the particle by 60° to the beam yields the reciprocal lattice zone [364] (Fig. 3.3b). This fact made it possible to unambiguously identify the carbine modification with space group P43m (elementary cell 3.56 Å). Within the limits of measurement error, interplanar spacing, as shown in Fig. 3.3, fit the respective values for carbine (Table 3.1). Microdiffraction patterns of other particles correspond to diamond zones [111] and [100], respectively (Fig. 3.3c, d). Such microdiffraction patterns correspond only to particles with hexagonal and cubic lattices. Interplanar spacings based on these microdiffraction patterns fit the corresponding values for cubic diamond with a = 3.55 Å (sp. gr. Fd3m). None of the other known cubic carbon phases have such interplanar spacing fitting in the diamond parameters within the limits of measurement error (Table 3.1). It should be mentioned that the diamond particles obtained in our experiment are unstable during the radiation and thermal impact of the electron beam. All microdiffraction images were obtained only with a Gotana camera under conditions of the minimal impact of the electron beam. Therefore, we ruled out long-term manipulations with the particles during electron microscopic investigations. The experimental products also include rare particles whose interplanar spaces after inclination by 60° fit those of lonsdaleite (Table 3.1; Fig. 3.3e). In this respect, it is opportune to recall that lonsdaleite has already been observed in the central part of some nanodiamonds (Huss 2005). In addition, platy hexagonal crystals corresponding to chaoite were identified (Simakov et al. 2004). Graphite particles, both well-crystallized and poorly crystallized, were also recorded in the run products. In particular, Fig. 3.3f shows one of these graphite particles and its microdiffraction pattern with {001} reflections, indicating a 2OH modification. It should be highlighted that some of the aforementioned diamond particles grew up to 1 micron in size (Fig. 3.3d), although the hydrothermal diamond nuclei were synthesized without the diamond substrate. Such a relatively large size could be related to the relatively long duration of the experiments in the presence of supercritical water. Moreover, the hydrothermal nanodiamond formation was attributed to the stabilizing role of nitrogen, which was present in the initial experimental reactants. In its absence, the formation of nanodiamond particles did not occur. Cubic diamond particles obtained by the aforementioned experiments in hydrothermal conditions resemble those previously obtained during the precipitation of carbon vapor from electric arc discharge (Matyushenko et al. 1976). The diagnosed set of carbon phases also corresponds to the set obtained by Fedoseev et al. (1976) in metastable diamond synthesis from carbon-containing gases on a substrate. In general, the performed experiments confirm the theoretical calculations presented above in Chap. 2. These experiments are also consistent with the experiments by Ivanova et al. (2016). They

22

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions

Fig. 3.3 Microdiffraction patterns and appearance of particles. Carbine: a axis of zone [2–15], b axis of zone [364]; diamond: c axis of zone [111]; d axis of zone [001]; e ring-textured pattern of lonsdaleite; f graphite, axis of zone [001]

synthesized carbon phases from similar starting materials at 500–800 °C and 0.5– 1 kbar. In the experimental products, the Raman band at 1330 cm–1 indicates the possible presence of a sp3-type carbon phase. Fullerene and fullerene-like phase formation was experimentally obtained from organics at 700–750 °C and 5 kbar (Simakov et al. 2001). Polypropylene (CH3 – [–CH(CH3 )–CH2 –]n –CH = CH2 ) with an admixture of potassium ferricyanide K3 [Fe(CN)6 ], dissociated into C–N complexes upon heating, was used as a source of free carbon. The experiments were conducted using installations of high hydrothermal pressure at the Institute of Experimental Mineralogy of the RAS. The design of an external heating electric furnace allowed us to create a 50-mm-long zone in the reactor, with a thermal gradient lower than 2–3 °C into which an ampoule with the initial material was placed. The temperature was measured with an accuracy of ±5 °C by a chromel–alumel thermocouple. The installation was set up into the experimental regime in 1.5–2 h. The samples were quenched by cold air. During 10– 12 min of quenching, the initial temperature decreased to 150–200 °C. The pressure was measured with an accuracy of ±0.1 kbar by a tube manometer. All experiments were conducted in welded gold ampoules with 0.2-mm-thick walls. Taking into account that the high-pressure vessels generate an oxygen fugacity close to that created by an NNO buffer mixture, we applied the double-ampoule method. The initial mixture (30 mg) was placed into the inner ampoule (diameter 3 mm, length

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions Table 3.1 Calculated and experimental lattices of cubic carbine, lonsdaleite, diamond, octahedral shape carbon (C8), graphite, and α-carbine. Calculations are provided in Figs. 3.3 and 3.5d. Experimental data were taken from ICDD cards № 6–0675, 19–268, 22–1069, 23–64, and from Matushenko et al. (1979)

hkl

dexp , Å

dcalc. , Å

1.48

1.47

23

Carbine 2–11 21–3

0.94

0.96

40–3

0.72

0.72

120

1.60

1.61

Lonsdaleite 100

2,18

2,18

002

2,05

2,06

102

1,48

1,49

110

1,24

1,24

Diamond < 220 >

1,26

1,26

220

1,26

1,26

040

0,88

0,88

3.025

3.03

C8 011 002

2.139

2.14

112

1.746

1.80

013

1.353

1.35

222

1.260

1.235

3.32

3.34

Graphite 002 100

2.1

2.12

110

1.2

1.23

201

1.03

1.05

α-Carbine 111

4.26

4.26

213

2.46

2.46

304

2.1

2.1

227

1.47

1.49

336

1.26

1.26

40 mm). The ampoule was welded up and placed into a larger ampoule (diameter of 5 mm, length 50 mm). Into the outer ampoule, 1000 mg of carbonyl iron and 100 mg of water were added to provide the oxygen fugacity corresponding to the IW-WM buffer mixture. To increase the duration of buffer mixture functioning, a nickel container (length 30 mm, diameter 5 mm) filled with a QFM buffer mixture was placed near the outer ampoule. The experiment duration was 5–7 days.

24

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions

Fig. 3.4 X-ray diffraction pattern of the synthesis. The major peaks are related to graphite (C) and α-carbine (α-C)

After treatment in concentrated HCl to remove possible metallic phases and iron oxides, experimental products were studied by the X-ray diffraction method on a Siemens D-500 diffractometer. The reflection data were obtained using CuKα radiation (Ni filter). The measurements were carried out within the range 2θ = 20°–60° with a step of 0.02° and an exposure of 12'' . Using this method, graphite and another hexagonal phase of C (α-carbine) were identified in the synthesis products (Fig. 3.4). After this, the synthesis products were treated with concentrated HClO4 to remove the graphite component. Then, the samples were dried, doped with 2–3 ml of chemically pure toluene, and subjected to ultrasonic treatment for 10 min at a frequency of 22 kHz on an MSE-100 W instrument (UK). A drop of the obtained suspension was placed on an electron microscope net with a carbon plate and, after evaporation of the volatile components, was observed on a JEM-100S transmission electron microscope (JEOL, Japan) at magnifications of 20,000–100,000 with recording on the FT-41MD film. As a reference preparation, we used carbon nanotubes obtained from a fullerene-producing device and kindly presented by ZAO Astrin. As a result of the experiments, fullerite multilayer formations with diameters from 20 to 80 nm were found. The diameter of the inner cavity or channel in these structures varied from 2 to 6 nm (Figs. 3.5 and 3.6). These features are similar to the short, closed multilayer carbon nanotubes formed during fullerene synthesis under vacuum discharge conditions at ~5000 o K. In addition, a small amount of thin (10–20 nm in diameter) elongated (0.5–1 μm) nanotubes were found (Fig. 3.6). A crystalline phase. Approximately 200 nm of an octahedral shape was also detected in our experiments (Fig. 3.5d). The microdiffraction of its patterns is close to the rare cubic phase of carbon C8 with the lattice parameter a = 4.3 Å (Table 3.1). It was previously diagnosed in diamond-like films synthesized at nearly 1600 °C from carbon plasma by Matushenko et al. (1979). Separate fullerene-like “onion” formations (Fig. 3.6b, c) appeared close to similar onion-like graphitic particles observed in wood charcoal produced at 700 °C (Hata et al. 2000) (Fig. 3.7). The association of the same onion-like particles and diamond

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions

25

Fig. 3.5 Microphotos and microdiffraction patterns of nanotube (a), hexagonal graphite (b), αcarbine (c), and carbon phase of octahedral form (d)

structures was detected in wood charcoals carbonatized at 700 °C (Ishimaru et al. 2001). As a final remark, the experimental data presented in Chap. 3 confirm the theoretical assumption provided in Chap. 2 that nanodiamonds could be formed from carbon-bearing fluids at low temperatures and pressures without diamonds or other seeds. The obtained results show that microdiamonds can be formed from fluids at P–T parameters corresponding to graphite stability. The provided experiments also show that fullerenes, onion carbon structures, and carbon nanotubes can be synthesized from reduced hydrocarbon fluids at temperatures lower than 1000 °C.

26

3 Experimental Data on Nanocarbon Formation Under Low P–T Conditions

Fig. 3.6 Microphotos of fullerite multilayer structures and carbon nanotubes at a magnification of 300 000x

Fig. 3.7 Carbon onion structure in wood charcoal produced at 700 °C (Hata et al. 2000)

References

27

References Chhowalla M, Aharonov RA, Kiely CJ et al. (1997) Generation and deposition of fullerene- and nanotube-rich carbon thin films. Philosop Mag Lett 75(5):329–335 Dresselhaus MS, Dresselhaus G, Eklund P (1996) The science of fullerenes and carbon nanotubes. Press, Acad Fedoseev DV, Deryagin BV, Varnin VP et al (1976) About polymorphism in the carbon and nitride bore systems. Dokl Akad Nauk SSSR 228(2):371–374 Harris PJF (1999) Carbon Nanotubes and related structures: new materials for the twenty-first century. Cambr Univ Press Hata T, Imamura Y, Kobayashi E et al (2000) Onion-like graphitic particles observed in wood charcoal. J Wood Scie 46(1):89–92 Huss GR (2005) Meteoritic nanodiamonds: messengers from the stars. Elements 1:97–100 Ishimaru K, Vystavel T, Bronsveld P et al (2001) Diamond and pore structure observed in wood charcoal. J Wood Scie 47(5):414–416 Ivanova LA, Shumilova TG, Medvedev VYa (2016) Experimental modelling of native carbon formation in a C–O–H fluid system. Dokl Earth Sci 466(2):196–198 Karavanski VA, Mel’Nik NN, Zavaritskaya TN (2001) Preparation and study of the optical properties of porous graphite. JETP Lett 74:186–189 Matushenko NN, Stel’nitsky VE, Gusev VA (1979) New dense modification of crystalline carbon C8. JETP Lett 30:218–221 Schrand AM, Hens SAC, Shenderova OA (2009) Nanodiamond particles: properties and perspectives for bioapplications. Crit Rev in Solid State Mater Scie 34:18–74 Simakov SK, Dubinchuk VT, Novikov MP et al (2010) Metastable nanosized diamond formation from fluid phase. SRX Geoscie 2010: ID 50424 Simakov SK, Dubinchuk VT, Novikov MP et al (2008) Formation of diamond and diamond-type phases from the carbon-bearing fluid at PT parameters corresponding to processes in the earth’s crust. Dokl Earth Scie 421:835–837 Simakov SK, Kalmykov AE, Sorokin LM et al (2004) Chaoite formation from carbon-bearing fluid at low PT parameters. Dokl Earth Scie 399:1289–1290 Simakov SK, Grafchikov AA, Sirotkin AK et al (2001) Synthesis of carbon nanotubes and fullerite structures at temperatures and pressures corresponding to natural processes of mineral formation. Dokl Earth Scie 376:87–89 Smith PPK, Buseck PR (1981) Graphitic carbon in the Allende meteorite: a microstructural study. Science 212:322–324

Chapter 4

Diamond Formation in the Oceanic Lithosphere

Currently, there are numerous reports on diamond findings in nontraditional geological occurrences, such as the oceanic crust and in volcanic rocks other than kimberlites. Regarding “oceanic” diamonds, there is a wide range of opinions regarding their formation (Dobrzhinetskaya et al. 2022; Doucet et al. 2021; Farré-de-Pablo et al. 2018; Galimov and Kaminsky 2021; Pujol-Solà et al. 2020; Litasov et al. 2019; Massonne 2019; Yang et al. 2018; 2020). As previously reported, microdiamonds have been found in ophiolite complexes worldwide. Ophiolites, as is known, are rocky masses already belonging to the lithosphere of former oceanic domains that have experienced tectonic uplift (namely, “obduction”) and placement in different continental areas, especially in collisional contexts. Extensive outcrops of ophiolites are known as complexes or massifs. Although in some cases mineralogical and structural transformations have occurred during ophiolite obduction, these transformations are rarely so extreme as to totally cancel their original petrologic characteristics. Therefore, ophiolites provide valuable insights into the composition and structure of the oceanic lithosphere (e.g., Dilek and Furnes 2014). In particular, diamonds are usually confined to ophiolite sections representing the oceanic mantle, mostly in “chromitite pods.” These diamonds were first identified in Tibetan ophiolites (Fang and Bai 1981) and were later found many times in the same rocks, as well as in their analogs worldwide. It was initially thought that the diamonds found in ophiolites were artifacts due to diamond-coated tools used for cutting or polishing rock samples. The debate on ophiolitic diamond formation, however, was rejuvenated after the discovery of other UHP minerals, such as coesite, together with superreduced phases (SuR), in many chromitites and associated peridotites of ophiolites worldwide (e.g., Griffin et al. 2016). Diamonds were also identified in rocks of Tolbachik volcano (Karpov et al. 2014; Galimov et al. 2016) and in lavas and pyroclastic material of other volcanoes in Kamchatka (Kaminskiy et al. 2016). Diamond nanocrysts are also known in serpentinite xenoliths from Sicily (Central Mediterranea area: Simakov et al. 2015), in melt inclusions in garnet peridotite xenoliths from Hawaiian volcanic products on Oahu Island (Wirth and Rocholl 2003) and in alkaline basalts from Langan (Cai et al. 2021; © The Author(s), under exclusive license to Springer Nature Switzerland AG 2023 S. Simakov et al., Nano and Micro Diamond Formation in Nature, SpringerBriefs in Earth Sciences, https://doi.org/10.1007/978-3-031-43278-1_4

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4 Diamond Formation in the Oceanic Lithosphere

Wang et al. 2020). In the first two cases, xenoliths represent fragments of past and extant oceanic lithosphere. The case of nanodiamonds related to serpentinite from Sicily is reported here in more detail because it sheds light on possible links between the serpentinization process and the formation of diamonds.

4.1 Nano- and Micron-Sized Diamond Formation in the Oceanic Lithosphere 4.1.1 Nanodiamond and Organic Compound Formation in Oceanic Serpentinite Systems Experimental and astrophysical data (Simakov 2010; Kouchi et al. 2005; Nakano et al. 2002) suggest that nanodiamonds could potentially be found in different Earth’s shallow crustal rocks whose formation has involved organic matter and aqueous fluids. Accordingly, Manuella (2013) proposed the hypothesis that serpentinitehosted hydrothermal systems, where organic matter and H2 O coexist over a wide range of temperatures, can be considered potential sites for the nucleation and growth of nanodiamonds.

4.1.2 General Notes on the Serpentinization of the Oceanic Lithosphere International marine geology expeditions have brought crucial advances in understanding the composition and tectonic evolution of the present (and hence the ancient) oceanic lithosphere (e.g., Pearce 2002; Dick et al. 2003; Boschi et al. 2006; Snow and Edmond 2007; Silantyev et al. 2011). Important discoveries particularly regard oceanic ridge systems where spreading develops at slow or ultraslow rates in a moderate, even starving, volcanic regime. For example, these oceanic domains include a large part of the Atlantic Ocean, part of the Indian Ocean, and the Arctic domain. Many nonvolcanic morphological reliefs have been observed in the abyssal plains of slow-spreading oceanic domains, far away from the main ridge, sometimes even emerging beyond the surface of the ocean, as in the case of the islets of Saint Peter and Saint Paul, lying in the equatorial Atlantic Ocean between Brazil and the coast of West Africa (Campos et al. 2010; Palmiotto et al. 2013). These structural highs mostly consist of strongly altered (e.g., serpentinites) mantle peridotites and minor gabbroic and volcanic rocks. In some cases, these reliefs are called “oceanic core complexes” (OCCs), displaying domed surfaces with characteristic frictional streaks (“corrugations”) related to low-angle detachment faults (Cannat and Casey 1995; Miranda and Dilek 2010; Haughton et al. 2019). Seawater-related hydrothermal systems are often associated with OCCs, as testified by a number of

4.1 Nano- and Micron-Sized Diamond Formation in the Oceanic Lithosphere

31

vents discharging aqueous fluids with different temperatures and compositions (e.g., Rona et al. 2010). More generally, fracture occurrences in the oceanic lithosphere, especially in slow/ ultraslow-spreading ridge systems, favor the infiltration of seawater at relatively great depths, causing pervasive hydration of the anhydrous minerals forming the ultramafic and mafic country rocks. Some hydration reactions, such as serpentinization, can reach relatively great depths in the oceanic lithosphere (up to 40 km in depth in the case of serpentinization, assuming an upper mantle temperature of 1350 °C) (Li and Lee 2006). Such a deep hydration front, typical of subduction contexts, can even occur in intraplate settings due to penetration of seawater through major fracture systems (e.g., “macrojoints”: Sokolov 2017; Sokolov et al. 2018). Hydration reactions must take into account the geothermal gradient of the affected area and the stability temperature of the particular hydrated mineral phase(s). Hydration of the oceanic crust is a typical exogenous, low-temperature process, whereas igneous activity is an archetypal, high-temperature, endogenous process. The interaction between the aforementioned opposite processes in an intraplate setting can develop in different ways. Heating of hydrous secondary minerals above their stability temperature due, for example, to the conductive cooling of a nearby igneous intrusion produces dehydration of the former. H2 O will then be released either in the gaseous or supercritical state, and in most cases, it leaves the system in the form of hydrothermal fluid. If such separation cannot occur, due to the lack of an adequate fracture system in the country rocks, the overpressure exerted by the confined vapor phase into the surroundings may act as a geological “bomb,” hence breaching the overburden and possibly producing an explosive eruption (Correale et al. 2019) and even a tuff-breccia pipe or diatreme structure (Manuella et al. 2016). Pressure (depth) plays an important role in dehydration reactions. Once a given pressure threshold is exceeded, “dehydration melting” occurs (e.g., Le Breton and Thompson 1988). More precisely, the breakdown of the particular hydrous mineral phase in the case of dehydration melting does not produce free vapor but one or more anhydrous minerals and a water-undersaturated silicate melt. In other words, a new magma forms (Rapp and Watson 1995). This magma is generally rich in volatiles, and hence, its chemical composition resembles those of so-called “evolute” magmas, such as trachytes (Viccaro et al. 2009). Serpentinization is the archetypal hydration process in the oceanic lithosphere. It denotes different chemical reactions producing serpentine group minerals at the expense of anhydrous or less-hydrous Mg-rich minerals, e.g., olivine, orthopyroxene, clinopyroxene, and amphibole. Serpentine minerals are trioctahedral 1:1 phyllosilicates with a tetrahedral (4-coordinated Si) layer alternating with an octahedral “brucite-type” (6-coordinated Mg) layer. The general chemical formula for serpentine is Mg3 Si2 O5 (OH)4 . The serpentine polymorphs are called lizardite, chrysotile, and antigorite. Lizardite is stable at 50–300 °C, antigorite is stable at 250–680 °C, and chrysotile metastably forms at 0–400 °C (Evans 2004; Ulmer and Trommsdorff 1999). Serpentine minerals often display grain sizes on the micrometer and even nanometer scale. Brucite Mg(OH)2 , talc Mg3 Si4 O10 (OH)2 , magnetite (Fe3 O4 ), and dihydrogen (H2 ) are also possible byproducts of serpentinization reactions. Brucite

32

4 Diamond Formation in the Oceanic Lithosphere

and talc are mutually exclusive, as the presence of one or the other depends on the silica activity in the system. The hydration of an iron-bearing olivine will make Mg-rich serpentine (and brucite), and the iron will form magnetite. The trivalent iron necessary for magnetite formation derives from the deprotonation of the water present in the system, and hence, dihydrogen is produced. If both olivine and orthopyroxene are reactants in equal molar proportions, no brucite or talc will form. For example, Mg-rich olivine (e.g., Fo90 –Fa10 ) and orthopyroxene (e.g., En90 –Fs10 ), as typical in mantle peridotites, react with water, giving serpentine and traces of magnetite and dihydrogen as products, as shown by the following reaction (Evans 2004): ) ( ) ( 1.2 Mg1.8 Fe0.2 SiO4 + 0.76 Mg0.9 Fe0.1 SiO3 + 2.088H2 O 3+ → Mg2.85 Fe2+ 0.11 Fe0.08 Si1.96 O5 (OH)4 + 0.042 Fe3 O4 + 0.176 H2

(4.1)

Dihydrogen produced by the aforementioned reaction can give rise to the catalytic reduction of aqueous carbon dioxide and carbonate present in the system, producing various hydrocarbon species, particularly methane, recalling Fischer–Tropsch (FT) hydrocarbon synthesis. nCO2 + (3n + 1)H2 → Cn H(2n+2) + 2nH2 O

(4.2)

there 1 < = n < + ∞. Magnetite and the aforementioned native metals and alloys, although produced in trace amounts during serpentinization events, act as effective catalysts for such an FT-type reaction (e.g., Charlou et al. 1998; Konn et al. 2009; McCollom 2013; Etiope et al. 2011). Methane formed as a product of the interaction of H2 and CO2 creates acoustic anomalies (e.g., bleaching) in the seismic record of the oceanic crust (e.g., Sokolov et al. 2018). Moreover, magnetite produced by serpentinization reactions has implications for magnetic anomalies in the oceanic crust (Bach et al. 2006). Where serpentine minerals prevail over the primary anhydrous ones, the rock is called serpentinite. Serpentinite textures are defined as pseudomorphic and nonpseudomorphic (Wicks and Whittaker 1977). The shapes of olivine and pyroxene crystals are preserved in the pseudomorphic texture and may contain relics of the primary minerals. Serpentine after olivine typically shows a mesh texture consisting of “cores,” representing the center of serpentinized olivines, and “cords” representing the serpentinized grain boundaries or fractures. Nonpseudomorphic textures either consist of a cross-cutting fabric due to elongated serpentine blades or a massive fabric due to interlocking equate serpentine grains. In both cases, the shape of the original minerals is not preserved (Wicks and Whittaker 1977). Serpentinite “kernels” consist of underformed serpentinized peridotite cores surrounded by serpentinite rims (O’Hanly 1966). The lack of interpenetrating textures in serpentinites indicates that serpentinization in the oceanic crust mostly develops under static conditions (Bach et al. 2006). Porosity in serpentinites is an important, although not fully clarified, issue. Considering a hypothetical ultramafic precursor with 58 volume % of olivine and 42 volume % of orthopyroxene, i.e., a 1:1 molar ratio of both mineral

4.1 Nano- and Micron-Sized Diamond Formation in the Oceanic Lithosphere

33

phases, complete serpentinization results in a volume increase of 45% (Bach et al. 2006). Although such a volume increase should clog any void in the newly formed serpentinite, hence hampering further water inflow, the same volume increase causes increasing stress within the rock, and hence, these stresses can lead to fracturing of the rock (e.g.,Rudge et al. 2010; Kelemen and Hirth 2012; Plümper et al. 2012; Tutolo et al. 2016), favoring the circulation of fluids. In this respect, the tomographic dataset on a natural serpentinite obtained by Pujatti et al. (2019) by means of “two focused ion beam scanning electron microscopy” (FIB-SEM) indicated that a channeled wormhole-like porous network at the micrometer scale occurs at the serpentinization reaction front. In particular, the porosity shows preferential orientation parallel to the grain boundaries of the olivine that undergoes serpentinization. On the other hand, pores at the nanometer scale are isotropically distributed throughout the fully replaced serpentinite matrix, also far away from the reaction front. In particular, the nanopores are arranged along directions subparallel to the coarser porosity. On these grounds, Pujatti et al. (2019) suggested that nanopores are the remnants of the previous microporosity. As already mentioned, serpentinization reactions depend upon different chemical– physical parameters, such as T, P, f O2 , f C O2 , the olivine/orthopyroxene ratio in the host rock, the water (W)/rock (R) ratio, and silica activity (Frost and Beard 2007) in the whole system. W/R ratio is inversely correlated with the pressure and hence with the depth below the seafloor because, as already mentioned, the pressure exerted by the overburden causes a decrease in porosity and permeability in the rocks, hence hampering the water diffusivity. Therefore, fully serpentinized peridotites mainly constitute the upper section of the oceanic lithosphere characterized by W/R > 1, whereas partially serpentinized peridotites are in the lower section (W/R < 1). The dihydrogen produced by serpentinization reactions, such as (4.1), generates extremely reducing conditions. Partially serpentinized peridotites are so reducing that they host Ni–Fe and Ni-Co alloys, native iron, and different SuR (Evans 2004). Native carbon phases, including nanodiamonds, are also putative byproducts of the serpentinization of the oceanic crust, as remarked by Manuella (2013).

4.1.3 Nanodiamonds in Serpentinites—A Case Study from Sicily and Other Occurrences The discovery of nanodiamonds in a widely serpentinized ultramafic xenolith (Simakov et al. 2015) could be considered the first evidence of their formation in a serpentine system, albeit fossil. The nanodiamond-bearing sample is a strongly serpentinized ultramafic rock extracted from a volcanic tuff breccia, late Miocene in age, in a zone called the “Hyblean Plateau” (Sicily, Central Mediterranean area: Fig. 4.1a). There is abundant geological and petrologic evidence that the sample is a fragment of rock torn from a buried serpentinite unit, incorporated as a xenolith into a diatreme-type volcanic system, and finally erupted to the surface in the late

34

4 Diamond Formation in the Oceanic Lithosphere

Miocene (Fig. 4.1b). The source serpentinite unit most likely belonged to the Ionian Tethys oceanic crust, Early Mesozoic in age (Scribano et al. 2006). Transmitted light observations on standard thin sections and thin specks of the untreated organic matter (e.g., Fig. 4.2) revealed the abundance of micrometer-sized black (opaque) carbonaceous flakes interspersed in the semitransparent organic matter (Scirè et al. 2011). The micro-Raman spectra clearly show visible sp3 bonds at 1300 cm−1 and sp2 bonds at 1600 cm−1 for the blebs (Simakov et al. 2015, Fig. 4.2a). In accordance with the method of Koniakhin et al. (2018), it corresponds to nanodiamond particles with a size of 1.3 nm. HRTEM observations of the aforementioned black carbonaceous flakes of serpentinized ultramafic samples showed that they contain aggregates of ultrananocrystalline diamonds ranging in size from 1 to 10 nm (mostly from 1 to 6 nm) (Simakov et al. 2015). The carbon isotopic composition of the treated asphaltene corresponds to -29.0 ‰ δ13 C (Simakov et al. 2015). Meanwhile, the carbon isotopic compositions of the extracted hydrocarbon fractions correspond to -29.8‰ and -32.8‰ δ13 C (Simakov et al. 2015). These values are close to the δ13 C values of gaseous hydrocarbon seepage from serpentinizing ultramafic rocks of Greece (Etiope and Schoell 2014), which confirms the shallow nanodiamond origin in serpentinite. Raman and photoluminescence spectra for another two asphaltene blebs containing black carbonaceous flakes display sp3 bonds from 1287 to 1312 cm−1 and sp2 bonds from 1580 to 1600 cm−1 (Simakov et al. 2021, Fig. 4.3b). In accordance with the method of Koniakhin et al. (2018), it corresponds to nanodiamond particles of size 1–1.6 nm. Raman (RS) and photoluminescence (PL) spectra were also recorded on an “inVia Raman microscope” (Renishaw) at room temperature at the Lebedev Physics Institute in Moscow for these blebs. Raman microscope attachment has excitation emissions of 785 nm and 1064 nm, which strongly reduced the intensity of the luminescent background in the spectra of the samples. A laser excitation source at 785 nm was applied, and the laser power was controlled by means of a series of density filters to avoid heating effects from 100% to 5 × 10–8 %. Excitation light was focused to a spot diameter of 2–5 μm. The absence of the laser heating effect on the sample was controlled through a microscope. The multichannel system of registration controlled a satisfactory signal/noise ratio. The resolution of the device in the entire spectral region was ~ 1 cm−1 . It showed peaks of sp2 hybridization at 1586–1607 cm–1 and peaks of sp3 hybridization at 1286–1324 cm–1 (average 1304 cm–1 ) (Fig. 4.3). Bands at 1586–1607 cm–1 may be due to the graphitic planes interconnected by sp3 bonds (Karavanski et al. 2001). The 1286–1324 cm–1 band can be attributed to nanodiamonds. In accordance with the method of Koniakhin et al. (2018), it corresponds to nanodiamond particles of size 1–2.5 nm. Diamonds of the same size have been obtained by synthesis (Stehlik et al. 2015) and detected in meteorites (Daulton 2006). In accordance with the “archetypal macrosopic molecule” classification of Dahl et al. (2003), nanoparticles of 1–2 nm sizes correspond to “higher diamondoids” formed in the temperature range of 150–350 °C under low-pressure conditions. Molecules of higher diamondoids (e.g., C22 and higher polymantanes) contain 4–11 diamond crystal cages. The latter consists of 88 atoms, which corresponds to a diamond sphere with a diameter of nearly 1 nm. The octamantane class with formula C34 H38 shows 18 isomeric structures with both chiral and achiral forms. All of

4.1 Nano- and Micron-Sized Diamond Formation in the Oceanic Lithosphere

35

Fig. 4.1 Location of the study area and geological sketch of the corresponding lithospheric column. a Satellite view of part of the Mediterranean Sea, taken from “Google Maps,” showing the location of the xenolith-bearing diatreme structures in the southern part of Sicily. b General indications of the composition of the lithospheric column in the study area as inferred from field geology data, core drillings, xenoliths in volcanic rocks (tuff-breccia pipes, late Miocene in age: Manuella et al. 2015), and geophysical data (Giampiccolo et al. 2017)

these higher diamondoids correspond to nanometer-sized, H-terminated diamonds of diverse shapes and sizes (Dahl et al. 2003). In accordance with the model of Badziag et al. (1990), they correspond to the transition of hydrocarbons to hydrogenterminated diamonds. Petrologic and geochemical characteristics of the sample were reported by Scirè et al. (2011) and Simakov et al. (2015). The occurrence of altered relicts of spinel grains (Fig. 4.2) suggests that in its earliest geologic history, the sample equilibrated in the upper mantle in the spinel peridotite stability field, where graphite is the stable carbon allotrope. The different generations of serpentine veins in the sample, along with their relationships with Fe-oxide and Fe–Ni sulfide micrograins, suggest that serpentinization developed in several stages at different temperatures and oxygen fugacities. The secondary mineral assemblage in the sample suggested the presence of seawater-derived hydrothermal fluid with dissolved CO2 . Simakov

36

4 Diamond Formation in the Oceanic Lithosphere

Fig. 4.2 Microphotographs (plane-polarized light) of a thin section of the sample (a, b, c) and (d) a tiny fragment of the condensed OM. Srp: serpentine group minerals (chrysotile and lizardite); Cpx: clinopyroxene (Cr-bearing diopside); Spl: spinel (picotite). Arrows in (b) indicate nanodiamondbearing opaque specks dispersed in the OM. A full explanation is given in the text Fig. 4.3 RS spectra of the sample obtained using a spectrometer at the excitation of 785 nm: 1330 cm–1 indicating the presence of nanodiamonds; 1600 cm–1 related to nanographite

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Fig. 4.4 Variation in the C–O–H system composition with oxygen fugacity at 300 °C and 2 kbar in accordance with the gaseous system of Zhang and Duan (2010). Legend: MH, CCO, QFM; WM is reproduced from Fig. 5 of Simakov et al. (2021)

et al. (2015), based on the phase relationships in the MgO-SiO2 -H2 O system (Evans 2004; Klein and Bach 2009), suggested that the highest serpentinization temperature was ~ 300 °C. During the formation of dihydrogen by reaction (4.1), f O2 changed from QFM to WM buffer. This corresponds to the variation in the C-O–H fluid composition from water-rich to methane-rich (Fig. 4.4). This fluid transition explains the significant emission of hydrogen and methane over the ophiolitic zones (Schrenk et al. 2013). In accordance with our physicochemical model of nanodiamond formation (see Chap. 2), in the H–C system, the formation of nanodiamonds is possible by the following reaction: CH4 → C + 2H2

(4.3)

The 1–6-nm diamonds could be formed by this reaction. As a result, the full replacement of olivine, partial replacement of orthopyroxene, and formation of a hydrocarbon medium containing immersed carbon particles follow (Fig. 4.2c). Then, the f O2 of the system evolved from reducing to significantly oxidizing conditions (e.g., above the MH buffer) (Fig. 4.4), with the water/rock ratio increasing concomitantly over time. The final stages in the history of the sample’s alteration were characterized by its incipient steatization, formation of clays after serpentine, and severe carbonatation of hydrocarbons and silicate minerals; these processes occurred at temperatures of nearly 150 °C and at increasing f O2 and f C O2 hence, decreasing f H2 conditions (Simakov et al. 2015). Diamond nuclei formation is possible at this stage in the range of the upper limit of carbon stability corresponding to the CCO buffer (see Chap. 1, Fig. 4.1b, c) by the reaction of Rudenko et al. (1993): CH4 + CO2 → 2C + 2H2 O Rare 10-nm nanodiamonds could be crystallized during these processes.

(4.4)

38

4 Diamond Formation in the Oceanic Lithosphere

The size of the nanodiamonds is reported to decrease with increasing temperature (see Chap. 2, Fig. 2.1). Since there is no agreement between the results of the cited authors, it is not possible to extrapolate univocal temperature values that are always compatible with the different stages of the serpentinization process (e.g., Evans 2004; Klein and Bach 2009). The organic matter of the samples presents characteristics typical of “petroleum asphaltenes” as n-heptane-insoluble macromolecules with unknown structures, composed mainly of polyaromatic carbon ring units with oxygen, nitrogen, and sulfur heteroatoms and aliphatic side chains of various lengths. Heteroatoms may be present both in heteroaromatic structures (such as pyrazine, pyrrole, carbozole, indole, and thiophene) and in substituting groups (i.e., hydroxyl, ether, ester, aldehyde, ketone, carboxylic, thiol, amine, and amide). It is worth noting that highmolecular-weight, sulfonated, nitrogenated, fluorogenated, and halogenated organic compounds of likely abiotic origin, compatible with the cracking of asphaltenes, were detected in fluid inclusions in different sulfide minerals from hydrothermal vents, often rooted in serpentinite systems, in the Atlantic Ocean seafloor (Tomilenko et al. 2022). The association of nanodiamonds with abiotic asphaltenes and serpentine in the Sicilian xenoliths provides indirect evidence of nanodiamond formation in serpentinite systems, as the rock sample represents a fragment of a buried serpentinite body belonging to the extinct Permian-Tethys ocean. Therefore, the presence of nanodiamonds coexisting with condensed organic matter and serpentine in a sample of ultramafic rock in a current oceanic context can be regarded as a very important discovery (Andreani et al. 2023). In particular, a nanodiamond-bearing troctolite sample has been drilled in the OCC known as the “Atlantis Massif,” which is mainly composed of strongly serpentinized mantle peridotites and mafic intrusive rocks exhumed along the Mid-Atlantic Ridge (Andreani et al. 2023). Nanodiamond-bearing organic matter, coexisting with serpentine minerals, brucite, and magmetite, occurs in microcavities in troctolite olivine grains. The authors suggest that such microcavities were previously fluid inclusions filled with an aqueous fluid, which promoted serpentinization reactions with the host olivine walls. These aforementioned results also suggest that nanodiamond formation is possible during the oxidative degradation process of crude oils (Simakov 2010), together with the formation of macromolecular hydrocarbons as well as bitumens, anthraxolites, and asphaltenes. The hydrocarbons and oil formed via FT-t reactions can be dissolved in water at 300–400 °C and 220 bar (Chekaluk and Filyas 1977; Price 1976). Their water solutions cross the hydrothermal nanodiamond domain on the C-H–O diagram and can be the sources for nanodiamond formation (see Chap. 3, Fig. 3.1b).

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4.1.4 Nanodiamonds in Serpentinites as Proxies for the Emergence of Life on Early Earth Finally, it is worth mentioning that the discovery of nanodiamonds in a sample of serpentinite containing condensed organic matter has allowed Simakov et al. (2021) to consider the still open problem of the appearance of bioorganic molecules under primitive Earth conditions. It is an established notion that surfaces of several mineral phases are also thought to be catalytic platforms stimulating the assemblage of complex bioorganic molecules relevant to the emergence of life on Earth (Bebié and Schoonen 2000; Holm et al. 1993; Hazen and Sverjinsky 2010). They can create water crystalline interfacial water layers that could be the medium for life starting (Szent-Gyorgyi 1971). The diamond surface could create crystalline interfacial water layers even at room temperature (Sommer et al. 2008), which could be a good medium for rapidly transporting positive charges in the form of hydrated protons. It has been demonstrated that a high concentration of hydrogen in the gas mixture is required to promote the formation of nanodiamond bonds, as opposed to the equilibrium form of carbon and graphite (Spitsyn et al. 1981; Matsumoto et al. 1982; Fedoseev et al. 1984). In a later stage of serpentinite system formation, the hydrated nanodiamond surface could create crystalline interfacial water layers. Hydrogen could be outgassed from nanodiamonds at temperatures as low as 100 °C (Landstrass and Ravi 1989) and create a high [H+ ] layer on cluster surfaces. Due to this effect, the transformation of the LUCA structure (Last Universal Common Ancestor of cells: Lane et al. 2010) to living cells might be possible on the surface of the carbonaceous cluster (Fig. 4.5).

Fig. 4.5 Rough drawing of the crystalline interfacial water layers on the diamond surface (reproduced from Fig. 6 of Simakov et al. 2021)

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4 Diamond Formation in the Oceanic Lithosphere

4.1.5 Micron-Sized Diamonds in Ophiolites—Insights into Their Formation Currently, several scientific groups from different countries have independently confirmed the existence of diamonds in ophiolites (Yang et al. 2014, 2015a, 2015b, 2017 and references therein reported). For instance, diamonds have been discovered in mantle peridotites and chromitites from six ophiolitic massifs along the 1300 kmlong Yarlung-Zangbo suture zone (Bai et al. 1993; Yang et al. 2014; Xu et al. 2015) and in the Dongqiao and Dingqing mantle peridotites in the Bangong-Nujiang suture zone (Robinson et al. 2004; Xiong et al. 2018). Recently, in situ diamond and coesite have also been reported in the Nidar ophiolite of the western Yarlung-Zangbo suture zone (Das et al. 2015, 2017). The abovementioned diamond-bearing ophiolites represent remnants of the Mesozoic Eastern Tethys oceanic lithosphere. Recent papers show that diamonds also occur in chromitites in the Pozanti-Karsanti ophiolite in Turkey and in the Mirdita ophiolite in Albania belonging to the western Tethys domain (Lian et al. 2017; Xiong et al. 2017; Wu et al. 2018). Similar diamonds and associated minerals have also been reported from Palaeozoic ophiolitic chromitites in the Central Asian Orogenic Belt of China and in the Ray-Iz ophiolite in the Polar Urals, Russia (Yang et al. 2015a, b; Tian et al. 2015; Huang et al. 2015). Importantly, in situ diamonds have been recovered in chromitites of both the Luobusa ophiolite in Tibet and the Ray-Iz ophiolite in Russia (Yang et al. 2007, 2014, 2015a). Many unusual minerals have been previously reported from podiform chromite deposits, both in situ and in mineral separates. Diamond-like micrograins were reported early in chromite ores in southern Quebec, Canada (Dresser 1913), but they were later recognized as synthetic periclases formed by heating of the samples (LeCheminant et al. 1996). In more recent times, there have been several reports of “real” diamonds in ophiolites, including those in the Eastern Sayan (Shestopalov 1938), Kamenusha massif in the Urals, and the Amasia-Sevan-Akera belt of Armenia (Kaminskiy and Vaganov 1977; Lukyanova 1980). Diamonds in the Pamali breccias in Indonesia and alluvial diamonds in British Columbia (Canada) and Copeton– Bingara (New South Wales) are thought to have been derived from ophiolites (Bergman et al. 1987; Pearson et al. 1989; Strnad 1991). The ophiolite-hosted micrometric diamonds contain mineral and fluid microinclusions consisting of H2 O, carbonates, chromite, magnesiochromite, magnetite, albite, hematite, moissanite (SiC), solid CO2 , coesite, Mn silicates such as tephroite (Mnrich olivine), spessartine (Mn garnet), MnO, native Mn, Ca-perovskite and metallic alloys such as Fe–Ni and Ni-Mn-Co alloys (Lian et al. 2018; Moe et al. 2016, 2017, 2018; Wu et al. 2019; Yang et al. 2015a, b, 2017) (Table 4.1). The carbon isotopic composition of the aforementioned microdiamonds varies from −18.9‰ to −28.30‰ of δ13 C with a leading mode at −25‰. Their nitrogen content varies from 108 to 589 ppm with a prominent mode in the range of 250 ppm (Xu et al. 2017). The mineral assemblages accompanying microdiamonds in their ultramafic host rocks mainly include native elements (e.g., native iron, native chromium, and graphite), alloys, and SiC. Wüstite, orthopyroxene, wadsleyite, and coesite pseudomorphous of

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stishovite and qinsongite (cubic boron nitride) also occur in the mineral assemblages, although less frequently (Table 4.1). Another interesting discovery concerning ophiolite complexes containing diamonds is that, in addition to the UHPT and SuR mineral phases, there are some minerals, such as zircon, corundum, rutile, and feldspars (Table 4.1), which are generally ascribed to the continental crust. Compared to diamonds in kimberlites and those in “UHP” metamorphic belts, microdiamonds from ophiolites represent a new occurrence of diamonds that requires significantly different physical and chemical conditions of formation. On the one hand, diamond inclusions such as coesite, Ca-perovskite, and moissanite could show their deep mantle formation. On the other hand, other inclusions, such as chromite, hematite, magnetite, feldspar, zircon, and magnesiochromite, indicate shallow-depth genesis. However, to date, there have been no reports documenting the presence of nanodiamonds in serpentinite samples obtained through dredging or coring during various marine geology expeditions worldwide. One possible explanation is that the search for nanodiamonds in oceanic serpentinite samples has not been conducted, as it requires specific investigations. Additionally, scientific coring in oceanic serpentinite systems typically reaches depths of only a few hundred meters, where the relatively high water-to-rock (W/R) ratio promotes the oxidation of reduced carbon species into carbonates. Moreover, Litasov et al. (2019), by analyzing literature data, concluded that the metal alloys present in ophiolite microdiamonds resemble those in synthetic diamonds, suggesting that these diamonds, likewise those from Tolbachik area, represent contamination from drilling tools or other diamond-coated instruments having, therefore, a synthetic origin. In this respect, Yang et al. (2020) remarked that Litasov et al. (2019) did not consider the petrologic context related to microdiamond occurrences and the entire set of mineral inclusions discovered in ophiolite-hosted diamonds and, more in general, that similarities between natural and artificial materials do not necessarily mean that the natural materials are synthetic. In this respect, reasonably assuming that ophiolitic microdiamonds can be regarded as natural occurrences (Yang et al. 2020), there are different hypotheses on their origin. For example, Ballhaus et al. (2017) suggested that some of the nominally UHP and superreduced minerals, including (micro)-diamonds, in worldwide ophiolites, can be formed by quick condensation of high-T plasma originating from lightning strikes on chromitite bodies exposed on the Earth’s surface. However, this hypothesis does not seem to have passed the scrutiny carried out by Xiong et al. (2019), who have shown, among other things, that diamonds in ophiolites contain a wide range of UHP and SuR mineral and fluid inclusions that have never been reported from mineral phases in worldwide fulgurite occurrences, and hence, the high temperature of lightning strikes would need to be sustained for hours to produce diamonds in the typical rocks of ophiolite complexes, including chromitiites. Apart from the lightning strike hypothesis, there are two main schools of thought on the origin of microdiamonds in ophiolites and, more generally, in the oceanic crust. (a) The most popular hypothesis suggests HPHT genesis in a deep-sited region of the oceanic mantle (e.g., Yang et al. 2014). As mentioned before, such a great depth of diamond formation has been inferred from the HPHT mineral inclusions (Table 4.1). Moreover, HPHT minerals (e.g., coesite pseudomorphose after stishovite,

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4 Diamond Formation in the Oceanic Lithosphere

Table 4.1 Synoptic list of minerals found in ophiolithic microdiamonds or accompanying microdiamonds in ophiolithic rocks Mineral name

Chemical formula

Crystal system

Specific gravity

Note

Awaruite

Ni3 Fe

Isometric

7.8–8.22

SuR

Albite

Na(AlSi3 O8 )

Triclinic

2.6–2.65

Crustal

Ca-perovskite

CaTiO3

Orthorhombic

3.98–4.26

UHTP

Calcite

CaCO3

Trigonal

2.7

–

Chromite

Fe2+ Cr3+ 2 O4

Isometric

4.5–4.8

–

Coesite

SiO2

Monoclinic

2.92

UHPT

Corundum

Al2 O3

Trigonal

3.98–4.1

–

Dolomite

CaMg(CO3 )2

Trigonal

2.84–2.86

–

Graphite

C

Hexagonal

2.09–2.23

SuR

Hematite

Fe2 O3

Trigonal

5.26

–

Kyanite

Al2SiO5

Triclinic

3.53–3.67

Crustal

Magnesiochromite

MgCr2 O4

Isometric

4.1–4.3

–

Magnesite

MgCO3

Trigonal

2.98–3.02

–

Magnetite

Fe2+ Fe3+ 2 O4

Isometric

5.17

–

Moissanite

SiC

Hexagonal

3.218–3.22

SuR

Native Chromium

Cr

Isometric

7.17

SuR

Native iron

Fe

Isometric

7.3–7.87

SuR

Ni-Mn-Co alloy

NiMnCo

–

–

**

Osbornite

TiN

Isometric

–

UHTP

Qinsongite

BN

Isometric

–

UHTP

Rutile

TiO2

Tetragonal

4.23

–

Spessartine

Mn2+ 3 Al2 (SiO4 )3

Isometric

4.12–4.32

Crustal

Stishovite

SiO2

Tetragonal

4.35

UHTP

Tephroite

Mn2+

Orthorombic

3.87–4.12

UHTP

Wadsleyite

Mg4 O(Si2 O7 )

Orthorombic

3.84

UHTP

Wüstite

FeO

Isometric

–

SuR

Zircon

ZrSiO4

Tetragonal

4.6–4.7

Crustal

2 SiO4

high-pressure forms of chromite, and qingsongites) were also found in the chromitites and peridotite host rocks. Therefore, Yang et al. (2014) reported that ophiolitic chromitite may form at depths of 150–380 km or even deeper in the mantle. The very light C isotope compositions of these ophiolitic diamonds and their mineral inclusions suggest recycling of ancient crustal materials into the mantle (>300 km) or down to the mantle transition zone via subduction. Yang et al. (2018) finally remarked that “The widespread occurrence of ophiolite-hosted diamonds and associated UHP mineral groups suggests that they may be a common feature of in situ oceanic mantle.”

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The aforementioned hypothesis regards only types of ophiolites that are believed to derive from SSZs, that is, from the upper plate at a convergent plate boundary, including intraoceanic subduction (e.g., Shervais 2001). In particular, the hypothesis considers the case of oceanic lithosphere having undergone deep subduction and subsequent exhumation. This geological scenery could be framed in the geodynamic model reported as “accretionary uplift” by Shervais (2001). Nevertheless, Shervais (2001) remarked that the aforementioned mechanism should not occur in the case of obduction of ophiolites on a passive continental margin. It should be noted, however, that the mantle ultramafics hosting both the diamonds and the chromitite bodies, although widely serpentinized, can be securely ascribed to the spinel peridotite facies, hence indicating a shallow mantle origin. In this respect, Xiong et al. (2019) suggested that diamonds were carried to shallow mantle levels by asthenospheric magmas due to a decompression melting process during slab rollback. In addition, the aforementioned authors reported that the SuR mineral assemblage coexisting with microdiamonds in ophiolites is due to “fluid percolation and crystallization of alteration-related minerals in the lithospheric parts of a (hydrated) mantle wedge, resulting in the formation of highly reduced microenvironments.” However, it must be mentioned that in some geological processes, including serpentinization, superreduced microenvironments are created even at low temperature and low-pressure conditions (e.g., Farré-de-Pablo et al. 2018). Accordingly, Golubkova et al. (2016) reported that moissanite can be deposited at low pressure and low temperature by ultrareducing fluids released during the initial stages of serpentinization of olivine in peridotites. Moissanite and native elements were also found in metamorphic quartz–muscovite dynamo-schists from the Baikal ledge Sarma region, formed at 660–700 °C and 3 kbar in the presence of essentially hydrocarbon fluid (Savelyeva et al. 2019). In situ moissanite was also observed in porphyroblastic garnet–staurolite micaschists from Bulgaria (Machev et al. 2019). Regarding UHTP minerals, such as coesite, coexisting with ophiolitic microdiamonds, it should be noted that coesite can form on crushed quartz at 450–900 °C and 5–20 kbar (Green 1972). The occurrence of minerals believed to be of continental crust origin coexisting with UHPT minerals (Table 4.1) in diamond-bearing ophiolites seems a puzzling issue in the frame of the deep-origin viewpoint. In this respect, the case of the zircon is worth mentioning in some detail due to the obvious geochronology implications. Interestingly, in some cases, there is a significant discrepancy between the zircon U–Pb age and the age of the ophiolite host rock. For example, Xiong et al. (2019) reported xenogenic zircon grains with ages ranging from 1718 to 465 Ma hosted in the diamond-bearing Purang Ophiolite (Tibet) Early Cretaceous in age (130 Ma). The authors hypothesize that the old zircons represent continental crustal material delivered into the upper mantle via previous subduction events and subsequently carried to shallow mantle levels by asthenospheric magmas. In contrast, experimental petrology results by Borisova et al. (2020) indicate that typical 100-micron zircon crystals dissolve rapidly (~10 h) and congruently upon reaction with basaltic melt at pressures of 2–7 kbar. Even if basaltic magma is absent and zircon dissolution does not occur, mantle-recycled crustal zircon grains may not retain their original Th-U–Pb isotope ratios, as reheating zircons at T > = 1000 °C dramatically increases

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the Pb2+ diffusion coefficient. For instance, Pilot et al. (1998) suggested that under shallow mantle conditions (~1000 °C), zircon grains of normal size would totally lose their radiogenic Pb in approximately 150 million years. At temperatures ≥ 1200 °C (90–130 km in depth), Pb would be lost in less than 0.05 Myr (Bea and Montero 2013). The zircon Th-U–Pb isotope ratios remain unaffected by high temperature only if zircon remains totally included in a mineral (such as olivine) where lead has low diffusivity (Bea et al. 2018). Regarding the occurrence of older-than-host zircon crystals, without inherited cores, in the present and “fossil” oceanic crust (e.g., Pilot et al. 1998; Bortnikov et al. 2008; Skolotnev et al. 2010), it must be mentioned that intracrystalline redistribution of radiogenic Pb can occur for different causes during the geologic history of the zircon grains, possibly generating meaningless ages (e.g., Kusiak et al. 2013). More precisely, radiogenic Pb-enriched patches will yield spuriously old 207 Pb/206 Pb ages that are not true ages but the result of sampling by the probe beam that integrates areas with both unsupported and supported radiogenic Pb. (b) An alternative viewpoint considers that microdiamonds originated at relatively shallow depths in the oceanic lithosphere as a consequence of serpentinizzation reactions. Farré-de-Pablo et al. (2018) provided evidence of microdiamond formation in sealed fractures of chromites during serpentinization of ultramafic rocks in the oceanic crust. In particular, diamond precipitated metastably at low pressure from reduced C-O-H fluids that infiltrated from the host peridotite at the onset of serpentinization processes. Chromite chemical variations across the diamond-bearing healed fractures indicate this formation at temperatures between 670°C and 515°C. Simakov (2018) also suggested that micrometer-sized diamonds could be formed from hydrocarbon fluids under supercritical water conditions during the serpentinization of mantle peridotites already tectonically uplifted at the ocean floor. Accordingly, intergranular micron-sized diamonds were found between olivine, serpentine, and amorphous carbon in the Ray-Iz ophiolite of the Polar Urals (Yang et al. 2015a, b, Fig. 8c). Following Makeyev (1992), the stability conditions of serpentine minerals from Ray-Iz chromites vary from greenschist to granulite facies in the range of 300–700°C. In addition, diamonds up to 300 nm in size were discovered within olivine-hosted fluid inclusions from low-pressure oceanic gabbro and chromitite samples of the Moa-Baracoa ophiolitic massif, eastern Cuba (Pujol-Solà et al. 2020). Diamonds are found within CH4 -bearing fluid inclusions forming linear arrays (healed fractures) in olivine and surrounded by serpentine. In particular, Pujol-Solà et al. (2020) suggested that the diamond-bearing fluid and the coexisting mineral assemblage (olivine + serpentine + magnetite) can form at pressure and temperature conditions of approximately 1000 bar and 350°C, respectively. As mentioned above, ophiolites represent fragments of “fossil” oceanic lithosphere (Coleman 1971). In this respect, it is opportune to recall that modern marine geology expeditions have brought crucial advances in understanding the composition and tectonic evolution of the extant oceanic lithosphere, having discovered, for example, large areas of the sea floor paved with serpentinites locally hosting active hydrothermal systems (e.g., Blackman et al. 2002; Boschi et al. 2006; Dick et al. 2003; Früh-Green et al. 2004; Ildefonse et al. 2007; Kelley et al. 2001, 2007; Miranda and

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Dilek 2010; Silantyev et al. 2011; Snow and Edmond 2007). Serpentinite-hosted hydrothermal systems are thought to be ideal places for the catalytic hydrogenation of dissolved inorganic carbon species (e.g., Szatmari 1989; Holm and Charlou 2001; Charlou et al. 2002; McCollom and Seewald 2007). In fact, dihydrogen, methane, and other light organic compounds are common components of fluids issuing from hydrothermal seafloor vents (e.g., Konn et al. 2009). Dihydrogen necessary for abiotic organic synthesis is mostly released by the oxidation of ferrous iron in mafic silicate minerals by the protons of water. Typical cases of H2 production are due to the hydration of olivine and orthopyroxene in the peridotites, giving serpentine and magnetite, as illustrated, for example, by reaction (4.1). More generally, this reaction explains why superreducing environments are formed during the serpentinization process. Although the plumbing of serpentinite-hosted abyssal hydrothermal systems is poorly known, it is an established notion that the basic functioning of these systems is due to the convective circulation of seawater-derived aqueous solutions through deep fracture systems (e.g., Rona et al. 2010). The heating of the percolating seawater (e.g., descending limb) at the temporary base of the hydrothermal system (e.g., Fig. 4.6) can be due to the presence of an underlying magmatic sill or by the serpentinization reaction itself (reaction 4.1), which exhibits exothermal behavior with ~35 kJ/mol H2 O (Evans 2004). In the first case (igneous sill), the circulating fluid can attain temperatures as high as 400°C, even reaching a supercritical state. In the second case (exothermal reaction), the temperature does not exceed 90 °C (e.g., Rona et al. 2010). When the hydrothermal solution rises fast as an ascending limb of the convective system through an open fracture of the host serpentinite, the fluid undergoes depressurization far below its vapor saturation pressure, giving rise to gas bubble nucleation (Fig. 4.6). Bubbling likely occurs on nucleation centers, such as solid particles suspended in the fluid. Some of such solid particles may be nanosized carbon particles, including nanodiamonds (Manuella 2013; Simakov et al. 2015; Simakov 2018 and references there reported) and a plethora of serpentinite-related minerals, including native metals, alloys, sulfides, and silicates. After the fluid has completely filled the channel, the inner pressure returns to its initial value, causing gas bubbles to implode (i.e., hydrodynamic cavitation) with sudden, extreme increases in temperature and pressure (Flannigan and Suslick 2005), which gives rise to a shock wave often accompanied by light emission (sonoluminescence), immediately followed by rapid cooling (Fig. 4.6). There are different experimental results on bubble sonoluminescence (e.g., Brenner et al. 2002; Chen et al. 2008). In particular, Didenko et al. (1999) reported that multibubble sonoluminescence (MBSL) spectra irradiating from a water/benzene mixture at 278 K indicate an emission temperature of 4300 ± 200 K. Moreover, experimental results by Flannigan and Suslick (2005) on single-bubble luminescence in an acidic fluid indicate the existence of a hot, optically opaque plasma core in the collapsing bubbles. On these premises, reasonably admitting that both organic and inorganic carbon occur in the bubble gas phase in serpentinite systems, the energy released by the collapsing bubble may promote reaction (4.4) and the following reactions:

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4 Diamond Formation in the Oceanic Lithosphere

Fig. 4.6 Conceptual sketch of possible microdiamond formation in oceanic serpentinite systems due to hydrodynamic cavitation. a Geological setting of a typical serpentinite-hosted hydrothermal system (not to scale cross-section (e.g., Ondréas et al. 2012). The heat source can either be a conductively cooling subjacent igneous body or an exothermal serpentinization reaction. b Detail of a section of the ascending limb of the hydrothermal system, which highlights an empty widening causing transient decompression of the rising fluid and consequent gas bubble nucleation (c) and growth (d). After filling the widening (e), the fluid continues to rise along the conduct, re-establishing the initial hydrostatic pressure and hence causing the implosion of the gas bubbles. Nanodiamonds already occurring in the fluid can be entrapped in the growing bubbles (f, g). The final implosion of the bubble (h) releases a relatively large amount of energy (sonoluminescence: Flannigan and Suslick 2005), which could favor the synneusis of the nanodiamonds and the formation of microdiamonds (i)

2CO → C + CO2

(4.5)

CO + H2 → C + H2 O

(4.6)

Already existing nanodiamonds can act as “crystal seeds” for newly formed reduced carbon, hence increasing in size. An “Ostwald ripening” process will finally form microdiamonds. The energy of the bubble implosion may also induce synneusislike phenomena between nanodiamonds (Fig. 4.6). It is worth noting that, in addition to the nanodiamond, any other mineral particle carried in suspension by the rising

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hydrothermal fluid can be a nucleation site for gas bubbles. Minerals forming nucleation sites for gas bubbles will therefore be affected by the strong pressure wave and transient flush of heat caused by the subsequent bubble implosion. Therefore, minerals that are stable under crustal P–T conditions will be suddenly transformed into HPHT phases (Table 4.1). This is one of the geological cases where pressure is not related to depth. In this respect, if the newly formed microdiamonds include one or more of the aforementioned HPHT minerals, such an association can be correctly considered syngenetic. Hydrodynamic cavitation is, therefore, a plausible mechanism for the origin of microdiamonds and their UHPT solid inclusions in the oceanic lithosphere (Galimov et al. 2016; Manuella et al. 2018 and references therein reported). On the above grounds, cavitation may also occur in multicomponent-multiphase silicate magmas bearing dissolved carbon, particularly during their ascent to the surface via volcanic conducts that exhibit thickness variations or other morphological irregularities (Galimov et al. 2016). As shown in Chap. 2, diamond growth was also possible in supercritical aqueous fluids (Capelli, 1995). CH3 detected in the asphalten serpentinites (Scirè et al. 2011) could be an important growth precursor in diamond synthesis. An additional, very important aspect of microdiamond formation in abyssal serpentinite systems regards the role of metal catalysts, such as native elements and their alloys. These can stimulate diamond growth up to hundreds of microns in size. (It is important to note that Tehuitzingo, South Mexican diamonds with no native elements, are only a few microns in size, while Luobusa and Ray-Iz diamonds with native elements and their alloys reach several hundred microns in size.) Experiments on low-pressure diamond synthesis apply catalyst-solvent phases (cubic BN, Si, SiC, Cu, Mo, Ni, and Pt), which are isostructural with diamonds (e.g., Mallika et al. 1999). In particular, transition metals (mainly Ni) promote the dissolution of graphitic carbon into the catalyst, forming interstitial carbides and carbon solid solutions, and then the conversion into diamonds (Fei et al. 2001; Mallika et al. 1999). These metals possess good carbon solubility (decreasing in the series Mn>Fe > Co > Ni > Cu > Zn) and form stable carbides, whose stability decreases in the same order as carbon solubility (Mallika et al. 1999). As previously mentioned, zircon grains were found, together with other crustal minerals, in some diamond-bearing ophiolites. In this respect, Meng and Zhang (2009) suggested that zircon crystals in serpentinized dunite veins from North Qaidam, northwestern China, were precipitated from hydrothermal fluid related to serpentinization of dunite. Moreover, Dubinska et al. (2004) investigated a number of zircon grains in the rodingite blackwall from the Jordanow–Gogolow serpentinite massif (Sudetic Ophiolite, Poland). From the petrological data and fluid-inclusion evidence, the authors concluded that zircons coprecipitated with corrensite and chlorite from low-temperature (270–300 °C), highly alkaline hydrothermal solutions genetically related to serpentinization. More generally, although zirconium is known as the archetypal of geochemically “immobile” elements, it is quite mobile in some hydrothermal settings where F, Cl, phosphate, sulfate, and carbonate anions act as ligands for Zr complexing (Rubin et al. 1993).

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As already reported, ophiolitic microdiamonds have often been found in podiform cromitites hosted in serpentinites. Chromitites are generally considered igneous cumulates closely associated with oceanic spreading processes (e.g., Li et al. 2002 and references therein reported). In contrast, such ore bodies can be formed during the serpentinization of mantle peridotites uplifted at crustal levels in oceanic fracture zones. In fact, seawater-driven serpentinization of peridotites also involves the alteration of primary Cr-Al-spinels (Fig. 4.7) (Huang et al. 2017), which release Cr and Al in circulating fluids. Chromium (Cr3+ ) can form abundant nano- or microparticles of chromite concealed in serpentine (Huang et al. 2017; Schindler et al. 2017). Subsequent circulation of hydrothermal fluids within the serpentinite bodies can mobilize such chromite particles, hence favoring their agglomeration to form lumps of a few micrometers or more in size and attaining gravitational separation and final deposition into open fractures of the serpentinite country rocks. The aforementioned fracture-filling chromitites could be the precursors of podiform chromitites. In particular, the fact that chromite and serpentine minerals exhibit different behaviors in response to mechanical stresses likely played an essential role in the formation of podiform chromitites. In this respect, it is opportune to recall a petrologic process occurring in the continental crust known as “metamorphic differentiation” (e.g., Mason 1978). Such solid-state “flowing” of the chromite micrograins in the serpentinite matrix favors their further agglomeration, eventually giving rise to fairly large, even exploitable, ore bodies. The aforementioned process is certainly induced by the plicative tectonic events associated with ophiolite obduction on passive continental margins. The latter circumstance may be substantiated by the sheared texture often displayed by worldwide chromitites, although proponents of the deep origin of these ore bodies ascribed their deformations to high-temperature plastic flow in the upper mantle (e.g., Li et al. 2009). Inclusions of highly magnesian (Mg# 95–98) mafic silicates and Ni-rich sulfides, such as millerite and godlevskite, are quite common in the chromite grains from podiform chromitites. Moreover, heazlewoodite and awaruite also generally occur in the matrix of chromitites (e.g., Uysal et al. 2009). The presence of the aforementioned minerals is consistent with a serpentinite parentage. Fig. 4.7 Microphotograph (scanning electron microscope, BSE mode) of part of a thin section of a serpentinized peridotite xenolith from the Hyblean area (Sicily, Italy). Srp = serpentine minerals; Spl = primary peridotite spinel (Cr2 O3 = 19.6 wt%); Chr = micrograins of secondary Fe-rich chromite (Cr2 O3 = 61 wt%). A full explanation is given in the text

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From the aforementioned arguments, it follows that ophiolitic diamonds were likely formed in serpentinite-hosted hydrothermal systems in slow/ultraslowspreading domains of the oceanic lithosphere, far before the geodynamic collisional events that led to ophiolite emplacement. The latter process probably played an important role in the redistribution of microdiamonds in their host rocks, including the podiform chromitites.

4.1.6 Nano- and Micron-Sized Diamond Formation in Hawaiian Salt Lake Crater Xenoliths The discovery of nano- and microdiamonds beneath Hawaii sheds new light on the metasomatic enrichment of the oceanic lithosphere. Nano- and micron-sized diamonds were detected in a mantle-derived garnet pyroxenite xenolith from the Salt Lake Crater (Hawaiian island of Oahu) (Wirth and Rocholl 2003). The depth of xenolith crystallization is estimated to be approximately 50–75 km (1000 °C and 20 kbar). Diamond is the major phase in the glass, ranging in size between a few and several hundred nanometers. Two varieties of nanocrystalline diamonds are identified. Polycrystalline aggregates of nanometer-size diamonds are tentatively explained as pseudomorphs after carbonate, while single crystals are suggested to have crystallized from the melt. Other nanocrystalline phases include native Fe and Cu, FeS, FeS2 , ZnS, AgS, and several Ti, Nb, Zr, Ir, In, Pd-rich phases of unknown composition. Far less abundant than polycrystalline aggregates are single crystals of microdiamond, ranging in size from 20 to 1000 nm. The study of the xenolith documented mid-ocean ridge basalt (MORB)-type noble gas compositions for the abundant secondary fluid and melt inclusions that occur along cracks and fractures within the pyroxenes (Wirth and Rocholl 2003). Therefore, diamond formation is connected with secondary mineralization of the xenoliths. In this sense, the mechanism of their formation is close to the formation of microdiamonds from Tehuitzingo serpentinite (southern Mexico). Diamond precipitated metastably at low pressure from reduced C-O–H fluids that infiltrated from the mid-ocean ridge structures. The following reactions may take place at the boundary of initial carbonates: CaCO3 → CaO + CO2

(4.7)

It follows the formation of free carbon and water-carbon dioxide fluid formation by reaction (4.5) (Wirth and Rocholl 2003; Frezzotti and Peccerillo 2007), which is favorable for metastable diamond formation (see Chap. 2, Fig. 2.4d) and, as a result, to the replacement of carbonate on nanodiamond aggregates. On the other hand, reduced fluids of methane-dihydrogen composition could also be favorable for native elements and metastable diamond formation (see Chap. 2).

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4.1.7 Conclusions The aforementioned results led to the general conclusion that diamonds can be metastably formed at shallow mantle depths in the oceanic lithosphere from fluids related to processes of secondary mineralization. This view calls into question the general notion of an ultrahigh-pressure origin for ophiolitic diamonds.

4.2 Carbonado Genesis The aforementioned data in the previous chapters and sections allow us to shed new light on the problem of carbonado genesis—rare diamond varieties encountered in ancient alluvial deposits nonconnected with kimberlite mineralization. Carbonado is the most enigmatic variety of all diamonds. These cryptocrystalline diamond aggregates lack mantle mineral inclusions and any affinity to kimberlite-clan rocks. There are two types of carbonado that are essentially different in composition and properties: common Brazilian-type carbonado and yakutite, a new variety (Kaminsky 1994). The sintered, polycrystalline microdiamond aggregates have a porous ceramic texture and a melt-like surface patina. Carbonados range up to 3167 carats, and Yakutites range up to 2.2 carats. The color of both varieties differs from light gray to almost black. Light yellow grains are observed only in yakutite. Frequently, the latter has black spots caused by graphite inclusions. Sometimes, finely dispersed graphite inclusions are responsible for a regular black color. Carbonados were found in 1.5-Ga metasedimentary rocks that are localized in Brazil (Bahia) and the Central African Republic (Bangui). Carbonado has been dated by Ozima and Tatsumoto (1997) and Sano et al. (2002) on samples derived from conglomerates that have been reworked over a period from at least 1.7 Ga to approximately 3.8 Ga. Both studies report ages of 2.6–3.8 Ga on implanted radiogenic lead. Yakutites were discovered in alluvial deposits of northern Yakutia. They were previously known as “hexagonal diamonds” (Bundy and Kaspar 1967; Hanneman et al. 1967) or “carbonado-like diamond” (Bartoshinsky et al. 1980). The carbonado is composed of small diamond crystallites. Crystals in carbonado are 30 to 100 μm across; some are much finer, others reach a millimeter in size, and all are embedded in a strongly bonded matrix of microdiamond to nanodiamond. The crystallite shape is octahedral and more rarely cubic, as shown by electron microscopy. Some of the coarser diamonds have defect lamellae and dislocation tangles and mineral inclusions of metals (Fe, Ni, Cr, Ti, Si), metal alloys (Fe–Ni, Fe–Cr, Ni–Cr, and W-Fe–Cr-V), and very unusual minerals, specifically moissanite (SiC) and osbornite (TiN) (Haggerty 2017; De et al. 1998). In contrast, yakutite crystallites are considerably smaller (0.1–1 microns) and more uniform (Kaminsky et al. 1985a). Polycyclic aromatic hydrocarbons (PAR) were encountered at the studding of organic matter, extracted from crushed samples of carbonado and yakutite (Kaminsky et al. 1985b). Carbonado is isotopically light, with δ13 C = –24–-31‰ (Ozima et al. 1991; Shelkov et al. 1997; De et al. 2001). Nitrogen

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concentrations range from 20 to 500 ppm, and δ15 N ranges from -3.6 to 12.8‰ with an average of 3.7‰ (Shelkov et al. 1997; Vicenzi and Heaney 2001; Yokochi et al. 2008). Currently, there are five main schools of their origin (Haggerty 2017): 1. Meteoritic impact; 2. Growth and sintering in the crust or mantle; 3. Subduction; 4. Radioactive ion implantation of carbon substrates; 5. extraterrestrial. Kaminsky (1994) concluded that carbonado formation could occur under low pressure in organic carbonaceous rocks (coals) near radioactive sources. However, no radioactive sources were noted in carbonados. Meanwhile, from Sect. 4.1, it follows that the crystal size, carbon isotopic composition, nitrogen content, association with moissanite, native metals, and metal alloys of the carbonado matrix correspond to serpentinite and ophiolitic diamonds. Osbornite is a very rare mineral that was first discovered in meteorites. In terrestrial rocks, he is known only in the Tibetan Luobusa ophiolite containing microdiamonds (Patil 2019). According to McCollom and Seewald (2013), prior to ~ 3 billion years ago, serpentinization was widespread. It would have provided an abundant source of metabolic energy for the first organisms on ancient Earth. The discovery of microfossils 3.8 billion years old by Pflug and Jaeschke-Boyer (1979) showed that serpentinites could exist during this period of time. It is possible to conclude that carbonado crystals could have formed during the final stage of serpentinization in the primary oceans of Earth from an organic-water system, similar to that occurring in some meteoritic parent bodies.

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Chapter 5

Nanocarbon and Microdiamond Formation in the Lithogenesis and Metamorphic Processes

5.1 Nanocarbon Formation in the Lithogenesis and Contact Metamorphism Processes 5.1.1 Nanodiamond Formation in the Lithogenesis and Contact Metamorphism Processes In the sedimentary rocks, diamonds (500–700 nm in size) were first detected by Dubinchuk et al. (1976). They were identified by TEM on the contacts of organic matter with radioactive centers in coals, lignites, carbonaceous shales, kerogens, and kerites. Later, Nuth and Allen (1992), Kouchi et al. (2005) showed that UV irradiation can stabilize the formation of nanodiamond particles from organic matter at low P–T parameters. Nakano et al. (2002) also predicted that diamonds could be formed in sedimentary rocks, particularly in the kerogen and coal fractions. Such a prediction was later confirmed by Tian et al. (2011), as nanodiamonds were found in association with Younger Dryas Boundary charcoals. Simakov (2015) suggested that nanodiamonds can also occur in some sedimentary rocks bearing different condensed hydrocarbon compounds, including bitumen and asphaltenes. In fact, it is believed that at catagenesis temperatures from 50 to 150 °C, kerogen transforms into bitumen (petroleum + asphalt), and diamondoid formation can take place (Wei et al. 2006). Shale gas mainly consisting of methane also formed in all studies of lithogenesis. The conditions of coal and oil formation at the stages of late catagenesis, metagenesis, and lithogenesis (Fig. 5.1) correspond to the conditions of the Hyblean (Sicily, Southern Italy) serpentinites formation (see 4.1 part; Simakov et al. 2015). Free carbon is formed at these stages due to the decomposition of hydrocarbons, and it can contain nanocarbon particles, as has been demonstrated in the Hyblean case. Nanodiamond formation is possible from shale gas presented at all stages of lithogenesis together with the formation of bitumen and anthracite coals by the reaction of methane destruction:

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2023 S. Simakov et al., Nano and Micro Diamond Formation in Nature, SpringerBriefs in Earth Sciences, https://doi.org/10.1007/978-3-031-43278-1_5

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Fig. 5.1 Scheme of OM reformation in the processes of lithogenesis in accordance with Tissot and Welte (1978)

CH4 → C + 2H2

(5.1)

The presence of UV radiation can catalyze this formation. On the other hand, their formation could be connected with diamondoids. They occur in crude oils and in gas condensates (from 35 to 2075 ppm, respectively; Nekhaev et al. 2010) and are also known to have significant contents in coals with vitrinite reflectances from 1 to 4 Ro (Wei et al. 2006, 2007), which corresponds to the middle and final stages of sediment lithogenesis (Tissot and Welte 1978). Diamondoids generally accumulate in the bitumen of coals and sedimentary rocks during geological processes and are widespread in organic-rich sedimentary rocks. There are three main phases of diamondoid evolution: diamondoid generation (phase I, Ro < 1.1%), diamondoid generation and enrichment (phase II, 1.1% < Ro < 4.0%), and diamondoid destruction (phase III, Ro > 4.0%) (Wei et al. 2006). Their transformation to nanodiamonds under P–T conditions of different lithogenesis processes is certainly possible (Dahl et al. 2003). Gebbiea et al. (2018) showed that the presence of diamondoid molecules in CVD syntheses at 750 °C and low pressure from methane leads to the formation of postcritical diamond nuclei consisting of 26 carbon atoms only. In coals, diamondoids

5.1 Nanocarbon Formation in the Lithogenesis and Contact Metamorphism …

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are stable up to 550 °C, and their decomposition becomes significant at temperatures higher than 600 °C and leads to the formation of 15% aromatics and 83% gaseous products (Wei et al. 2006). Decomposition is also possible at lower temperatures from 150° to 450 °C in the presence of clays, metal catalysts (Ni, Al, Cr, Pt, and Pd), and UV radiation (Bagrii 1989). The presence of UV radiation and clays, such as Ni, Al, Cr, Pt, Pd, and U is known in coals (Dai et al. 2022; Vyalov and Nastavkin 2019; Vyalov et al. 2021). There are several hypotheses on the formation of different types of coal in nature (e.g., Vyalov et al. 1998). It is possible to distinguish the main schools of thought. The first is connected with regional metamorphism. The Moscow, Donetsk, Kuznetsk, and Gorlovsky Basins represent suitable case studies. Coal metamorphism increases as the thickness of coal-bearing deposits increases, and thence, the most metamorphosed anthracites were formed at depths of up to 10–14 km. The second school of thought suggests that the coal formation is genetically connected with contact metamorphism with igneous intrusions in different geological times (mainly on the border of the Permian and Triassic and in Triassic periods). Taimyr and Tunguska anthracites and graphite coal, Karelian shungites, and many others correspond to this type. As a result, coals of various genetic types can coexist in the same coal basins (for example, Kuzbass and Gorlovsky). From the aforementioned data, it follows that nanodiamonds could have formed during the different stages of diagenesis and contact metamorphism. To verify this assumption, Raman observations were performed for the Moscow brown coals, Donbass and Kuzbass antracite, and Taimyr coals samples. It is known that the application of Raman spectroscopy in the characterization of the inorganic composition of coal is considered a real possibility by Potgieter-Vermaak et al. (2011). Raman (RS) and photoluminescence (PL) spectra were recorded on an “inVia Raman microscope” (Renishaw) at room temperature at the Lebedev Physics Institute in Moscow. Raman microscope attachment has excitation emissions of 785 nm and 1064 nm, which strongly reduced the intensity of the luminescent background in the spectra of the samples. A laser excitation source at 785 nm was applied, and the laser power was controlled by means of a series of density filters to avoid heating effects from 100% do 5 × 10–8 %. Excitation light was focused to a spot diameter of 2–5 μm. The absence of the laser heating effect on the sample was controlled through a microscope. The multichannel system of registration controlled a satisfactory signal/noise ratio. The resolution of the device in the entire spectral region was ~1 cm−1 . The Moscow basin is located in the central part of the Easter-European platform. His spreading area is nearly 120,000 km2 . It contains 12 coal-bearing regions and 30 areas and deposits. The age of sedimentary layers corresponds to the Lower Carbon (C1v and C1t ). The thickness of coal-bearing strata is up to 180 m. Brown coals were formed at depths from 110 to 200 m, and their metamorphism corresponds to upper diagenesis and low catagenesis with temperatures of 50–70 °C. The average contents of natural humidity, coal ash, and sulfur are equal to 32, 31, and 3–5%, respectively. The studied sample of brown coal from the Gryzlovsky deposit has an

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

Ad of 4% and an Std of 3%, and its vitrinte reflectances (Ro) correspond to 0.43– 0.47% (Fig. 5.2a). In the range of the basin, there are sands with carbonaceous plant detritus containing U–Mo–Re mineralization (Vikentyev and Kailachakov 2020). Raman and photoluminescence spectra recorded on the aforementioned “Renishaw inVia Raman” microscope showed peaks of sp2 hybridization at 1588 cm–1 and peaks of sp3 hybridization at 1323 cm–1 . The ~1323 cm−1 bands can be attributed to nanodiamonds (Fig. 5.3). The Donetsk coal basin (Donbass) is located on the slopes of the Ukrainian and Voronezh crystalline shields. Its length reaches 650 km with a maximum width of 200 km. The thickness of sedimentary layers of the lower, middle, and upper Carboniferous reaches 5–18 km in the central part of the structure. The peak metamorphism of sedimentary rocks corresponds to greenschist facies. Approximately, 310 coal seams are distributed in 5 lower Carboniferous, 7 middle Carboniferous, and 3 upper Carboniferous formations. The thickness of coal seams varies widely from 0.5 to 2.5 m. Coals are mostly humus vitrinite. The average content of natural coal ash ranges from 7 to 20%. In the Donetsk coal basin, there are all the main types of coal, from subbituminous to anthracite, which were formed during regional metamorphism (without the visible influence of igneous rocks) during the subsidence of the coal-bearing strata of the so-called “Donetsk aulacogen.” According to Levenshtein (1962), anthracites were mainly formed in the interval temperatures of ~190–350 °C at depths from 6.5 to 8 km, which corresponds to a pressure of approximately 1.5–2 kbar. Meta-anthracites, with a reflection index of more than 5%, formed at depths of up to 10 km or more. According to Vyalov (1998), anthracites of regional metamorphism formed at temperatures up to 400 °C. The studied anthracite of the Donetsk basin is from the Victoria mine of the C24 coal-bearing suite, from the Dronovsky coal seam. Coal contains 53% inertinite and 47% vitrinite. Anthracite contains approximately 1% pelitomorphic clay matter (Fig. 5.2b). The sample was formed at 200– 250 °C and has a vitrinite reflectance (Ro) of 3.23%. Raman and photoluminescence spectra recorded on the aforementioned “Renishaw inVia Raman” microscope showed peaks of sp2 hybridization at 1596 cm–1 and peaks of sp3 hybridization at 1326 cm–1 (Fig. 5.3). The Gorlovsky Coal Basin is located in Western Siberia near the Ob River. It extends 120 km in the northeastern direction with an average width of 1.5–7.5 km. It contains 55 coal seams and interlayers with thicknesses from 10–14 to 26–41 m. The coal-bearing Balakhonskaya Series has ages from the middle Carboniferous to early Permian. The coals of the basin are represented by anthracites of high quality due to their low-ash, low-sulfur, and high-carbon contents. Dynamometamorphism probably played a significant role in the formation of Gorlovsky anthracites. The regional metamorphism is recorded by the values of vitrinite reflectance (Ro%) of the stratigraphically underlying anthracite strata higher than those of the upper strata. The presence of olivine dolerite sills in the Gorlovsky quarry also indicates a possible thermal metamorphism contribution in the coal formations. The temperature of anthracite formation is estimated to be in the range of 300–350 °C (Vyalov et al. 1998). Their Ro varies from 3.16 to 5.09%, with an average of 4.67% (Vyalov 1998). The organic matter of the studied Gorlovsky anthracite sample of 40–60% consists

5.1 Nanocarbon Formation in the Lithogenesis and Contact Metamorphism …

65

Fig. 5.2 Microphotographs of Moscow brown coal with telenite and corpovitrinite (80x) (a); Taimyr meta-anthracite in reflected light and immersion with vitrodentrinite, sporinite, and inertodentrinite (365x) (b); Taimyr coal graphite (770x) (c); Donbass antracite in reflected light with main vitrinite–collinite gray-light groundmass with fragments of intertinite in left part (100x) (d); Gorlovsky antracite with semifusinite and inertodetrenite (270x) (e); shungite “shiny” (270x) (f)

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

Fig. 5.3 RS spectra of the samples obtained using a spectrometer at the excitation of 785 nm: 1330 cm–1 indicating the presence of nanodiamonds; 1600 cm–1 related to nanographite. 1—coal graphite from Taimyr, 2—shungite, 3—Donbass anthracite, 4—Gorlovsky antracite, 5—anthraxolite from Taimyr, 6—Moscow brown coal

of inertinite macerals composed mainly of semifusinite and funginite (Fig. 5.2c). It showed micro-Raman peaks of sp2 hybridization at 1600 cm–1 corresponding to nanographite and peaks of sp3 hybridization at 1320 cm–1 (Fig. 5.3). The Taimyr Coal Basin is situated in the northern part of Russia on the Taimyr Peninsula. It extends 1000–1100 km in the sublatitudinal direction with an average width of 100–150 km from Yenisei Bay to the Laptev Sea coast. The basin has a Permian age. In the western part of the basin, coking coals prevail, while in the eastern part, weakly caking and lean coals prevail. Anthracites are developed on the thermal contact with igneous rocks and are also related to coal graphites. Seregen meta-anthracite deposits were formed as a result of intensive contact metamorphism of coal on the boundary of trap intrusions. Similarly, anthraxolites and coal (carbon) graphite were formed on the boundary with dolerite and gabbro–dolerite intrusions in the Triassic period. The formation temperatures of these meta-anthracites and coal graphites were estimated as 500–600° and 600–650 °C, respectively (Amarsky and Vyalov 1995; Vyalov 1996, 1998). The studied meta-anthracite sample is composed of 84% vitrinite and 16% intertinite (Fig. 5.2d). Its organic matter contains 96% carbon, 2% sulfur, and 2% hydrogen. The coal graphite sample is composed of crystallites 2–4 μm in size (Fig. 5.2e) derived from partial crystallization of amorphous organic matter. It contains 1% sulfur. Its organic matter fraction contains 97.5% carbon and 0.5% hydrogen. Both samples have a high index of reflectivity: metaanthracite up to 6.54% and coal graphite up to 7.74%. Using Raman spectroscopy, we studied the samples of anthraxolite and graphite coal (Fig. 5.3). The graphite sample showed peaks of sp2 hybridization at 1583 and 1614 cm–1 , which correspond to

5.1 Nanocarbon Formation in the Lithogenesis and Contact Metamorphism …

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nanographite, and peaks of sp3 hybridization at 1313 cm–1 . The anthraxolite sample showed peaks of sp2 hybridization at 1602 cm–1 , which corresponds to nanographite, and peaks of sp3 hybridization at 1318 cm–1 (Fig. 5.3). In summary, peaks at 1600 cm–1 (and more), corresponding to nanographite, were recorded in the Taymyr and Gorlovsky samples. The peaks at 1313–1326 cm–1 corresponding to ultrananocrystalline diamonds were recorded in all studied samples of coals.

5.1.2 Fullerene and Fullerene-Like Phase Formation During Metamorphic Processes Currently, fullerenes and fullerene-like phases (fullerites) occur in different geological situations, such as in the matrix of carbonaceous chondrite meteorites (Buseck and Hua 1993), in thin films within fractures in Karelian shungites (Buseck et al. 1992), in pillow lavas that cut across black shales (Jehlicka et al. 2000), in coals (Fang and Wong 1997; Osawa 1999), in fossil dinosaur eggs (Wang et al. 1998), and in a carbonaceous impact breccia from the Canadian Sudbury Structure (Becker et al. 1994, 1996). Some authors suppose that fullerenes can form complexes with metals in endogenic fluids (Vinokurov et al. 1997). These complexes can transport metals and take part in the kimberlite and carbonatite formations. Despite many reports of natural fullerene findings, their origin remains elusive (Buseck 2002). Becker et al. (2001) state that fullerenes are “highly resistant to metamorphism” and “a robust tracer in the geological environment.” The most typical fullerene association in natural objects is association with graphite. There is also report of larger natural fullerenes, sometimes called “giant fullerenes” (Smith and Buseck 1981). These are large and closely resemble the concentric carbon layers in the “carbon onions” that are presumed to be related to fullerenes (Ugarte 1992; Kroto 1992). To date, the main occurrence of fullerenes has been found in shungite, a carbonrich rock in Karelia, near Lake Onega, northwestern Russia. It is a Palaeoproterozoic metasedementary rock of volcanic and metosomatic origin. Shungites are represented by four major rock types: tuffs, siltstones, dolostones, and cherts. The estimated age of the shungite deposition is ca. 2000 Ma (Melezhik et al. 1999). Their chemical composition is complex, as all these rock types may be mixed in different proportions. Shungite rocks were later affected by contact metamorphism induced by the intrusion of dolerite sills and by greenschist facies regional metamorphism related to the 1.8 Ga Svecofennian orogeny, resulting in a complex metamorphic history. The paragenetic chlorite–actinolite–epidote assemblage reflects a temperature of 300– 350 °C. More generally, the Palaeoproterozoic succession of the northern Lake Onega area includes seven formations: Palozerskaya, Jangozerskaya, Medvezhegorskaya, Tulomozerskaya, Zaonezhskaya, Suisarskaya, and Vashezerskaya (Sokolov 1987). Currently, geologists still debate the origin of shungites. There are several different

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

genetic concepts for their origin. One view suggests an abiological nature of shungite with a major source of carbon coming from volcanogenic magmatic systems, e.g., (Sokolov and Kalinin 1975). A biogenic hypothesis was advanced by Volkova and Bogdanova (1986) and was very specific, describing shungite as a product of the accumulation of higher plants. They characterized shungite as coal. Timofeev (1924) reported a metamorphic origin for Karelian shungite. Shungite “shiny” contains 97.5% carbon (in kerogen), 0.4% sulfur, and 1.3% ash and is characterized by a high index of reflectivity of 7.2%. His molecular structure is close to the structure of industrial cokes and some types of glassy carbon (Fig. 5.2e). Chazhengina and Kovalevski (2013) studied shungitwe samples on a dispersive Nicolet Almega XR Raman spectrometer with a green laser (532 nm, Nd-YAG). The shungite Raman spectra showed two well-resolved bands: D1 at approximately 1350 cm−1 and G band at approximately 1580 cm−1 . In summary, they concluded that shungite formation was due to partial crystallization of the sapropele organic substance under contact-thermal metamorphism at 400–420 °C, or even higher, since the presence in them of unusual microstructure known as “bucky balls,” which are believed to be formed at higher temperatures than the above. Raman (RS) and photoluminescence (PL) spectra for the shungite sample were recorded on a “Renishaw inVia Raman” microscope. The sample showed peaks of sp2 hybridization at 1600 cm–1 , which corresponds to nanographite, and peaks of sp3 hybridization at 1311 cm–1 (Fig. 5.3). The ~1311 cm−1 bands can be attributed to nanodiamonds. The HRTEM images of shungite are similar both to kerogen heated above 2000 °C (Jehlicka and Rouzaud 1993) and to natural and synthetic cokes (Kovalevski et al. 2001), although these C species are formed in different environments. In drillhole 2, the thermal effect of dolerite sills can be traced by variations in the microstructure. Chazhengina and Kovalevski (2013) detected an unusual microstructure in shungite at the contact of a dolerite sill (Fig. 5.4a). It corresponds to hollow “bucky balls” or fullerene-like structures (Chhowalla et al. 1997, Fig. 2). These fullerene-like structures correspond to fullerene “bucky balls” synthesized in experiments on the walls of carbon nanotubes at 700–750 °C and 500 MPa (see Chap. 3, Fig. 3.5 and 3.6). The slightly elongated structures oriented along one direction of carbon lay short with bent stacks (or fractions of closed shells) of the Maksovo shungite sample (Avdeev et al. 2006, Fig. 4) correspond to fullerene-like “onion” formations synthesized in the aforementioned experiments (see Chap. 3, Fig. 3.5b, c). Some onion-like structures are very close to natural “giant fullerenes” found in the Allende meteorite (Smith and Buseck 1981) and to onion-like graphitic particles observed in wood charcoal (see Fig. 3.7 in Chap. 3). The same onion-like particles associate with diamond structures in the wood charcoals carbonatized at 700 °C (Fig. 5.4b). On the other hand, the P–T parameters of the experiments described in Chap. 3 (700–750 °C) correspond to the P–T of contact metamorphism and temperatures of coke formation. Therefore, the model of fullerene formation in Karelian shungites at the contact metamorphism on the boundary of carbonaceous sedimentary rocks and dolerite sills is more likely than the hypothesis of their formation at greenschist facies metamorphism or at accumulation of higher plants.

5.1 Nanocarbon Formation in the Lithogenesis and Contact Metamorphism …

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Fig. 5.4 Microstructure of unusual hyper fullerene-like microstructure of shungite in accordance with Chazhengina and Kovalevski (2013) (a) and in the wood charcoals carbonatized at 700 °C in accordance with Ishimaru et al. (2001) (b)

“Renishaw inVia”—Raman spectra (RS) and experimental data reported in Chap. 3 allow us to shed new light on the problem of fullerene genesis. We can conclude that rare forms of carbon-multiwall fullerites and nanotubes could be formed from hydrocarbon gases at P–T parameters corresponding to contact metamorphism conditions.

5.1.3 Conclusions In conclusion, the formation of nanocarbon phases is possible at all stages of lithogenesis (diagenesis, catagenesis, and metagenesis). The formation of nanosized diamond particles can be connected here with diamondoids and shale gas. UV radiation, clays, such metals as Ni, Al, Cr, Pt, and Pd, can catalyze their formation. Nanodiamond formation due to contact metamorphism episodes connected with intensive diamondoid and shale gas breakdown at temperatures greater than 600 °C is more likely. Our study also showed an inverse correlation between the Raman sp3 peak and the coal formation temperature (Fig. 5.5a). This corresponds to Beyssac et al.’s (2002) suggestion that the Raman spectrum of carbonaceous material can be used as a geothermometer for regional metamorphism. Using the results of Koniakhin et al. (2018), it is possible to roughly estimate the size of nanodiamonds in the coals and shungites (Fig. 5.5b). Unfortunately, the shape of the Raman peak is the result of lattice defects, laser heating, and other factors, which makes it difficult to precisely estimate the particle size. For example, the Raman and HRTEM data of Hyblean serpentimites show the uncertainty of this method (see Sect. 4.1). Meanwhile, the obtained Raman sp3 peaks of coals and shungites mainly correspond to those of

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Fig. 5.5 Correlations of RS spectra of the samples a and nanodiamond sizes determined by the method of Koniakhin et al. (2018) b with the temperatures of the sample formation

carbon phases from the Hyblean serpentinites (lower than 1332 cm–1 ). Therefore, the presence of ultrananocrystalline diamonds is possible. On the above grounds (see Chap. 3), it is possible to conclude that the studied (Karelian) shungites consist of a complex of various carbon structures formed at temperatures as high as 700 °C and more in the temperature range of contact metamorphism (pyrometamprphism).

5.2 Micron-Sized Diamond Formation During Metamorphic Processes 5.2.1 Worldwide Metamorphic Diamonds Micron-sized diamonds have been found in shallow metamorphic rocks, such as in rocks from Kokchetav (Sobolev and Shatsky 1990), Dabie Shan (Xu et al. 1992), and northern Qaidam, China (Yang et al. 2003); the Western Gneiss Region, Norway (Dobrzhinetskaya et al. 1995; Van Roermund et al. 2002); Erzgebirge, Germany (Massonne 1999, 2003; Stockhert et al. 2001; Dobrzhinetskaya et al. 2006a); Sulawesi, Indonesia (Parkinson and Katayama 1999); northern Greece (Mposkos 2003; Mposkos and Kostopoulos 2001; Beyssac and Chopin 2003); southwest Japan (Nishiyama et al. 2020) and the Southern Urals (Bostick et al. 2003). Their carbon isotopic composition varies from −3 to −30%o of δ 13 C (the main part of analyzes lies in the range from 0 to −25%o ) and has a maximum at −10%o (Cartigny 2005; Cartigny et al. 2001; De Corte et al. 1998; Lavrova et al. 1999). Their nitrogen content varies from 0 to 11,150 ppm (the main part of the analyzes lies in the range from 0 to 8000 ppm) with a prominent mode at ~1800 ppm. The nitrogen isotopic composition corresponds to metasediments. It varies from −1.8 to 12%o of δ 15 N, displaying a

5.2 Micron-Sized Diamond Formation During Metamorphic Processes

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maximum at 12%o (Cartigny 2005). Rare UHP minerals have also been found in these rocks. The crustal metamorphic microndiamonds may contain minerals (such as graphite or carbonates) and fluid inclusions (De Corte et al. 1999).

5.2.2 Kokchetav Diamond Deposit To date, the main known occurrence of metamorphic microdiamonds is connected with the Kokchetav massive (Northern Kazakhstan). In 1967, several small yellowishgreen cubic diamond crystals were found in Palaeogene Ti-Zr placers near the northern margin of the Kokchetav Massif (Kashkarov and Polkanov 1972). Later, in the 1970s, diamonds were discovered in the weathering profile of carbonate– silicate rocks in the shoreline of Kumdykol Lake (Kumdykol), garnet-biotite graphite-bearing gneisses and a variety of other rocks, from ultramafic (e.g., gametpyroxenites) to felsitic rocks. Exploration and feasibility studies occurred during 1980–86. In 1980–86, the diamond reserves were estimated at approximately 2.5 billion carats at an average content of approximately 20 cr/t. In 1990, diamond occurrence was found in Barchi, approximately 15 km NW of the Kumdykol deposit. Both deposit areas are comparable in their geological setting and diamond potential. The Kumdykol deposit and the Barchi occurrence are located in the western part of a metamorphic belt that outcrops at the center of the Kokchetav Massif. The latter is a large (180 km × 150 km) Precambrian structure located amidst Palaeozoic structures of the Central Asiatic Fold Belt (Fig. 5.6). This massif was considered either a fragment of the Kazakhstan—Tien Shan microcontinent or a pre-Riphean continental block ruptured from the marginal part of Proto-Gondwana in the Late Riphean (Mossakovsky et al. 1993). In most recent publications, the Kokchetav Massif is considered a subduction-collision zone that underwent a multistage Vendian— Early Ordovician geodynamic evolution (Dobretsov et al. 2006). Within the massif, Precambrian formations occur in three blocks: Kokchetavsky, Zagradovsky, and Shatsky. All known diamondiferous bodies lie within the Kokchetavsky Block. The Kokchetav massif (Fig. 5.7) is composed of several Precambrian rock series, island arc-related Cambrian to Ordovician volcanic and sedimentary rocks, Devonian volcanic molasse, and Carboniferous ± Triassic shallow-water and lacustrine deposits. All of these rocks were intruded by granites of varying ages. The central part of the massif, in which eclogites and diamondiferous rocks occur, is composed of the Zerenda rock series, whose age is supposed to be Pre-Riphean (Kushev and Vinogradov 1978; Schertl and Sobolev 2013). None of the eclogites contain diamonds. The basal eclogite-gneissic-shale member, a component of the Zeredinskaya Series, makes up the basement of the Kokchetav Massif. This member consists of biotite, garnet-biotite, and cordierite-garnet-biotite gneissic rocks and high-aluminous garnet-kyanite-muscovite shale hosting eclogite and amphibolite

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bodies, with minor marble. The rocks are metamorphosed to amphibolite and granulite facies. Finds of the UHP and HP indicator minerals (diamond, K-rich clinopyroxene, and coesite) in metamorphic rocks gave a reason to propose a UHP metamorphic grade for this deposit (Sobolev and Shatsky 1990; Sobolev et al. 1991). The age of the UHP metamorphic event (as estimated using zircon and garnet-pyroxene pairs) is 527–537 Ma (Claoue-Long et al. 1991; Jagoutz et al. 1990; Hermann et al. 2006). The oldest basement structures of the Kokchetav Massif are granite-gneissic domes with Palaeozoic granites in their cores and the Zeredinskaya Series rocks along the margins. Middle and Late Riphean platform cover deposits discordantly overlie these rocks. The platform complex includes phyllite-like carbonaceous shales with limestone intercalations and quartzite formations. The alkaline mafic/ultramafic rocks (syenite, biotite pyroxenite, and carbonatite) framing the Kumdykol deposit are dated at 570 ± 50 Ma (Sm/Nd; Lavrova et al. 1997). The youngest structures of the Kokchetav Massif are Devonian to Early Carboniferous superimposed troughs and graben synclines filled with rhyolite porphyry and a coal-bearing molasse red-bed sequence. Many of the abovementioned alkaline mafic/ultramafic intrusive bodies are confined to long-lived, deep-seated faults, mostly identified from geophysical data. The existence of these faults (at least since Early Riphean), in many respects, governed the essentially linear-block pattern of the study area. Diamondiferous rock exposures tend to be located at marginal parts of the UHP units, adjoining either other domains (e.g., Barchi) or low-pressure areas (Kumdykol). Within the latter domain, diamondiferous rocks follow a general NE– SW strike as a narrow patch 1-km long. The widths of their outcrops range from 250 m in the NE part to 50 m in the southwest. Contoring of diamondiferous rocks at Barchi is incomplete to date, but a linearity of the outcrop pattern has also been observed here. At Kumdykol, diamondiferous rocks occupy 0.2 km2 of the total 90 km2 areal extent of the UHP rocks, i.e., ca. 0.16%. In the Barchi area, the proportion is roughly the same. Despite detailed geological prospecting, no other diamondiferous objects have been found within the metamorphic belt. Thus, diamondiferous bodies are highly localized. Figure 5.8 shows a major productive zone of the deposit. Outside, the content of diamonds sharply decreases. The northwestern boundary of the zone follows a belt of granite injections (presumably a marginal part of garnet-muscovite granite body). Here, numerous garnet-muscovite injections penetrate the gneissose matrix. The thickness of individual bodies ranges from several tens of centimeters to several meters (Fig. 5.9a). The thickness of these injections increases in the NW direction. Migmatite bodies limit the productive zone from SE. Banding in migmatite is distinct due to alternation of thin leuco- and melanocratic bands (Fig. 5.10b). Major minerals are garnet, biotite, and quartz-feldspar forming an aggregate with zoisite and epidote as admixtures. Gneiss makes up 70–80% of the productive zone, and in its SW part, gneiss becomes the only rock variety. Gneissosity in unaltered rock varieties is quite distinct (Fig. 5.9c). It strikes 25–35° and dips at 65–85°. Major minerals are garnet, biotite,

5.2 Micron-Sized Diamond Formation During Metamorphic Processes

73

Fig. 5.6 Structural elements of the Central Asiatic Foldbelt (after Mossakovsky et al. 1993) (a), location of the Kokchetav Massif and the Kumdykol and Barchi deposits (b). 1—Microcontinents: K—Kokchetav, U—Ulutau, Am—Aktau-Mointi, I—Iliisk, NT—North Tien Shan, MT—Median Tien Shan, D—Dzhungar; 2—Accretion-related (mosaic) fold systems of Caledonian age: Bn—Baikonur, Kt—Karatau, DN—Dzhalair-Naiman, As—Atasu, Tt—Tekturmas, Sp— Stepnyaksk, B—Boshekul, E—Erementau, Ch—Chinghiz; residual or superimposed basins: 3— Caledonian: A—Aghadyr, Ach—Anuisk-Chuisk; 4—Variscan: Pb—Pribalkhash; 5—Linear fold systems (Variscan): TA—Turkestan-Alai, Z—Zaisansk; 6—Major faults; 7—Central part of the Kokchetav metamorphic belt

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Fig. 5.7 Geological scheme of the central part of the Kokchetav metamorphic belt (Dobretsov et al. 1998). 1—Northern domain (Vendian-Cambrian island arc); 2—Kokchetav microcontinent domain; 3—White lake domain (high-pressure/low-temperature equivalents of HP members in megamelange domain); 4—Granite dome domain (Late Ordovician granites, Devonian granites, and grano-syenites); 5–7—tectonic units of megamelange domain: 5—Ultrahigh pressure/HP unit with high temperature (HT) eclogites, 6—medium pressure (MP) unit with Al-rich metasediments, 7—Low pressure (LP) unit (Daulet); 8—Diamond-bearing rocks

muscovite, feldspar, and quartz. Mesocratic gneiss varieties additionally contain clinopyroxene, zoisite, and amphibole (accounting for 50; 5—adit horizon; 6—drillholes; 7—faults (reproduced from Fig. 4 of Pechnikov and Kaminsky 2008)

frequently blasto-mylonitic or blasto-cataclastic. In the latter case, porphyry clasts of garnet and feldspar are surrounded by narrow bands of fine-grained quartz and small mica scales, either replaced with chlorite or occurring in a chlorite-sericite-carbonate matrix. In intensely altered varieties, chloritization becomes typical of garnet. The mineral neoformations are zoisite, epidote, titanite, and rutile. Metasomatites carry veinlets and disseminations of pyrite and graphite. Graphite particles are present both in the matrix and in garnet porphyroblasts (most frequently along fissures). The thickness of metasomatic bodies ranges from several meters to several tens of meters. Observations of coesite microinclusions in zircons from garnet-biotite gneiss, along with clinopyroxene (both ordinary and potassic) and titanite (highly aluminous) microinclusions in garnets from some rocks (Sobolev et al. 1991, 1994; Claoue-Long et al. 1991; Sobolev 2006), have generated considerable interest among researchers. After the first discovery of coesite in the Kumdykol deposit (Sobolev et al. 1991), its intergrowth with diamond was confirmed by Raman spectroscopy (Sobolev et al. 1994). Later, coesite was identified as an inclusion in zircon in eclogitic rock from the Barchi occurrence (Korsakov et al. 1998) and as inclusions in garnet from mica schists of the Kulet occurrence (Shatsky et al. 1998b). In addition to coesite, Mg-Ca

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Table 5.1 Average chemical composition of diamondiferous rocks from the Kumdykol deposit (wt.%) Component

SiO2 TiO2 A12 O3

Rocktype Garnet-biotite gneiss (n = 12)

Quartzose rock (n = 11)

Silicate-carbonate rock (n = 12)

Garnet-pyroxene rock (n = 5)

65.56

74.59

16.97

48.02

0.58

0.28

0.12

0.56

12.82

10.69

2.72

9.12

Fe2 O3

1.42

1.35

1.21

4.13

FeO

4.83

2.18

2.14

8.55

MnO

0.26

0.11

0.23

0.56

MgO

4.39

2.59

15.25

6.85

CaO

2.48

2.40

28.90

18.45

Na2 O

0.83

0.48

0.20

0.70

K2 O

3.02

2.90

0.73

0.19

H2 O

2.64

1.58

2.35

0.00

CO2

0.20

0.06

28.17

1.48

P2 O5

0.16

0.58

0.14

0.06

C

0.26

0.35

0.14

–

(pyrope-grossular), low-Fe garnet is also considered a relict UHP mineral (Sobolev et al. 2001, 2007). The spatial distribution of diamonds has no distinct lithological control. Generally, the economic diamondiferous zones display roughly pod-like patterns with long axes striking SW, and linear diamond-rich zones alternate with barren zones (Fig. 5.10). Diamondiferous linear zones are largely confined to fault-related metasomatite with sulfides and graphite. Compositions of economic diamondiferous rocks vary from silicate to essentially carbonate varieties (Table 5.1).

5.2.3 Kokchetav Diamond-Bearing Crustal Rocks—An Overview Diamonds were first described in gamet-pyroxene rocks and garnet-biotite gneiss only as microinclusions in garnet and zircon, as well as admixtures in pseudomorphs after garnet, i.e., chlorite-sericite aggregates (Sobolev and Shatsky 1987). More recently, diamonds were found in biotite and quartz from garnet-biotite gneiss, in the chlorite-tremolite-actinolite matrix of quartzose rocks, in hornfels from pyroxeneamphibole-feldspar-quartz varieties, and in clinopyroxene, garnet and phlogopite pyroxene-phlogopite-carbonate rocks (Ekimova el al. 1992, 1994; Dobrzhinetskaya

5.2 Micron-Sized Diamond Formation During Metamorphic Processes

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el al. 1994; Lavrova et al. 1999; Ogasawara et al. 2000). Relatively large (200– 300 μm) crystals were found in the intergranular space of gamet-pyroxene rocks (Korsakov and Shatsky 2004), in quartz-feldspar aggregates (Korsakov et al. 2004), and in fissure-filling carbonates (Lavrova et al. 1999). Diamonds in metamorphic rocks occur in a variety of textural contexts and mineralogical assemblages (Fig. 5.11). Clustering of diamond micrograins occurs even at the thin-section scale. Intergrowths of two or more microcrystals are also possible. There are reports of oriented linear chains of 3–6 crystals along microfissures of the rock (Pechnikov and Kaminsky 2008). Microcrystals of 5–10 μm have the form of 2D patterns; 3D botryoidal aggregates were also observed (Fig. 5.11b). From thin-section studies, it follows that the frequency of diamond occurrence in garnet is higher than that in other minerals. Garnet has the highest diamond contents, up to one thousand and more microdiamonds per grain (Lavrova et al. 1999; Yoshioka et al. 2001; Ogasawara 2005). Intergranular chlorite-sericite or chlorite-sericite-carbonate filling is another “diamond-friendly” environment, irrespective of whether it replaces garnet. Diamonds often form intergrowths with graphite, coesite, clinopyroxene, rutile, titanite, kyanite, KFS, biotite, phengite, and carbonates (Shatsky and Sobolev 1993). The most frequent combination is graphite + diamond. The host minerals contain numerous complex inclusions of 1–3 diamonds as intergrowths with quartz, phengite, phlogopite, albite, KFS, rutile, apatite, and titanite. biotite, chlorite, and graphite. The TEM results demonstrated that diamonds contained numerous oxide, carbonate, and silicate nanoparticles (Dobrzhinetskaya et al. 2001, 2003, 2006b). Using the TEM techniques, Langenhorst (2003) observed diamonds coated with thin (ca. 100 nm) chlorite layers. The most complicated and variegated nano- and microinclusions were spotted in garnet and zircon from felsic gneiss (Dobrzhinetskaya et al. 2003).

5.2.4 Mineralogical Features of Kokchetav Metamorphic Diamonds Morphology. The following morphological varieties are distinguishable among the diamonds from metamorphic rocks: octahedra, cubes, cube-octahedron combinations, and skeletal and spheroid crystals (see Fig. 5.12). The last two varieties are considered “metamorphic” diamonds only. No rhombic-dodecahedral crystals typical of diamonds from kimberlites and lamproites were found (Imamura et al. 2002; Ogasawara 2005). Some regularity is traceable in the diamond morphology by the hosting rock type (Shatsky et al. 1998a; Lavrova et al. 1999). Cubes and combined skeleton crystals are most typical of gamet-biotite gneiss. Quartzose varieties additionally carry the smallest (several microns) and most imperfect crystals (rosettes, star-like, and spheroid species), whereas garnet-pyroxene varieties host the largest (100–500 μm) cubic crystals. Most of the diamonds studied are grayish-green or yellowish-green,

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

Fig. 5.11 Photomicrographs showing interrelationships between rock-forming minerals and microdiamonds, Kumdykol deposit. a Altered biotite gneiss: diamonds in chlorite-sericite aggregates (thin section 33/26); b pyroxene-carbonate rock: accumulation of diamond crystals in garnet (thin section 430-3a); c amphibole-biotite gneiss rich in graphite and sulfide: diamond in amphibole (thin section 31/39); d the same rock: diamond in quartz (thin section 31/39); e garnet-biotite gneiss: diamond in biotite (thin section 33/26 k); f pyroxene-carbonate rock: diamonds in carbonate agregate (thin section 430-3a); g altered garnet-biotite gneiss rich in graphite and sulfide: diamond inclusions in carbonate and chlorite-sericite fissure, common appearance (thin section 38/20); g the same thin section detailed fragment: diamond inclusions in carbonate grain and chlorite-sericite aggregate. a–f bar scale is 30 μm on these photomicrographs, g—bar scale is 0.2 mm, g' —bar scale is 10 μm. Amph—amphibole, Bi—biotite, Car—carbonate, Chl—chlorite, Cpx—clinopyroxene, Dia—diamond, Fls—feldspar, Gar—garnet, Gr—graphite, S—sulfide, Ser—sericite, Q—quartz (reproduced from Fig. 6 of Pechnikov and Kaminsky 2008)

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Fig. 5.12 Specific types of microdiamonds. a—Skeletal crystals in a garnet-biotite gneiss; b— Spheroids in silikate-carbonate rocks; c—“Star-shaped” type microdiamonds in dolomite marble (Ogasawara, 2005). The bar scale is 10 μm on all photographs

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

pale to dark-colored, with a minor proportion of milky white, gray, and colorless (octahedral only) crystals. De Corte et al. (1998, 2000) found that some diamonds from gamet-clinopyroxene rocks contained water (OH and H2 O absorptions), carbonate, and graphite inclusions. In addition, a highly potassic C–H–O fluid with dissolved phosphate, chloride, and sulfate/sulfide was identified as nanometer-size inclusions in microdiamonds from dolomite marble, garnet-quartz-pyroxene rock and gneisses (Hwang et al. 2005, 2006). Metamorphic diamonds have a wide range of δ 13 C values, from −8.9 to −27%o , and δ l5 N values from +5.3 to +25%o , similar to those of eclogite-hosted diamonds in the kimberlites (Fig. 5.13). They are also similar to microdiamonds from the Akluilak minette dykes (Cartigny et al. 2004). The formation of methylene (=CH2 ) and methyl (–CH3 ) groups detected in the Kokchetave-diamonds shows the presence of organic matter at their formation (Khachatryan and Baryshev 2022). As shown in Chap. 4, organic matter formation takes place during secondary alteration and sulfidization. Diamonds from different metamorphic rock types differ in their carbon isotopic patterns (Pechnikov et al. 1993). As noted above, diamonds from gneiss have “lighter” isotope compositions relative to the pyroxene-carbonate and gametpyroxene rocks. Table 5.2 summarizes carbon isotope data of diamonds from various sources (Pechnikov et al. 1993; De Corte et al. 1998; Lavrova et al. 1999; Cartigny et al. 2001) and compares them to data of diamonds and associated graphite from the Kumdykol deposit. Since the average size of the diamond crystals is ca. 30 μm, the aliquots analyzed (0.1 g and more) could have contained up to several thousand microdiamonds. Thus, the data presented in Table 5.2 reflect average δI3 C values. Such differences in δ13 C values found in graphite and diamond from the same rock type cast a shade on the graphite-diamond genetic hypothesis, as if the latter is true, the isotopic compositions of these two substances should be similar (Galimov et al. 1973). Progress in SIMS analysis of the carbon isotopic composition enabled individual determinations on diamond grains with a diameter as little as 10 μm (Imamura et al. 2004). Having studied grains of R- and S-types of diamonds (in terms of Ishida et al. 2003) from the gamet-pyroxene-carbonate rocks, the researchers obtained δ13 C data, which differ with morphologic types. The R-type diamonds displayed smaller variations in δ13 C (−8 to −16%o , avg. −11.8%o ) relative to the S-type diamonds. Nuclei of the S-type diamonds display δI3 C ranges from −9 to − 20%o (avg. −15.3%o ), whereas the rims of the grains display values from −15 to − 27%o (−22%o ). This fact favored a two-stage hypothesis of S-type diamond growth (Ishida et al. 2003). In general, high concentrations of paramagnetic nitrogen centers occur as single atoms (6.9 × 1018 cm−3 ). are typical of microdiamonds (ca. two decimal orders higher than measured in kimberlite diamonds, (see Finnie et al. 1994; Cartigny et al. 2001). All metamorphic diamonds contain nitrogen as a major structural impurity in C- and A- defects, so these diamonds fall into the Ib-IaA type (Finnie et al. 1994; Taylor et al. 1996; De Corte et al. 1998, 2000; Nadolinny et al. 2006). The nitrogen content in diamonds, as estimated by FTIR spectroscopy, ranges from 500 to 11,150 ppm. Excessive nitrogen that remains undetermined by FTIR spectroscopy most likely

5.2 Micron-Sized Diamond Formation During Metamorphic Processes

83

Fig. 5.13 Carbon and nitrogen isotopic compositions of Kumdykol diamonds compared with published values for diamonds of known provenance (Lavrova et al. 1999). 1—Kumdykol diamonds; 2—diamonds from kimberlites (reproduced from Fig. 5 of Pechnikov and Kaminsky 2008)

occurs in diamonds as fluid inclusion filling. Such discrepancies between the FTIR and bulk combustion data are unknown in the mantle-derived atoms and can be indicative of the metamorphic origin of diamonds (Cartigny et al. 2001). Taking into account the different fractionation curves between graphite and carbonate as a function of temperature, Cartigny et al. (2001) concluded that Kokchetav microdiamonds crystallized at temperatures between 540 and 640 °C.

5.2.5 Processes of Kokchetav Microdiamond Formation Currently, the origin of metamorphic diamonds is hotly debated. At least three different genetic concepts for the origin of the Kokchetav Massif diamonds exist:

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

Table 5.2 Carbon isotopic composition of coexisting diamond, graphite, and carbonate in diamondiferous rocks of the Kumdykol deposit (δ13 C%o ) Gneiss

Silicate-carbonate rock

Garnet-pyroxene rock

Quartzose rock

Diamond −16.60 −18.70 −15.40 −17.01 −17.69 −16.5 −17.10 − 17.50 −16.60 −15.40 −16.80 −17.54 −16.70 δ13 Cav = −16.88

−12.4 −11.9 −10.6 − −10.71 −10.65 −10.34 −11.60 −17.30 12.59 −10.50 −10.32 −9.50 −12.59 −10.22 δ13 Cav = −10.5 −10.21 −11.10 −10.10 δ13 Cav = −11.12

Graphite −23.10 −24.10 −23.5 − -15.30 -16.70 -16.0 19.80 δ13 Cav = −16.0 −20.90 −25.20 −23.70 −25.2 −23.10 −18.40 δ13 Cav = −22.7 Carbonate −6.6 −4.3

−5.9

mantle (i), crustal-mantle (ii), and crustal fluid-metasomatic (iii). The mantle hypothesis (i) states that existing diamondiferous metamorphic rocks originate from mafic– ultramafic melts, originally intruded into the crust and later metamorphosed; eclogites are considered xenoliths of these mantle rocks (Marakushev et al. 1998). Diamonds found in eclogite, as the proponents of this hypothesis believe, originate from contamination during the sample treatment at Yakutian mills, but once reported, this statement had never been confirmed. The crustal-mantle hypothesis (ii) proposes that primary crustal rocks, when dragged deeply into a subduction zone, undergo UHP metamorphism responsible for diamond formation (Sobolev and Shatsky 1987; Sobolev et al. 1991). Subsequently, diamondiferous rocks were presumed to occur under amphibolite-facies conditions. This hypothesis relies upon the fact that the HP minerals (e.g., diamond, coesite, K-clinopyroxene, aluminous titanite, and high-Si phengite) observed only as microinclusions in garnet and zircon are relict minerals. Taking into consideration that diamondiferous rocks are mostly gneisses and carbonate rocks (not eclogite), these authors concluded that UHP metamorphism was a key genetic control. Mantle ultrapotassic fluids played a great role in diamond formation according to experimental data under high (75 kbar) pressure (Pal’yanov et al. 2007). Both the crustal-mantle and mantle hypotheses are connected with UHP conditions. The term UHP metamorphism was born as a result of coesite occurrences in metamorphic rocks. Chopin (1984) discovered metamorphic coesite in pyrope quartzites of the Dora-Maira Massif, Italian Western Alps. In the same year, Smith (1984) described coesite from eclogites of the Western Gneiss Region in Norway.

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85

Fig. 5.14 P–T plot relative to the diamond–graphite transition curve of Bundy et al. (1960) for Alpe Arami peridotites (◯ ) and eclogites (●). P–T parameters were calculated by the sensors of Brey and Kohler (1990), Harley (1984), Simakov (2008), Ellis and Green (1979). Analyzes for the calculations were taken from Ernst (1977, 1978)

The occurrences indicate that supracrustal rocks formed or embedded at shallow levels of the continental or oceanic lithosphere have experienced recrystallization under PT conditions above the low-P stability limit of coesite. Reinecke (1991, 1998) described coesite inclusions from eclogites and metasediments of Lago di Cignana, Italien Western Alps. Meanwhile, the first notification of the high P–T parameters corresponding to the upper mantle generation of the Alpe Arami metamorphic layers associated with pyropic garnet-bearing Iherzolite was done by Ernst (1977, 1978). The estimated P–T on the basis of Grt-Opx thermobarometers for these garnet-bearing Iherzolites correspond to 907–1088 °C and 36–47 kbars (Fig. 5.14). The calculated P–T for the Dora-Maira and Alp Arami eclogites on the basis of Grt-Cpx thermobarometers correspond to 840–880 °C and 24–55 kbars (Fig. 5.14). The contents of UHP minerals in the Kokchetav metamorphic rocks are orders of magnitude lower than their diamond contents. The P–T estimation for Kokchetav GrtCpx diamond-bearing rocks corresponds to 750–1030 °C and 17–47 kbar (Fig. 5.15). It coincides with Schertl and Sobolev’s (2013) previous P–T estimates corresponding to P > 43 kbar in the range of 900–1000 °C. Meanwhile, it is significantly lower than diamond-forming ultrapotassic fluids in accordance with the experimental data of Pal’yanov et al. (2007). From the provided estimations, it follows that the calculated P–T parameters of the Western Alp eclogites and peridotites correspond more to the field of thermodynamic diamond stability than the Grt-Cpx diamond-bearing rocks of the Kokchetav deposit (Figs. 5.14 and 5.15). Then, according to the models of the UHP metamorphic diamond formation, the microdiamond deposit should also be present in the Western Alps. Meanwhile, there is only one report of microdiamond presence detected by Raman spectroscopy in fluid inclusions of Lago di Cignana garnetites (Frezzotti et al. 2014). The proportion of the diamondiferous Kokchetav rocks connected with UHPs (garnet-pyroxene and silicate-carbonates) in the total volume of diamondiferous rocks does not exceed 10–15%. Even if we exclude these

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

Fig. 5.15 P–T plot relative to the diamond–graphite transition curve of Bundy et al. (1960) for GrtCpx diamond-bearing metamorphic parageneses from Kokchetav. P–T parameters were calculated by the sensors of Simakov (2008), Ellis and Green (1979). Analyzes for the calculations were taken from Sobolev and Shatsky (1990), Shatsky et al. (1995), Lavrova et al. (1999)

rocks from the metamorphic sequence, this will have practically no effect on the diamond-bearing rock volume. Gneiss carries the greater part of diamonds. It is possible to conclude that the Kokchetav diamond deposit is not connected with UHP conditions. The crustal fluid-metasomatic hypothesis (iii) is based on the facts that the Kumdykol deposit and Barchi occurrence are confined to NE-trending fracture zones. The spatial distribution of diamonds has no lithological control; diamondiferous rocks vary in their composition from silicate to essentially carbonate varieties with a wide spectrum of mineral assemblages in these rocks. Garnet-biotite gneiss makes up approximately 80% of the diamondiferous zone. Diamonds occur in metamorphic rocks, which undergo dislocations and metasomatism controlled by the fracture zones. Diamondiferous linear zones are largely confined to fault-related metasomatite with sulfides and graphite. Diamonds in host rocks occur as inclusions in all major minerals and intergranular spaces. Diamond and graphite display a close paragenetic relationship (frequently forming intergrowths), and their spatial distribution has common structural controls. Crystallization of diamonds occurred synchronously with the deposition of graphite. Metasomatites connected with the fluid ascended along permeable pathways. They caused secondary alteration with chlorite-sericitization, chlorite-sericitecarbonatization, graphitization, carbonatization (as a multitude of thin veinlet witnesses), and sulfidization. The temperatures of diamond formation by Cartigny et al. (2001) correspond to these processes. As shown in Chaps. 3 and 4, organic formation in these processes stimulates diamond nucleus formation. The morphology of the mineralized body (i.e., the concentration model of diamondiferous bodies within the explored part of the tract) displays a direct correlation with the fluid pathways. These processes and temperatures coincide with the processes of micron-sized diamond formation from CO2 –H2 O–CH4 fluids during the secondary alteration of

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87

southern Mexico ophiolitic chromitites (Farré-de-Pablo et al. 2018) described in Chap. 4. In accordance with the temperature estimations of Cartigny et al. (2001), the temperatures of the Kokchetav diamond formation correspond to greenschist facies (540–640 °C). As shown in Chap. 4, at these temperatures, micron-sized diamonds can be formed from fluids during serpentinization. The isotopic composition of Kokchetav graphite carbon is depleted in heavy isotopes, while diamond is enriched (Lavrova et al. 1999). This may be due to the simultaneous formation of graphite and diamond at low P–T parameters from fluid (Fedoseev et al. 1971). Water-carbon dioxide with a lower hydrocarbon composition of Kokchetav diamond fluid inclusions (De Corte et al. 1998) corresponds to diamond nuclei formation at greenschist P–T parameters (see Chap. 2). They could be formed by the following reaction: CH4 + CO2 → 2C + 2H2 O

(5.2)

It is known that the nitrogen content in the upper mantle is low; meanwhile, in crustal sediments, it can reach high levels in the greenschist facies of metamorphism (Busigny and Bebout 2013). Low δ13 C volumes of Kokchetav diamonds correspond to these sources. Nitrogen entrance into the diamond structure at low P–T parameters leads to their stabilization (see Chaps. 2 and 3). It could be entranced from sediment ammonia by the following reaction: +3 NH+ + 4H− 4 →N

(5.3)

Further diamond nuclei growth up to micron sizes could be connected to produce reactive intermediates CH3 , CH2 , CH3 OH, CH2 OH, HO2 , H2 O2 , H, OH formed at organic formation and from fluids at supercritical water stability (P–T exceeding 375 °C and 22 MPa, respectively) and stimulating diamond growth at metastable conditions (Antal et al. 1987). The aforementioned data show that Kokchetav microdiamonds can be formed during alteration hydrothermal-metasomatic processes at P–T corresponding to the greenschist facies. Their high contents of nitrogen and metasediment values of δ15 N confirm this conclusion (Cartigny 2005).

5.2.6 Conclusions As a result, it is possible to conclude that Kokchetav diamonds were formed during crustal metamorphic hydrothermal-metasomatic processes from hydrocarbonnitrogen-bearing fluids at P–T parameters corresponding to greenschist facies. They are products of an open disequilibrium system at a relatively low-temperature level and pressure. These conditions are responsible for the carbon isotope disequilibrium of the crystallization environment and low aggregation grade of the diamond

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5 Nanocarbon and Microdiamond Formation in the Lithogenesis …

structure, micrometer-size dimensions of diamonds, and prevalence of imperfect crystalline patterns. It is questionable a generalized UHP origin for Kokchetav diamonds.

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Fedoseev DV, Galimov EM, Varnin VP et al (1971) Carbon isotope fractionation in the process of physic-chemical synthesis of diamond. Dokl Akad Nauk SSSR 201:1149–1151 Finnie KS, Fisher D, Griffin WL et al (1994) Nitrogen aggregation in metamorphic diamond from Kazakhstan. Geoch Cosm Acta 58:5173–5177 Frezzotti M-L, Huizenga J-M, Compagnoni R et al (2014) Diamond formation by carbon saturation in C–O–H fluids during cold subduction of oceanic lithosphere. Geoch Cosm Acta 143:68–86 Galimov EM, Prohorov VS, Fedoseev DV et al (1973) Heterogenic isotopic effects of carbon at synthesis of diamond and graphite from gas. Geokhimiya 416–425 Gebbiea MA, Ishiwatab H, McQuadea PJ et al (2018) Experimental measurement of the diamond nucleation landscape reveals classical and nonclassical features. PNAS 115:8284–8289 Harley SL (1984) An experimental study of the partitioning of the Fe and Mg between garnet and orthopyroxene. Contrib Mineral Petrol 86(2):359–373 Hermann J, Rubatto D, Korsakov A et al (2006) The age of metamorphism of diamondiferous rocks determined with SHRIMP dating of zircon. Russ Geol Geoph 47:513–520 Hwang S-L, Shen P, Chu H-T et al (2005) Crust-derived potassic fluid in metamorphic microdiamond. Earth Plan Sci Lett 231:295–306 Hwang S-L, Chu HT, Yui T-F et al (2006) Nanometer-size P/K-rich silica glass (former melt) inclusions in microdiamond from the gneisses of Kokchetav and Erzgebirge massifs: diversified characteristics of the formation media of metamoiphic microdiamond in UHP rocks due to host-rock buffering. Earth Plan Sci Lett 243:94–106 Imamura K, Yoshioka N, Ogasawara Y (2002) Morphology and distribution of microdiamonds in dolomite marble from Kumdy-Kol. In: Parkinson CD, Katayama I, Liou JG, Maruyama S (eds) The diamond-bearing Kokchetav Massif, Kazakhstan. Universal Academy Press Inc., Tokyo, pp 93–102 Imamura K, Ogasawara Y, Yurimoto H et al (2004) Carbon isotope compositions of microdiamond in UHP marble. In: Abstract of 32nd international geological congress, vol 1, pp 720–721 Ishida H, Ogasawara Y, Ohsumi K et al (2003) Two stage growth of microdiamond in UHP dolomite marble from Kokchetav Massif, Kazakhstan. J Metam Geol 21:515–522 Ishimaru K, Vystavel T, Bronsveld P et al (2001) Diamond and pore structure observed in wood charcoal. J Wood Scie 47(5):414–416 Jagoutz E, Shatsky VS, Sobolev NV (1990) Sr–Nd–Pb isotopic study of ultrahigh PT rocks from Kokchetav Massif. EOS Trans Am Geoph Un 71:1707 Jehlicka J, Rouzaud JN (1993) Transmission electron microscopy of carbonaceous matter in Precambrian shungite from Karelia. In: Parnell J, Kucha H, Landais P (eds) Bitumens in ore deposits. Springer, Berlin, pp 53–60 Jehlicka J, Ozawa M, Slanina Z et al (2000) Fullerenes in solid bitumens from from pillow lavas of Precambrian age (Mitov, Bohemian massif). Full Sci Teh 8:449–452 Kashkarov IF, Polkanov YA (1972) Some specific features of diamonds from titaniferous placers of Northern Kazakhstan. Novye Dannye o Mineralakh SSSR 21:183–185 Khachatryan GK, Baryshev AN (2022) Nitrogen and hydrogen in diamonds: consequences of minerageny. TsNIGRI, Moscow Koniakhin SV, Utesov OI, Terterov IN et al (2018) Raman spectra of crystalline nanoparticles: replacement for the phonon confinement model. J Phys Chem C 122:19219–19229 Korsakov AV, Shatsky VS (2004) Origin of graphite-coated diamonds from UHP metamorphic rocks. Dokl Earth Sci 399:1160–1163 Korsakov AV, Shatsky VS, Sobolev NV (1998) The first finding of coesite in eclogites of the Kokchetav Massif. Dokl Akad Nauk 360:77–81 Korsakov AV, Theunissen K, Smirnova LV (2004) Intergranular diamonds derived from melting of crustal rocks at ultrahigh-pressure metamorphic conditions. Terra Nova 16:146–149 Kouchi A, Nakano H, Kimura1 Y et al (2005) Novel routes for diamond formation in interstellar ices and meteoritic parent bodies. Astrophys J 626:L129–L132 Kovalevski VV, Buseck PR, Cowley JM (2001) Comparison of carbon in shungite rocks to other natural carbons: an X-ray and TEM study. Carbon 39:243–256

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Pechnikov VA, Bobrov VA, Podkuiko YuA (1993) Isotopic composition of carbon of diamond and accompanying graphite from metamorphic rocks of the Northern Kazakhstan. Geokhimiya I:150–153 Potgieter-Vermaak S, Maledi N, Wagner N et al (2011) Raman spectroscopy for the analysis of coal: a review. J Raman Spectrosc 42:123–129 Reinecke T (1991) Very high pressure metamorphism and uplift of coesite-bearing metasediments from the Zermatt-Saas zone, Western Alps. Eur J Mineral 3:7–17 Reinecke T (1998) Prograde high- to ultrahigh-pressure metamorphism and exhumation of oceanic sediments at Lago di Cignana, Zermatt-Saas zone, Weatern Alps. Lithos 42:147–190 Schertl H-P, Sobolev NV (2013) The Kokchetav Massif, Kazakhstan: “Type locality” of diamondbearing UHP metamorphic rocks. J Asian Earth Scie 63:5–38 Shatsky VS, Sobolev NV (1993) Some genetic aspects far diamond in metamorphic rocks. Dokl Akad Nauk 331:217–219 Shatsky VS, Sobolev NV, Vavilov MA (1995) Diamond-bearing metamorphic rocks of the Kokchetav Massif (Northern Kazakhstan). In: Coleman RW, Wang H (eds) Ultrahigh pressure metamorphism. Cambridge University Press, pp 427–455 Shatsky VS, Rylov GM, Efimova ES et al (1998a) Morphology and real structure of microdiamonds from the metamorphic rocks of the Kokchetav massif, kimberlites and alluvial placers. Russ Geol Geophys 39:942–955 Shatsky VS, Theunissen K, Dobretsov NL et al (1998b) New evidence of ultrahigh-pressure metamorphism in the mica schists of the Kulet area of the Kokchetav massif (Northern Kazakhstan). Russ Geol Geophys 39:1041–1046 Simakov SK (2008) Garnet-clinopyroxene and clinopyroxene geothermobarometry of deep mantle and crust eclogites and peridotites. Lithos 106(1/3):125–136 Simakov SK (2015) Different sizes of diamond formation in natural processes. Dokl Earth Sci 461:419–421 Simakov SK, Kouchi A, Mel’nik NN et al (2015) Nanodiamond finding in the Hyblean shallow mantle xenoliths. Sci Rep 5:10765 Smith DC (1984) Coesite in clinopyroxene in the Caledonides and its implications for geodynamics. Nature 310:641–644 Smith PPK, Buseck PR (1981) Graphitic carbon in the Allende meteorite: a microstructural study. Science 212:322–324 Sobolev NV (2006) Coesite as an indicator of ultrahigh pressures in continental lithosphere. Russ Geol Geoph 47:95–104 Sobolev NV, Shatsky VS (1987) Carbon mineral inclusions in garnets from metamorphic rocks. Russ Geol Geophys 7:77–80 Sobolev NV, Shatsky VS (1990) Diamond inclusions in garnets from metamorphic rocks: a new environment for diamond formation. Nature 343:742–746 Sobolev NV, Shatsky VS, Vavilov MA et al (1991) Coesite inclusions in zircon from diamondbearing gneiss of the Kokchetav massif: first find of coesite in metamorphic rocks in the USSR. Dokl Acad. Nauk SSSR 321:184–188 Sobolev NV, Shatsky VS, Vavilov MA et al (1994) Zircon from high-pressure metamorphic rocks from folded areas as a unique container of diamond, coesite and coexisting mineral inclusions. Dokl Akad Nauk 334:488–492 Sobolev NV, Schertl H-P, Burchard M et al (2001) An unusual pyrope-grossular garnet and its paragenesis from diamondiferous carbonate-silicale rocks of the Kokchetav Massif, Kazakhstan. Dokl Earth Sci 380:791–794 Sobolev NV, Schertl H-P, Neuser RG et al (2007) Relict unusually low iron pyrope-grossular garnets in UHPM calc-silicate rocks of the Kokchetav massif, Kazakhstan. Int Geol Rev 49:717–731 Sokolov VA (1987) The Jatulian superhorizon. In: Sokolov VA (ed) The geology of Karelia. Nauka, Leningrad, pp 51–59 Sokolov VA, Kalinin YuK (1975) Karelian shungites and their utilization. Karelia, Petrozavodsk

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Stockhert B, Duyster J, Trepmann C et al (2001) Microdiamond daughter crystals precipitated from supercritical COH + silicate fluids included in garnet, Erzgebirge, Germany. Geology 29:391–394 Taylor WR, Canil D, Milledge HJ (1996) Kinetics of lb to laA nitrogen aggregation in diamond. Geoch Cosm Ada 60:4725–4733 Tian H, Schryvers D, Claeys P (2011) Nanodiamonds do not provide unique evidence for a Younger Dryas impact. PNAS 108:40–44 Timofeev VM (1924) On genesis of shungite from the Onega region. Proc Leningrad Soc Naturalists 4:40–50 Tissot BP, Welte DH (1978) Petroleum formation and occurrence: a new approach to oil and gas exploration. Berlin–Heidelberg–New York Ugarte D (1992) Curling and closure of graphitic networks under electron-beam irradiation. Nature 359:670–671 Van Roermund HLM, Carswell DA, Drury MR et al (2002) Microdiamond in a megacrystic garnet websterite pod from Bardane on the Island of Fjortoft, western Norway: evidence for diamond formation in mantle rocks during deep continental subduction. Geology 30:959–962 Vikentyev IV, Kailachakov PE (2020) The unique deposit of rhenium in the coal-bearing carboniferous sands of the Russian plate. Communication 1. Geological Structure. Lithol Mineral Resourc 55(3):177–191 Vinokurov SF, Novikov UN, Usatov AV (1997) Fullerenes in the geochemistry of endigenic processes. Geohimiya 9:937–944 Volkova IB, Bogdanova MV (1986) The shungites of Karelia. Int Geol Rev 6:1343–1351 Vyalov VI (1996) The petrology of carbon graphites. Lithol Mineral Res 31(1):70–78 Vyalov VI (1998) The petrology of anthracites. Lithol Mineral Res 33(3):271–284 Vyalov VI, Nastavkin AV (2019) Concentration levels of industrially valuable trace elements in coals. Solid Fuel Chem 53(5):314–318 Vyalov VI, Golitsyn MV, Golitsyn AM (1998) Anthracites of Rassia and the world. Nedra, Moscow Vyalov VI, Nastavkin AV, Shishov EP (2021) Distribution of industrially valuable trace elements associated with germanium in the coals of the Pavlovsk deposit (Spetsugli Section). Solid Fuel Chem 55(1):14–25 Wang ZX, Li XP, Wang WM et al (1998) Fullerenes in the fossil of dinosaur egg. Full Sci Teh 6:715–720 Wei Z, Moldowan JM, Jarvie DM et al (2006) The fate of diamondoids in coals and sedimentary rocks. Geology 34:1013–1016 Wei Z, Moldowan JM, Shuichang Z et al (2007) Diamondoid hydrocarbons as a molecular proxy for thermal maturity and oil cracking: Geochemical models from hydrous pyrolysis. Org Geochem 38:227–249 Xu ST, Okay AI, Ji S et al (1992) Diamond from the Dabie Shan metamorphic rocks and its implication for tectonic setting. Science 256:80–92 Yang J, Xu Z, Dobrzhinetskaya LF et al (2003) Discovery of metamorphic diamonds in central China: an indication of a >4000-km-long zone of deep subduction resulting from multiple continental collisions. Terra Nova 15:370–379 Yoshioka N, Muko A, Ogasawara Y (2001) Extremely high diamond concentration in dolomite marble. In: Extenmded abstract of UHPM workshop 2001 at Waseda University, pp 51–55

Chapter 6

Diamonds in Kimberlites and Their Xenoliths: A Reappraisal

6.1 Introduction Diamonds of any sizes typically occur in kimberlites and their xenoliths. It is an established notion that the formation of macrodiamonds generally take place under pressure (P) and temperature (T) conditions compatible with a section of the subcontinental mantle from approximately 100 km to 750 km and hence in the field of diamond thermodynamic stability (Meyer and Boyd 1972; Nimis 2022; Smith et al. 2016; Stachel and Luth 2015; Sobolev 1977). Subcratonic lithospheric diamonds represent the primary source of over 99% (by mass) of world diamond production (Stachel and Harris, 2008). Sublithospheric diamonds formed at great depths of up to 750 km are very rare. The carbon isotopic composition of macrodiamonds varies from + 3 ‰ to −40 ‰ of δ13 C, the main part of values lying in the range from 0 to −25‰, with a prominent mode at −5‰ (Kirkley et al. 1991; Cartigny 2005; Stachel et al. 2022). Their nitrogen contents vary from ~ 0 to 3830 ppm. The majority of values lie in the range from 0 to 1500 ppm, with a median value of 160 ppm (Stachel and Harris 2009; Stachel et al. 2022). The nitrogen isotopic composition, indicated as δ15 N, varies from −18‰ to + 15‰ with a prominent mode at −2 ‰ (Cartigny 2005; Stachel et al. 2022). There is no fundamental difference between micro- and macrodiamonds, apart from size (McCandless et al. 1991; Sobolev et al. 2004; Reutsky and Zedgenizov 2007). Meanwhile, extralarge diamonds (e.g., type IIa) have very low nitrogen content (Moore 2009; Gurney and Helmstaedt 2012), whereas micron-sized diamonds are enriched in nitrogen (Pattison and Levinson 1995).

© The Author(s), under exclusive license to Springer Nature Switzerland AG 2023 S. Simakov et al., Nano and Micro Diamond Formation in Nature, SpringerBriefs in Earth Sciences, https://doi.org/10.1007/978-3-031-43278-1_6

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6.2 Diamond Inclusions Mantle macrodiamonds usually contain different mineral solid, fluid, brien, and melt inclusions (Arai 1986; Izraeli et al. 2001; Stachel and Harris 2009). Currently, the notion that diamond formation coincides with the formation of their inclusions is the most popular. On the other hand, there is a school of thought (Nestola et al. 2017) suggesting that inclusions in diamonds cannot be reliable proxies for their origins, as a given mineral type growing in a medium that hosts a heterogeneous assemblage of microxenoliths can trap one or the other mineral species following the laws of chance. Solid inclusions in diamonds allowed us to recognize three types of rocks as likely substrates for diamond formation in the lithospheric mantle: peridotite, eclogite, and websterite. Peridotite is thought to be the most important source rock of diamonds, as approximately 65% of diamond inclusions are consistent with a peridotite origin, followed by eclogite (33%), and websterite (only 2%) (Stachel1 and Harris 2008). The known P–T estimates for the peridotite and eclogite inclusions in diamonds are in the ranges of 650–1600 °C and 30–100 kbar, which correspond to depths from 100 to 300 km (Fig. 6.1). Rare sublithospheric inclusions can form at depths of 360–750 km (Smith et al. 2016; Nestolla et al. 2023). From the comparative evaluation of the P–T conditions of the diamond formation in the lithospheric mantle and the experimental results on their growth, it follows that mantle diamonds could be formed from metal–carbon and silicate-carbonate melts and water-carbon dioxide fluids (see Fig. 2.1 in Chap. 2). Studies on the physical properties of diamond matter have shown that there are three main zones in the crystals: central, periphery, and intermediate (Beskrovanov 1986). In this respect, the mineral assemblage “sulfide + native iron + wustite + graphite” in the inner core of macrodiamonds may provide evidence of redox conditions in their nucleation stage (Bulanova 1995). It is worth noting that such a rare type of “central” inclusion displays analogies with the “iron meteorite” mineral composition (Arai 1986; Bulanova and Zayakina 1991; Bulanova et al. 1979). Looking into more detail, the aforementioned inclusions consist of native iron, blebs of solidified iron-nickel-carbon-sulfur melt, taenite, native chromium, periclase-wustite, and moissanite (Bulanova et al. 1979; Gorshkov et al. 1995; Harris and Gurney 1979; Otter and Gurney 1989; Smith et al. 2016; Sobolev et al. 1981; Wilding et al. 1991). Such an occurrence indicates that the diamonds grew in a reduced environment from a redox-sensitive metallic liquid phase (Crookers 1909; Simakov 1987; Smith et al. 2016). Sulfides are widely represented in diamonds and mantle xenoliths from kimberlites, thereby indicating that they play a special role in deep mineral formation (Sobolev et al. 1981; Yefimova et al. 1983; Bulanova 1995). On the other hand, sulfur is an inhibitor in the synthesis of diamond from Fe–Ni–C melts (Wentorf and Bowenkerk 1961; Tsuno and Dasgupta 2015; Palyanov et al. 2020). According to Zhimulev et al. (2012), diamond can crystallize in the mantle with low sulfur concentrations and is accompanied by sulfide and metallic minerals. Rohrbach et al.

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Fig. 6.1 Field of known P–T parameters obtained for diamond peridotitic and eclogitic inclusions on the basis of Grt-Opx, Grt-Cpx, and Cpx thermobarometers (I) (data taken from Haggerty 1986; Harris 1992; Nimis 2022; Simakov 2008, 2018a; Shirey et al. 2013; Stachel and Luth 2015) and “diamond window” (gray color) relative to diamond–graphite boundary (solid line) (Bundy et al. 1961) and continental geotherms (dotted lines with surface heat flow values in mW/m2 ; Hasterok and Chapman, 2011). Legend: LS—lherzolite-water solidus (Wyllie and Ryabchikov 2000), WSwebsteritic solidus (Mysen and Boettcher 1976)

(2011, 2014) postulated that most of the deep-reducing mantle contain a small Fe– Ni–C melt fraction and that these melts should be ubiquitous in the mantle. Adopting bulk carbon contents of 50–500 ppm in the mantle would result in the phase association (Fe, Ni)3 C + metal + diamond. In accordance with these studies, fertile mantle composition may reach Fe metal saturation and Fe–Ni–C melt formation at mantle depths greater than 220 km. The calculated P–T- f O2 parameters of pyrope garnet inclusions from the central zones with native iron of Yakutian diamond indicate their formation at mantle depths of 200–270 and redox conditions of IW buffer and even lower (Fig. 6.2). This confirms the suggestions of Rohrbach et al. (2011, 2014). Only a few μm-sized Fe–Ni droplets were observed in garnets from polycrystalline diamond aggregates (Jacob et al. 2004) hosted in kimberlitic eclogite and peridotite xenoliths (Stachel and Harris 2009), which corresponds to the suggestion of Rohrbach et al. (2014). These μm-sized Fe–Ni droplets are compatible with the formation of μm-sized diamonds only. The “iron-meteorite-like” inclusions in diamonds are sometimes associated with majoritic garnet, perovskite, and magnesiowustite, which are thought to represent superdeep lower mantle paragenesis (Zedgenizov et al. 2001; Hutchinson et al. 2001; Kaminsky et al. 2000; Moore and Gurney 1989; Smith et al. 2016; Wilding et al. 1991). The presence of a pure forsterite end-member olivine inclusion implies that localized environments of the transition zone–lower mantle region can be very reducing and metal-rich (Nestolla et al. 2023). Smith et al. (2016) detected the assemblage constituted by majoritic garnet, calcium silicate perovskite, Fe-Ni-C-S melt inclusions, and reduced volatiles in the inclusions in type II gem diamonds, whose sizes span from 1 to 30 carats. On the other hand, the rare presence of carbonate inclusions in the central part of some diamonds indicates an oxidizing environment (Bulanova and Pavlova 1987; Logvinova et al. 2019). Their formation is possible

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Fig. 6.2 P–T- f O2 plots for the pyrope garnet inclusions from central paragenesis of Yakutian diamonds (·) estimated using the garnet-(olivine) and garnet-(orthopyroxene) thermobarometers of Creighton (2009) and Ryan et al. (1996) and garnet-(olivine-orthopyroxene) oxygen barometer of Simakov (2013) relative to diamond–graphite boundary (solid line) (Bundy et al. 1961) and eutectic of Fe–C system (Fe-CL dashed line) in accordance with Nakajima et al. (2009) and Hirayama et al. (1993). The analyzes for the calculations were taken from Sobolev et al. (1981). Legend: QFM, WM, and IW in accordance with Frost (1991)

from carbonate melts at temperatures of 1150–1500 °C near the boundary of the lithosphere and asthenosphere (see Fig. 2.1 in Chap. 2). Epigenetic inclusions are found in many diamonds as fracture fills and replaces primary phases. They include serpentine, calcite, dolomite, amphibole, acmite, graphite, hematite, magnetite, kaolinite, perovskite, Mn-ilmenite, sulfides, xenotime, goethite, apatite, mica (phlogopite, biotite), microcline, sellaite, spinel and talc (e.g., Meyer and McCallum 1986; Davies et al. 2004; Taylor and Anand 2004; Banas et al. 2007; Klein-Bendavid et al. 2007; De Stefano et al. 2009; Korolev et al. 2018). In particular, serpentine, as an epigenetic mineral, has been described in diamonds many times. Interestingly, Tal’nikova (1995) described the mineral association of serpentine-chromite-calcite-sulfides in Yakutian diamond from the Udachnaya pipe despite the lack of cracks in the diamond surface. The author suggested that it was probably protogenic and entrapped from a kimberlite melt during final growth in the Earth’s crust. Klein-Bendavid et al. (2007) identified Mg–Fe silicate with a structure similar to that of serpentine in fibrous Yakutian diamond. In accordance with the observations of Moore (2014) and Cartigny et al. (2014), the formation of fibrous diamonds was linked to the kimberlite magmatic event just prior to kimberlite eruption. Organic matter is also present in diamonds. Khachatryan (2017) studied impurities of organic matter locally bearing nitrogen by infrared spectroscopy in diamond crystals from Yubileynaya pipe (Yakutiya) and Juina (Brazil). Crystals from both sites contain impurities of graphite in the marginal zones. For the Yakutiyan diamond, there is a tendency to decrease the aggregated nitrogen form (%Ng ) from the center to the marginal zone regardless of the total nitrogen concentration. This feature corresponds to the decrease in crystallization temperature in accordance with the model by Taylor and Milledge (1995). Infrared spectroscopy also shows the presence of organic matter in the form of methylene (CH2 ) and methyl (CH3 ) groups of

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saturated hydrocarbons, always in the marginal zone. The diamond from the Juina region (Brazil) contains a high concentration of completely aggregated nitrogen in the central part (nearly 1000 ppm) and relicts of organic matter in the marginal part. The organic matter occurs both in nitrogen-rich and nitrogen-poor zones of the crystal. These data, combined with the relatively light isotope carbon composition of the studied diamonds (−9.6‰ and −22.0‰ of δ13 C correspondingly), show mainly the increase in organic matter in the late stage of Juina diamond formation and therefore at decreasing temperatures. Interestingly, Melton and Giardiny (1974; 1975) detected low percentages of methanol and ethanol in the fluid inclusions of some African diamonds, and Sobolev et al. (2018; 2019a, b) reported alcohol, ketone, carboxylic, and carbon acid occurrences in the fluid inclusions of Siberian and Ural diamonds. On the basis of mineral inclusions in Yakutian diamonds, Garanin and Kudryavtseva (1990) concluded that they formed in four different environments: (1) nucleation and growth under extremely reducing and high-pressure conditions; (2) P–T-X conditions related to eclogite formation; (3) P–T-X conditions related to the garnet-facies peridotite stability field; and (4) P–T-X parameters related to Ti-bearing chromite, serpentine, calcite, apatite, and phlogopite stability under relatively oxidized conditions. Detailed investigation of a great number of diamonds confirms the aforementioned scenario. The majority of the Orapa cluster diamonds (80%) show complex growth histories in their CL images, FTIR data, and C-N isotopes, confirming a multistage evolution (Gress et al. 2021). In this respect, it is worth mentioning that the absolute age of diamond formations, as deduced from solid inclusions, corresponds to that of the old mantle, as well as to the age of kimberlite formations (Gress et al. 2021; Burgess et al. 2002; Richardson 1986; Richardson et al. 1984, 1993; Smith 1983; Navon 1999). Suites of eclogite-related silicate inclusions give Sm– Nd isochron ages from 1 to 2 Ga, whereas peridotitic silicate inclusion suites give isochron or model ages from 1.9 to 3.4 Ga (Richardson 1986; Shirey and Richardson 2011). Nimis et al. (2020) concluded that diamonds from the Cullinan kimberlite pipe may have very different ages and may have been formed by different processes under different thermal regimes, as they contain different inclusion types.

6.3 Diamond Formation from Fluids in the Upper Mantle There is plenty of evidence that the formation of diamonds is associated with fluids of probable mantle origin (Pokhilenko et al. 2015; Sobolev et al. 2019b; Stachel and Luth 2015) and could be connected with the subduction of oceanic crust with highly saline fluids (Weiss et al. 2015). The main fluid components in kimberlite diamonds are H2 O, H2, N2 , CO2 , CH4 and CO (Giardiny and Melton 1975; Melton and Giardiny 1974; 1975; Schrauder and Navon 1993; Izraeli et al. 2001; KleinBendavid et al. 2007; 2009; Tomlinson et al. 2006; Jablon and Navon 2016; Nimis et al. 2016). Chlorine-phosphate-fluorine medium was suggested for fibrous and cloddy diamonds on the basis of the findings in their mineral phases and brain

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halides, phospate, and fluorine-containing cuspidine (Klein-Bendavid et al. 2007; Izraeli et al. 2001; Wirth et al. 2009; Kaminsky and Schreber 2013). Chlorine and fluorine were also detected in the organic compounds of the fluid inclusions of the Yakutian and Ural diamonds (Sobolev et al. 2019a, b). They are noted in oxygenated hydrocarbons (n-pentyl methylphosphonofluoridate (C6 H14 FO2 P)), arenes (fluoro (C6 H5 F) and fluoromethyl (C7 H7 F) benzenes, fluorobenzyl alcohol (C9 H6 ClF3 O2 )), and nitrogenated organics (C7 H6 ClFN2 O and C6 H9 F2 NO). In addition, traces of hydrocarbons, sulfur, and sulfonated compounds (sulfur dioxide (SO2 ), carbon disulfide (CS2 ), dimethyl disulfide (C2 H6 S2 ), and thiophenes (C4 H4 S)) have been discovered as fluid inclusions in natural diamonds (Klein-Bendavid et al. 2007; Sobolev et al. 2019a, b). Although H2, N2 , CO, or CO2 can be the dominant fluid species, the average composition of diamond-related fluids has a hydrous nature (Giardini and Melton 1975; Izraeli et al. 2001; Bartoshinsky et al. 1987). A model for the C-H-O-N-S gaseous system has been generally employed to consider diamond formation processes in the deep lithospheric mantle and asthenosphere (Cartigny et al. 2014; Haggerty 1986; Simakov 1983, 1998, 2006; Sobolev et al. 2019b; Sokol et al. 2017). Diamond preservation in mantle processes depends upon the oxygen conditions. The limit of the carbon presence in the C-H–O system by oxygen fugacity (CCO buffer) corresponded to the following reactions (Jakobsson and Oskarsson 1994): CO2 → C + O2

(6.1)

CO → C + O.5O2

(6.2)

At higher f O2 values, free carbon transfers to CO2 and CO. The optimum conditions for carbon precipitation and diamond growth in the C-H–O system correspond to water-rich fluids (Simakov 1994, 1998, 2018b; Simakov and Stegnitskiy 2021; Stachel and Luth 2015). They could be formed by the following reaction: CH4 + O2 → C + 2H2 O

(6.3)

It corresponds to the average composition of fluids extracted from diamonds (Giardini and Melton 1975; Izraeli et al. 2001; Tomlinson et al. 2006; Nimis et al. 2016) and experimental data of diamond syntheses from fluid at the mantle P–T parameters (Sokol et al. 2009). At lower f O2 , the partial pressures of [CH4 + H2 ] exceed the partial pressures of H2 O, and the main part of the free carbon transfers to methane (Fig. 6.3). The oxygen fugacity of the CCO buffer increases with decreasing pressure and temperature relative to the QFM buffer (Luth 1993). The minimum atomic carbon content in the C–H–O system increases with temperature and decreases with increasing pressure at a fixed f O2 (Simakov 2018b). Grutter et al. (2006) have shown that P–T estimates of diamond inclusions and diamond-bearing xenoliths mainly lie

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Fig. 6.3 Diamond saturation field as a function of water content in a coexisting fluid at 1000 °C and 45 kbar in accordance with (Simakov 1998). CCO corresponds to equilibriums (6.1) and (6.2), and max H2 O and min C% correspond to equilibrium (6.3) in the text

in the field of “cold” palaeogeotherms with 36–41 mW/m2 values, while diamondfree xenoliths lie in the field of “hot” palaeogeotherms with values higher than 41 mW/m2 . Griffin and Ryan (1995) highlighted the “diamond window” for diamond-rich kimberlites as the range of temperatures between the intersection of typically cratonic geotherms (35–40 mW/m2 ) and the basement of the lithosphere (Fig. 6.1). Nearly, 80% of the P–T estimated for the diamond inclusions and diamond-bearing peridotite xenoliths in Simakov (1998) and Stachel and Luth (2015) corresponds to this window. By our calculation, fluid compositions with minimum contents of carbon correspond to “cold” geotherms, while fluids with higher carbon contents correspond to “hot” geotherms (Fig. 6.4). Water-rich fluids correspond to WM buffer at mantle P–T parameters and partially to the QFM buffer at low crustal P–T parameters (Fig. 6.5). Fig. 6.4 Relationship between the available minimum atomic carbon contents of the C–H–O fluid system and the P–T of conductive geotherms by Hasterok and Chapman (2011) in accordance with (Simakov 2018b)

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Fig. 6.5 T- f O2 plot of oxygen buffers of water-rich fluid and fluid optimum (max H2 O) for diamond growth in the 35 mW/m2 (a) and 41 mW/m2 (b) conductive geotherms by Hasterok and Chapman (2011). Legend: QFM, WM, and IW in accordance with Frost (1991), CCO–oxygen fugacity calculated for the upper limit of carbon stability in the C–H–O system; CO2 -H2 O and H2 O-CH4 boundaries of the water-rich fluid

As is well-known, the oxidation state of the mantle assemblage can be estimated by reactions between oxygen and iron-bearing minerals. For spinel-bearing peridotites, which usually originate at a depth range of 30–50 km, oxythermometer calculations use the spinel-orthopyroxene-olivine equilibrium reaction (Ballhaus et al. 1991; Wood 1991). For garnet-bearing peridotites and eclogites, which are thought to originate at depths mainly between 50 and 220 km, the oxythermmobarometry equilibriums use the garnet-orthopyroxene-olivine and garnet-clinopyroxene-silica equilibriums (Gudmundson and Wood 1995; Stagno et al. 2013, 2015; Simakov 1998, 2006). Estimated P–T- f O2 parameters and C-H–O fluid compositions for garnetorthopyroxene-olivine mineral assemblages indicate that at depths lower than 200 km, the formation of diamond inclusions and xenoliths also occurred under redox conditions corresponding to the range of the IW buffer and even lower (Simakov 1998; Stachel and Luth 2015) and therefore just beyond the optimal conditions for the formation of diamonds from fluid. At these parameters, a methane-rich fluid is stable in the C-H–O system. The oxygen fugacities of the diamond-bearing xenoliths formed at depths up to 200 km are mainly higher than those of the IW buffer (Simakov 1998; Stachel and Luth 2015) and correspond to the optimal conditions for the formation of diamonds from fluid. In accordance with the estimation of Simakov (1998), 80% of diamond-bearing peridotite xenoliths correspond to H2 Orich fluid containing a minimum percent of atomic carbon at temperatures lower than 1100° C. Calculated P–T- f O2 parameters and fluid compositions for more than 9000 pyrope-rich garnet xenocryst grains from South African, Angolan, and Yakutian pipes of different diamond grades also indicate that optimal conditions for diamond growth and preservation occur in presumed water-rich mantle fluids (Simakov and Stegnitskiy 2021). It is possible to conclude that the calculated compositions of

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the fluid could be an additional indicator of the diamond ability in the kimberlites and lamproites and can be used as a valid mineralogical-petrological method for prospecting kimberlitic diamond deposits. Stachel and Harris (2009) concluded that in the mantle, there are two models of peridotite diamond formation: reduction of carbonates and oxidation of methane. The first leads to an increase in δ13 C values from −5 o /oo to 0, and the second leads to a decrease in δ13 C from −5 o /oo to −10 o /oo . These models coincide with two types of central inclusions in diamonds (carbonate and iron metallic) and with variations in fluid compositions from oxidized to reduced diamond inclusions (Giardini and Melton 1975; Izraeli et al. 2001; Sobolev et al. 2019a). The oxidation of mantle fluids mainly decreases with depth from QFM to IW buffers (Ballhaus 1993; Simakov 1998, 2006; Stachel and Luth 2015). Therefore, the upper mantle diamonds could be formed by reaction (6.1) corresponding to higher δ13 C values of carbon or by reaction (6.3) corresponding to lower δ13 C values. Estimated P–T- f O2 and fluid compositions of eclogitic parageneses indicate that most of the diamond inclusions and diamond-bearing xenoliths also equilibrated at conditions below those of the CCO buffer, near WM buffer (Aulbach et al. 2022; Simakov 2006). Nitrogen and hydrocarbons play an important role in the natural processes of diamond formation. The free C-H–O dominating fluid phase cannot be stable at depths of 200–300 km because the presence of such a fluid would promote peridotite partial melting, thereby dissolving much of the fluid in the melt (Ballhaus and Frost 1994). In the asthenosphere, CO2 , H2 O, and some CH4 are dissolved in the melt (Taylor and Green 1987). Therefore, nitrogen could be an important component in the fluid phase of the asthenosphere. The primordial asthenospheric fluid from which the initial diamonds began to grow appears to consist of H2 , NH3 , N2 , and CH4, which mainly corresponds to the C-H-N gaseous system (Simakov 1998, 2006; Stachel and Luth 2015; Sokol et al. 2017). The lowest oxygen fugacity at which diamonds may occur corresponds to the formation of solid carbon after the following reaction (Harris and Gurney 1979): Mg2 SiO4 + SiC + O2 → 2MgSiO3 + C For

Muas

En

(6.4)

In this stage, the interaction of nitrogen with methane may be a possible reaction for diamond formation (Simakov 1983): N2 + 1.5CH4 → 1.5C + 2NH3

(6.5)

Nitrogen is included in the diamond structure by the following reaction (Sobolev et al. 1966): NH3 → N+3 + 3H−

(6.6)

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The decomposition rate of NH3 to H− and N3+ decreases with decreasing temperature. The ammonia concentration is inversely proportional to H2 O and CO2 (Simakov 2006). These facts agree with the tendency for nitrogen content in diamonds to decrease with increasing oxygen fugacity or temperature (Simakov 2006). In accordance with reaction (6.6), the NH3 content in the fluid decreases with decreasing N2 and CH4 contents. Meanwhile, N2 , NH3 , and CH4 contents all increase with depth (Simakov 1998, 2003; Sokol et al. 2017). In accordance with Kaminsky and Wirth (2017), major reservoirs of nitrogen on Earth should be expected in the core and in the lowermost mantle. On the other hand, nitrogen can be dissolved in mafic and ultramafic magmas in the presence of carbon under reduced conditions by the following reaction (Strehletov et al. 1990): Mg2 SiO4 + 3C + N2 = MgO + MgSiN2 + 3CO For Per Msn

(6.7)

Simakov (1998) suggested that nitrogen could be dissolved under very reduced conditions below the IW buffer, where ammonia is a more stable fluid phase. In the transition zone, most of the nitrogen could be dissolved in the melt. This is confirmed by the fact that sublithospheric diamonds include high proportions of type II diamonds characterized by low average nitrogen contents (Cartigny 2005). Therefore, it is possible to conclude that f O2 in the lithospheric upper mantle under the Archean cratons varies over a range of five to six log units, which agrees with the previously obtained results for spinel and garnet peridotite xenoliths (Ballhaus 1993; Simakov 1998, 2006). The upper mantle is zoned regarding the redox conditions, and the degree of its reduction state mainly increases with depth from the lithosphere to the asthenosphere. Meanwhile, the lithosphere may also show an opposite tendency. In addition, the same results show that the average calculated fluid compositions for peridotite and eclogite diamond inclusions are water-rich and close to the average composition of the gaseous inclusions in natural diamonds (Fig. 6.6). Diamond formation from fluids in the upper mantle is generally related to aqueous fluids containing a minimum percent of atomic carbon formed at pressures and temperatures corresponding to “cold” geotherms.

6.4 Diamond Formation in the Postmagmatic Processes of Kimberlites After magma crystallization and fluid separation, the processes of postmagmatic alteration and replacement reactions, including serpentinization, chloritization, sulfidization, and carbonatization, occurred. Replacement reactions, such as serpentinization, are particularly favored if external aqueous fluids are abundant, as can occur in the oceanic lithosphere. As mentioned above, secondary mineralization also occurs in the open and sealed cracks of diamonds. It includes minerals formed

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Fig. 6.6 Average compositions of the calculated fluids for African, Australian, South American, and Yakutian peridotite and eclogite diamond inclusions (1) and diamond-bearing xenoliths (2) (data from Simakov (1998, 2006)) in comparison with the composition of fluid extracted from South African, Brazilian, and USA diamonds (3) (data from Giardini and Melton (1975))

both in oxidizing conditions (e.g., calcite, dolomite, hematite, magnetite, etc.) and reducing conditions, such as graphite and sulfides. Serpentine has been described in diamonds many times. Mg–Fe silicate with a structure similar to that of serpentine was identified in the fibrous Yakutian diamond (Klein-Bendavid et al. 2007) formation, which could be linked to the last kimberlite magmatic event (Moore 2014; Cartigny et al. 2014). There are three main serpentine minerals: lizardite, chrysotile, and antigorite (Schwartz et al. 2013). Experimental studies show that antigorite is stable under relatively high-pressure conditions up to 60 kbar (Ulmer and Trommsdorff 1999). Maximum P–T parameters of antigorite stability were detected for silica-enriched serpentinites from Cerro del Almirez (southern Spain) formed at 680 °C and 19 kbar, serpentinized ultramafic rocks of the Chinese southwestern Tianshan metamorphic belt formed at 37 ± 7 kbar and 510– 530 °C and silica-enriched serpentinites from Cerro del Almirez (southern Spain) formed at 680 °C and 19 kbar (Padron-Navarta et al. 2010; Shen et al. 2015). As mentioned in Chap. 3, the process of serpentinization (i.e., the hydration of Mg-rich olivine and orthopyroxene) could lead to the formation of hydrogen (reaction 6.8) and hence hydrocarbons and other organic compounds. 3Fe2 SiO4 + 3Mg2 SiO4 + 6H2 O → 2Mg3 Si2 O5 (OH)4 + 2Fe3 O4 + 2H2 + 2SiO2 Fa

Fo

Srp

Mgt

(6.8)

Serpentine and magnetite were indeed detected in the epigenetic diamond inclusions (Meyer and McCallum 1986). As mentioned above, the dihydrogen byproduct can lead to the formation of different hydrocarbons, generally with low molecular weights, through Fischer–Tropsch-type (FT-t) organic synthesis: nCO2 + (3n + 1)H2 → Cn H(2n+2) + 2nH2 O there 1 < = n < + ∞

(6.9)

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6 Diamonds in Kimberlites and Their Xenoliths: A Reappraisal

As shown in Chap. 3 (Fig. 3.1b), the compositions of hydrocarbon and oil–water solutions formed via these reactions can approach the hydrothermal nanodiamond domain on the C-H–O diagram and can be the sources for nanodiamond formation (see Chap. 2). It is interesting to note that yearly Giardini et al. (1982) came to the conclusion that the genesis of kimberlitic diamonds and oil are connected with each other, and Vasiliev et al. (1968) highlight association of some Yakutian kimberlite pipes with crude oil and hydrocarbons. The main role in the FT-t process is played by catalysts, which are native metals and oxides of transition metals (Foustoukos and Seyfried 2004; Fu et al. 2007; Lazar et al. 2012; Szatmari 1989). The participation of sulfides may also catalyze the processes of organic formation (e.g., Lazar et al. 2012; Scribano et al. 2019), as sulfur at concentrations lower than 0.2 wt% acts as an activity promoter for FT-t catalysts (e.g., Manuella et al. 2018). Looking into more details, carbon dioxide reacts with H2, giving rise to CO, which binds to the catalyst surface, forming carbonyl groups (C = O). The latter is reduced to carbide and then to methylene groups (=CH2 ), which, in some cases, become methyl groups (−CH3 ) (McCollom and Seewald 2007). CH3 detected in the ND-bearing asphaltene patches in serpentinite xenoliths from Sicily (see Sect. 4.1) could be an important growth precursor in diamond synthesis at low P–T parameters (see Chap. 1). The formation of methanol, alcohol, hydrocarbons, ketones, fluoromethyl (C7 H7 F) benzene, nitrogenated organics containing F and Cl, and carboxylic and carbonic acids detected in diamonds (Melton and Giardiny 1974, 1975; Sobolev et al. 2018, 2019a,b) also takes place during these processes (Manuella 2018; Shock and Canovas 2010; Tomilenko et al. 2022). Methanol synthesis is possible at 320–400 °C and 0.1–0.3 kbar by the following reaction: CO2 + 3H2 → CH3 OH + H2 O

(6.10)

Alcohol can be synthesized from methanol by the next reaction: CH3 OH + CO + 2H2 → C2 H5 OH + H2 + 0.5O2

(6.11)

As shown in Chap. 3, the formation of nano- and microdiamonds from alcohol is possible at low P–T values corresponding to serpentine formation. Nanodiamond formation is possible at 220 °C and the saturated vapor pressure from water-organic solutions (Alzahrani and Alkahtani 2023). The growth of diamonds on the seeds is possible at temperatures of 100–350 °C at pressures of 0.1–0.4 kbar from waterorganic solutions (Borshevsky 1995). Serpentinization leads to nano- and micronsized diamond formation in the range of temperatures of 150–650 °C and pressures of 1–16 kbar (Simakov et al. 2015; Farré-de-Pablo et al. 2018; Pujol-Solà et al. 2020). The epigenetic graphite present in the cracks of the diamonds can also be formed during the processes of olivine serpentinization in kimberlites under reduced conditions (Pasteris 1981). Sverjensky et al. (2014) proposed a new quantitative theory of diamond formation as a consequence of the reaction of deep fluids with the rock types that they encounter during migration. Diamond can form at the drop in pH during water–rock interactions

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without changes in oxidation state. The model shows that fluid can react irreversibly with eclogite, generating diamond and secondary minerals due to a decrease in pH at almost constant oxygen fugacity. In accordance with their suggestion, diamonds may precipitate from fluids containing organic C species as follows: CH3 COO− + H+ → 2C + 2H2 O

(6.12)

Some pyrite desulfidization reactions are possible in hydrothermal systems (Scribano et al. 2019). The anodic reaction produces ferrous and sulfate ions (Caldeira et al. 2010): FeS2 + 8H2 O → Fe+2 + 2SO2− + 16H+ + 14e− 4 Pyr

(6.13)

Reaction (6.13) leads to a decrease in pH and could follow carbon (diamond) precipitation from organic components. It is possible to conclude that serpentinization and other types of secondary alterations in kimberlites can lead to nano- and microdiamond formation and the growth of diamonds on the seeds of mantle crystals at the depth of the upper mantle and Earth’s crust. On the other hand, serpentinization may lead to diamond dissolution processes at the contact of reduced hydrogen fluids with crystals and methane formation. CH4 generation at shallow depths is discussed in Etiope and Sherwood Lollar (2013). The formation of either methane or free carbon in the C-H–O system depends upon the P–T- f O2 parameters. Methane could be formed from free carbon and water by reaction (6.3) under decreasing pressure and oxygen fugacity (Simakov 2018a). The formation of methane, organics, and diamond could be connected with sulfides. They can catalyze organic and methane formations via FT-t reactions as a consequence of the serpentinization process. Then, in a water-rich environment, desulfidization reactions (6.13) lead to a pH decrease and possible diamond precipitation by reaction (6.12). This means that sulfides can change the pH regime in hydrothermal systems. In accordance with the model of Sverjensky and Huang (2015), the pH-change mechanism for diamond formation represents a mechanism in addition to potential changes in redox, temperature, and pressure. As already mentioned, optimum conditions for diamond growth correspond to water-rich fluids containing a minimum percent of atomic carbon. Indeed, combined pH decreases and f O2 increases could result in multiple cycles of diamond precipitation and dissolution. Thus, it is possible to highlight the third hydrothermal kimberlitic stage of diamond formation in the shallow mantle and Earth crust at P–T corresponding to the serpentine stability field (Simakov and Stegnitskiiy 2022). It is associated with secondary mineral formation from primary minerals of the kimberlite and therein xenoliths. In this stage, fluid compositions can vary from reduced to oxidized conditions, and the formation of various organics takes place. They are more depleted in nitrogen and ammonia than the upper mantle ones and mainly correspond to the C-H–O-S gaseous system. This stage explains the presence of methanol, alcohol,

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ketones, carbon acids, fluor- and chlorin-containing arenes, nitrogenated organics, and other organics in the fluids extracted from diamonds (Sobolev et al. 2018, 2019b) and the increase in organic matter in the diamonds in the later stage of their formation (Khachatryan 2017). The P–T of antigorite stability corresponds to the P–T of the experimental diamond growth from organic matter and to the P–T of micron-sized diamond formation in the chromite ores from the serpentinization of Tehuitzingo and Potosi serpentinites (Fig. 6.7). The formation of nano- and microdiamonds from fluids and the growth of previously formed diamond crystals are both possible at this stage, significantly influencing the total kimberlite diamond grade. It is confirmed, for instance, by the relationships between micro- and macrodiamonds in individual deposits (Krebs et al. 2016; Pattison and Levinson 1995) and the positive correlation of diamond grade and serpentine on Katoca pipe (Stegnitskiy 2006). The diamonds formed at the hydrothermal kimberlitic stage from organics should be characterized by a light carbon isotopic composition. More precisely, the δ13 C of serpentinite nanodiamonds formed in the range of 150–300 °C approximates −29‰ (Simakov et al. 2015). Interestingly, diamonds with close δ13 C values are known for Siberian, Ural, South African, Canadian, and Angolan crystals with eclogitic and unknown origins (Cartigny 2005; Kosman et al. 2016; Sobolev et al. 2018, 2019b; Stachel et al. 2022). Close to serpentinite nanodiamond δ13 C values are also known for Arkansas microdiamonds that are more enriched in light carbon than macrodiamonds (McCandless et al. 1991). It is important to note that Jericho diamonds, whose mineralogical association includes sulfides, are enriched in light carbon and have low nitrogen content (De Stefano et al. 2009), which corresponds to extralarge type IIa diamonds.

Fig. 6.7 P–T plot of antigorite stability, fields of microdiamond formation in chromites, and experimental data of diamond formations from OM. Legend: I—Antg-upper limit of antigorite stability, II—lizardite–antigorite boundary in accordance with (Evans 2004; Schwartz et al. 2013; Ulmer and Trommsdorff 1999); 1, 2—fields of microdiamond formation in the Tehuitzingo and Potos chromites (Farré-de-Pablo et al. 2018; Pujol-Solà et al. 2020); 3,4—fields of experimental diamond formations from organic matter (Borshevsky 1995; Simakov et al. 2008)

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6.5 The Genesis of the Extralarge Type IIa Diamonds Currently, the genesis of the extralarge type IIa diamonds of more than 100 carats still raises many contentious questions. They include the largest gemstone, such as the Cullinan (3106 ct), extracted from the Cullinan kimberlite and the famous alluvial stones from India, including the famous Koh-I-Noor. These large diamonds mainly occur in African kimberlite mines (e.g., Cullinan, Monastery and Jagersfontein, Letseng la Terai, Mothae, Kao, AK6 (now Karowe), Jwaneng (Orapa)), Canada (Jericho), and India (Gurney and Helmstaedt 2012; Moore 2009, 2014). Several stones were found in Yakutian pipes Nurbinskaya, Mir, Zarnitsa, and Udachnaya. Their main features are as follows: high degree of resorption; low δ13 C; irregular shape, occasionally with poorly preserved dodecahedral faces; low nitrogen contents (N < 10 ppm); absence of mantle-derived silicate and oxide inclusions; presence of sulfides and graphite; and nonassociation with small diamonds (Bowen et al. 2009; Gaillou et al. 2012; Gurney and Helmstaedt 2012; Levinson et al. 1992; Moore 2009). Their δ13 C (‰) signature mode spreads in the range of −17 to −21‰ between the average signatures of upper mantle diamonds and nanodiamonds hosted in a serpentinite xenolith and abiotic methane generated by serpentization of ophiolite massifs (Fig. 6.8). Fig. 6.8 Carbon isotope diagram of type IIa diamonds, typical mantle diamonds, transition-zone diamonds, serpentinite nanodiamonds, and abiotic methane generated by serpentization of the Othrys ophiolite. The data were taken from Banas et al. (2017), Cartigny et al. (2014), Etiope et al. (2013), Milledge et al. (1983), Simakov et al. (2015), and Tappert et al. (2005)

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Moore (2009, 2014) argues for the existence of multiple sources of type II diamonds and the websterite diamond suite as the main source for extralarge diamonds. Smith et al. (2016) proposed their formation in Fe–Ni-C-S melts in deep mantle eclogites. The presence of majoritic garnet and calcium silicate perovskite in the studied type II gem diamonds with sizes from 5 to 40 carats could show their formation at depths of 360–750 km. Meanwhile, the total absence of silicon in their metal inclusions (table S2 in Smith et al. 2016) could show their formation at relatively low pressure (Takafuji and Hirose 2005). As noted above, diamond can crystallize in the mantle with low sulfur concentrations and is accompanied by sulfide and metallic minerals (Zhimulev et al. 2012), which is consistent with the rare meteorite paragenesis of diamond (Arai 1986). However, experiments have shown that the formation of diamonds greater than 1000 carats requires the presence of iron bubbles dozens of centimeters in size (Borzdov et al. 2000; Sumiya et al. 2001). As a result, it is possible to conclude that the model of Smith et al. (2016) cannot explain the distinctive features of these extralarge gems (i.e., the absence of mantle-derived silicate and oxide inclusions in them, their integrity and lack of association with smaller size diamonds). Shirey et al. (2020) proposed that their formation is due to the coalescence of mantle fluids in large quasiplanar voids. From Ballhaus and Frost (1994), it follows that the formation of voids is possible in the late stage of kimberlitic magma crystallization during fluid discharge from melts. Currently, the Cullinan pipe represents the ideal site to investigate the origin of large diamond stones since it contains the highest worldwide quantity of famous and most studied large gem diamonds (Gurney and Helmstaedt 2012). Several stages of diamond formation can be shown for Cullinan eclogitic diamond inclusions. As previously mentioned, Nimis et al. (2020) concluded that Cullinan diamonds “contain different inclusion types, may have very different ages (even within the same kimberlite), and may have been formed by different processes (e.g., fluid-driven vs. melt-driven) under different thermal regimes.” Moreover, Korolev et al. (2018) concluded that Cullinan eclogites may have formed as a result of several short-term heating events. P–T- f O2 parameters calculated for eclogite and peridotite inclusions in diamonds from the Cullinan pipe show two types of oxygen fugacity trajectories (Fig. 6.9). The first one corresponds to diamonds with a high content of nitrogen under low f O2 decreasing with depth (limb I). Their crystallization was lower under f O2 conditions than in the IW buffer in equilibrium with the methane–nitrogen-rich fluid of the C-H-N gaseous system. In accordance with the experimental data of Hirayam et al. (1993), these diamonds could be connected with Fe–C melts. The second one corresponds to diamonds with a relatively low content of nitrogen under relatively high f O2 decreasing with depth (limb II). Their crystallization occurred under redox conditions above the IW buffer in equilibrium with water- and carbon dioxide-rich fluids. The observed oxygen fugacity trajectories can be explained by mantle plume formation in the region of the Bushveld igneous province (Korolev et al. 2018). The Cullinan eclogites may have formed as a result of several subduction events that built the Kalahari craton (Korolev et al. 2018). Upper overlap in the Cullinan pipe (Khar’kiv et al. 1998) supports the pausing of kimberlite magmas. The data

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111

of Korolev et al. (2018) strengthen the IIa-diamond formation during short-term heating events. The occurrence of transition zone minerals in the Cullinan diamonds (majorite and CaSiO3 phases) suggests that plume melts could have intruded from deep levels of more than 300 km. The high temperature related to the first event reduced the plume melts and fluids with diamond seeds, which were already uplifted to the lithosphere from the great depth of the transition zone. The calculated P–T- f O2 parameters (limb I in Fig. 6.9) suggest that at this stage, the diamonds could have formed from metal-silicate melts. The possibility of the formation of type I diamonds with high nitrogen content from iron melts was shown in the experiments of Borzdov et al. (2002), Yu et al. (2008), Babich et al. (2012) and Palyanov et al. (2010, 2013). At the fast decrease in the P–T parameters, only CH4 and N2 discharged from the melts. This leads to a decrease in f O2 and an increase in the ammonia content in the fluid (Simakov 1998; Sokol et al. 2017) and, as a result, to an increase in the nitrogen content in the diamonds by reaction (6.5). During the second event (limb II on Fig. 6.9), the portion of reduced fluids with diamond nuclei was uplifted to the lithosphere from the asthenosphere. In this stage, fluids contact subducted rocks and carbonates (Taylor and Green 1989), which could lead to diamond formation from carbonate–silicate melts. During these processes, the reaction Dol + Coe = liq + CO2 could be realized, which follows the diamond formation in the field of WM buffer by the reaction: CH4 + CO2 → 2C + 2H2 O

(6.14)

With decreasing P–T parameters, CO2 and H2 O would also separate from the melt. Thus, these second-stage fluids are enriched in H2 O and CO2 and depleted in nitrogen and ammonia, which mainly corresponds to the C-H–O-N gaseous system. From the aforementioned data, two main events of Cullinan eclogite and peridotite diamond formation could be proposed in the mantle: (1) crystallization of diamonds under very low f O2 conditions connected with metal-silicate melts and (2) diamond crystallization and growth from fluid under higher f O2 conditions corresponding to water-rich fluids. The first corresponds to melt-driven processes, and the second corresponds to fluid-driven processes by Nimis et al. (2020). This scheme explains the presence of iron-nickel-carbon–sulfur melt, accompanied by a thin fluid layer of methane and majoritic garnet inclusions in type II diamonds (Smith et al. 2016). Additionally, it explains the presence of hydrous silicic fluid films in the Cullinan gem-quality diamonds (Nimis et al. 2016). The obtained P–T- f O2 parameters for Cullinan diamond inclusion formations show the relationship between their f O2 and nitrogen content of the diamonds. From the provided calculations, it follows that the nitrogen content of diamonds increases with decreasing oxygen fugacity (Fig. 6.10) and that the formation of type IIa diamonds is connected with fluid-driven processes. Cullinan diamond secondary minerals are similar to the set of secondary minerals in mantle xenoliths and kimberlites. Such a secondary mineral association is represented by amphibole (actinolite-tremolite series, edenite or hornblende), spinel,

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Fig. 6.9 T- f O2 and P- f O2 plots for the Cullinan eclogitic (1) and peridotitic (2) diamond inclusions. The P–T values were estimated using the garnet–clinopyroxene thermobarometers of Nakamura (2009) and Simakov (2008) and the garnet-clinopyroxene-silica oxygen barometer of Simakov (2006). For the peridotitic inclusions, the P–T- f O2 values were estimated using the garnet–olivine and garnet-(orthopyroxene) thermobarometers of O’Neill and Wood (1979) and Ryan et al. (1996) and the garnet–olivine-(orthopyroxene) oxygen barometer of Simakov (1998). The analyzes for the calculations were taken from Gurney et al. (1985) and Korolev et al. (2018). Data on the nitrogen content of Deines et al. (1989) and Korolev et al. (2018) in ppm are plotted Fig. 6.10 The relationship between calculated f O2 and nitrogen content (in ppm) for the Cullinan eclogitic diamonds. The symbols are the same in Fig. 6.9

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serpentine, talc, mica (phlogopite, biotite), microcline, apatite, and possibly sellaite (Korolev 2017; Korolev et al. 2018). They replace primary inclusions (pyroxene and olivine) and are localized in open and sealed cracks. P–T- f O2 parameters calculated for secondary hercinyte spinels indicate that their formation occurred near the websterite solidus at 780–870 °C and 13–15 kbar and at oxygen fugacities 1– 2 units lower than the QFM buffer. It corresponds to carbon dioxide and watercarbon dioxide fluid compositions of the C-H–O gaseous system of Zhang and Duan (2010) (Fig. 6.11). The obtained parameters coincide with the calculated P–T- f O2 for Cullinan spinel peridotite xenoliths (Simakov 2018a) widespread in the Cullinan pipe (e.g., Khar’kiv et al. 1998). They also coincide with metasomatic events recorded in a spinel-harzburgite xenolith from the Kimberley cluster of the central Kaapvaal Craton estimated by (Konzett et al. 2013) at 750–760 °C at 30 kbar and at oxygen fugacity in the upper QFM buffer on 0.9–1.2 log units. This fact indicates that in the final stages, the Cullinan kimberlite magma was enriched in H2 O and CO2 volatiles. In contrast to the old age of many eclogitic diamond suites (e.g., Richardson 1986), the Cullinan eclogitic diamond suite is identical to the age of the host kimberlite (Smith 1983; Navon 1999). Crystallization of such “young” eclogitic diamonds is logically also linked to the kimberlite magmatic event. It is interesting to note that the percentage of large type IIa diamonds in the worldwide diamonds (2%) corresponds to the percentage of world diamond inclusions formed under P–T conditions in the range of the websterite solidus at 630–680 °C and 30–34 kbar (Stachel and Luth 2015, Fig. 6.2). On the other hand, estimated pressures and temperatures coincide with the field of serpentine stability (Fig. 6.7). Smart et al. (2011) proposed Jericho high-Mg0 eclogite diamond growth from methanogenically mediated organic matter. Harte and Gurney (1981) and Moore (2009, 2014) also proposed that diamond megacrysts formed in volatile-enriched pegmatitic veins at the time of or shortly prior to the eruption of the host kimberlite and represent very late-crystallizing megacryst phases. Metasomatic events observed at the boundary of the host kimberlites and intrusive traps in the Nyurbinskaya pipe (Khar’kiv et al. 1998) could confirm these processes. According to Speciuss et al. (2017), Nyurbinskaya’s diamonds are characterized by multistage growth from initial parental mantle fluids to later metasomatic fluids with lighter carbon. As shown before, the enrichment of fluids with relatively light carbon may be due to abiogenic organic formation in the hydrothermal stage of kimberlite alteration. P–T- f O2 parameters calculated for Monastery Opx, Cpx, Opx-Cpx, Ilm-Opx megacrysts and lherzolite xenoliths indicate that their formation occurred in the websterite solidus and antigorite stability field at 650–850 °C and 25–40 kbar (Fig. 6.12a). Interestingly, the obtained P–T parameters coincide with the estimated P–T for Orapa websterite diamond inclusions (Fig. 6.12a), Jericho spinel-garnet coarse peridotite xenoliths (Kopylova et al. 1999) and with the temperatures of the larnite and wollastanite formation (520–670 °C) discovered in type IIb diamonds (Moore and Helmstaedt 2019; Smith et al. 2018). In accordance with models by Haggerty (1975) and Gurney et al. (1993), variations in MgO and Cr2 O3 compositions in ilmenites correspond to the conditions of diamond stability that coincide with the obtained f O2 parameters for Monastery ilmenite-orthopyroxene megacrysts

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Fig. 6.11 P–T- f O2 plots of Cullinan secondary hercinyte spinel relative to websteritic (Mysen and Boettcher 1976) solidus. The P–T was estimated using the Spl-Ol thermometer of O’Neill and Wall (1987) using the model of “hypothetical olivine” (Simakov 2014) with fictive olivine composition of Fo75 corresponding to Cullinan diamond inclusion (Korolev et al. 2018). For the pressure estimation, the spinel barometer of Marakushev (1984) was used (A). f O2 were estimated for the spinels on the basis of the Spl-Opx-Ol oxybarometer of Ballhaus et al. (1991) and the model of the C-H–O gaseous system of Zhang and Duan (2010) using the same procedure. The calculated P–T- f O2 values correspond to the ranges of 780–870 °C and 13–15 kbar and to water-carbon dioxide compositions compatible with the C-H–O gaseous system. Legend: Antg, WS, QFM, WM, IW, CCO, CO2 -H2 O, and H2 O-CH4 —the same as in Figs. 6.1, 6.2, 6.5, and 6.7

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(Fig. 6.12b). The calculated f O2 values of the Ilm-Opx megacryst formation correspond to methane and water-rich fluids in the range of 650–800 °C and coincide with the oxygen fugacity of the Grt-Ol Orapa websterite diamond inclusion, while the f O2 conditions of peridotite xenoliths correspond to CO2 -rich fluids at approximately 1100 °C. In particular, the estimated P–T- f O2 parameters indicate that the degree of reduction decreases with depth in the final stage of kimberlitic magma formation (Fig. 6.12b). This contradicts previous estimates of the oxygen-reduction conditions of mantle peridotite parageneses (Ballhaus 1993; Simakov 1998, 2006; Stachel and Luth 2015). Considering the rarity of silicate inclusions, extralarge diamonds could have formed under P–T conditions near or below the websterite solidus on the body of the hydrothermal kimberlitic stage. As shown above, methane and other organics can be formed at the hydrothermal kimberlitic stage of diamond formation under reduced conditions. The formation of methane, organics, and diamond could be connected with sulfides. They can catalyze organic and methane formations via FT-t reactions during serpentinization. Then, in a water-rich environment, desulfidization reactions (6.13) lead to a pH decrease and possible diamond precipitation by reaction (6.12). This means that sulfides can change the pH regime in hydrothermal systems. In accordance with the model of Sverjensky and Huang (2015), the pH-change mechanism for diamond formation represents a mechanism in addition to potential changes in redox, T, and P. Indeed, combined pH decreases and f O2 increases could result in multiple cycles of diamond precipitation and dissolution, which explain extralarge type IIa diamond resorption features. On the other hand, it is possible to conclude that the formation of Cullinan extralarge type IIa diamonds must occur under relatively oxidized conditions (Fig. 6.10). The formation of either reduction or oxygen fluid components in the C-H–O system depends upon the P–T- f O2 parameters. Methane could be formed by reaction (6.3) as the pressure and oxygen fugacity decrease, while water and free carbon, on the contrary, increase (Fig. 6.13). Cyclical variations in P–T- f O2 conditions and fluid compositions in the system may cause resorption and dissolution of small diamonds during cyclical growth-dissolution processes, which leads to the gradual build-up of large stones (Zhao et al. 1997). Diamonds growing in a fluid can increase at the final stages of kimberlite pipe formation with increasing free carbon precipitation from C-H–O fluids with decreasing temperature (Stachel and Luth 2015; Simakov 2018b). This process is very likely to the interaction between a new portion of intruded kimberlitic magma and formed serpentinites, analogous to the serpentinite diatremes in the Navajo Volcanic Field (Colorado Plateau) (Smith 2010). Such an interaction could have made available huge amounts of fluids in quasiplanar voids necessary for the extralarge diamond and pegmatitic vein formations by the reaction (6.14). The presence of upper overlaps of the flush of mantle fluids in the Orapa (Field et al. 1995) and the Premier (Khar’kiv et al. 1998) pipes leads to the local fluid P–T increase up to the diamond thermodynamic stability. The formation of methane, organics, and diamond could be connected with sulfides. They can catalyze organic and methane formations via FT-t reactions during serpentinization. Then, in a water-rich environment, desulfidization reactions (6.13) lead to a pH decrease and possible diamond precipitation by reaction (6.12). This

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Fig. 6.12 P–T- f O2 plots of Monastery Opx (◯), Cpx (●), Ilm-Opx (☐), Cpx-Opx (∎) megacrysts and lherzolite (Δ) formation in the final stage of magma crystallization and Orapa Grt-Opx websterite diamond inclusions (▲) relative to the diamond–graphite boundary (Bundy et al. 1961), lherzolite-water (Wyllie and Ryabchikov 2000) and websteritic (Mysen and Boettcher 1976) soliduses and antigorite stability. The P–T values were estimated using the clinopyroxene thermobarometers of Nimis and Taylor (2000), orthopyroxene thermobarometers of Carswell (1991) and Nimis and Grutter (2010), garnet-orthopyroxene thermobarometers of Harley (1984) and Nickel and Green (1985), a modified version of the ilmenite-(olivine) thermometer of Andersen and Lindsley (1981) using the model of “hypothetical olivine” (Simakov 2014) (For = 83) and the clinopyroxeneorthopyroxene thermometer of Taylor (1998) (A). f O2 were estimated using the garnet–olivine(orthopyroxene) and garnet–orthopyroxene-(olivine) oxygen barometers of Simakov (1998) and modified ilmenite–orthopyroxene-(olivine) oxygen barometers of Eggler (1983) at 30 kbar. The analyzes for the calculations were taken from Jacob (1977) and Gurney et al. (1984). Legend: Antg, LS, WS, QFM, WM, IW, CCO, CO2 -H2 O, and H2 O-CH4 —the same as in Figs. 6.1, 6.2, 6.5, and 6.7

means that sulfides can change the pH regime in hydrothermal systems. In accordance with the model of Sverjensky and Huang (2015), the pH-change mechanism for diamond formation represents a mechanism in addition to potential changes in redox, T and P. Indeed, combined pH decreases and f O2 increases could result in multiple cycles of diamond precipitation and dissolution, which explain extralarge type IIa diamond resorption features.

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Fig. 6.13 Changes in the carbon content of diamond-saturated fluid with oxygen fugacity, P and T variations. Solid lines—molar parts of the main components of the C–H–O fluid system, dashed line—atomic percent in the fluid

From the obtained P–T- f O2 parameters for Monastery megacrysts and peridotite nodules, it is possible to conclude that the formation of extralarge stones was connected from one hand with oxidized fluid separated from crystallized kimberlite magma and from another hand with reduced organics formed at the hydrothermal stage by the reaction (6.14) (Fig. 6.14). Smart et al. (2011) also proposed that oxidized and reduced fluids occur during the formation of Jericho diamonds. In accordance with the aforementioned data, it is possible to conclude that extralarge type IIa diamonds have multiple origins from the lower mantle to crustal depths. Their main stage of formation took place on the boundary of the lithosphere and crust from separated fluid voids at the P–T corresponding to the range of websteritic solidus from the pegmatitic veins solidified along the contacts of kimberlite magma and crustal rocks. Their main sources were reduced hydrocarbons and other organics formed at the hydrothermal stage of kimberlite formation and discharged oxidized fluids from crystallized kimberlite magma. The presence of sulfides catalyzes their formation. The presented model explains the main distinctive features of the aforementioned extralarge diamonds (great sizes, absence of mantle silicate inclusions, integrity, low δ13 C signatures, and nitrogen contents, high degree of resorption, presence of sulfides and nonassociation with small diamonds).

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Fig. 6.14 Scheme of giant type IIa (octahedron) and volcanic methane (square) formations in accordance with variations in carbon content in diamond-saturated fluid from pressure

6.6 Conclusions As a result, three main stages of diamond formation associated with kimberlites in the mantle should be considered (Fig. 6.15): 1. Asthenospheric melt-driven stage of diamond formation at depths greater than 200 km. This corresponds to the crystallization of the initial diamonds in the asthenosphere mainly under very low f O2 conditions from metal-silicate melts in equilibrium with methane–nitrogen-rich fluids of the C-H-N gaseous system. At this stage, the main part of micron-sized diamonds containing silicate inclusions was formed.

Fig. 6.15 Three main stages of diamond formation in the mantle and the Earth’s crust. I—asthenospheric, II—lithospheric, III—mantle–crustal relative to diamond–graphite boundary (Bundy et al. 1961). Solid lines: A—field of known P–T parameters obtained for diamond peridotitic and eclogitic inclusions (see Fig. 6.1), B—possible field of hydrothermal-driven postmagmatic stage of diamond formation in accordance with Fig. 6.7. Legend: 1—metal-silicate melts, 2—carbonate melts

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2. Lithospheric fluid-driven stage of diamond formation at depths mainly from 70 to 200 km. This corresponds to the crystallization of diamonds with lherzolite and eclogite parageneses in equilibrium with water- and carbon dioxide-rich fluids from the C-H–O-N gaseous system and carbonate–silicate melts. At this stage, the main parts of mantle diamonds were formed. 3. Shallow mantle/crustal hydrothermal-driven postmagmatic stage of diamond formation at depths lower than 70 km on the final stage of kimberlite formation and alteration. It corresponds to the secondary mineral formations by the initial kimberlite and xenolith minerals. The growth and dissolution of previously formed diamond crystals and the formation of nano- and microdiamonds from hydrothermal C-H–O-S fluids occurred at this stage.

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Chapter 7

Conclusions

Experimental data and theoretical models reported in a vast, well-established literature, and summarized in this monograph are consistent with the notion that diamond is a typical deep-seated mineral that is often related to kimberlites and mantle xenoliths. On the other hand, a large set of data suggests totally different scenarios for the formation of diamonds, especially those at the nanoscale. Experimental results indicate that nanocarbon formation occurs at relatively low P–T conditions from organic matter. Nanodiamonds and diamond-like phases were synthesized at 500 °C and a total pressure of nearly 1 kbar in a fluid composed of nitrogen-bearing organic matter (Simakov et al. 2008). More precisely, diamonds were identified as 70–80 nm-sized particles of different forms; some of them reached 1 μm. Multiwall fullerite structures and typical carbon nanotubes were also synthesized at 700–750 °C and 5 kbar from organic matter (Simakov et al. 2001). Nanosized diamond particles can crystallize naturally under conditions of relatively low pressure and temperature, such as in interstellar clouds and protoplanetary disks. Their formation is associated with diamond-like nanocarbon particles called “diamondoids.” Moreover, nanodiamond formation is possible at the catagenesis and metagenesis stages at depths ranging from 3 km (in the case of mobile belts) to 6 km (Simakov 2018) (Fig. 7.1). In this respect, ultrananocrystalline diamonds can be connected with coals formed during regional and contact metamorphism. In addition, micron-sized diamonds can be formed during alteration hydrothermal-metasomatic processes under P–T conditions corresponding to greenschist facies at depths of 8–25 km (Fig. 7.1). Rare forms of carbon-multiwall fullerites and nanotubes could be formed from gaseous hydrocarbons at P–T parameters corresponding to contact metamorphic conditions. Their origin is connected with the hydrocarbon-rich fluids that are formed in contact metamorphism. Nano- and micron-sized diamonds could also form during the exothermic hydration reactions and other hydrothermal mineralization processes affecting the mafic and ultramafic rocks (e.g., Simakov et al. 2015). In particular, nanodiamonds and micron-sized diamonds can form in the oceanic lithosphere during serpentinization of tectonically uplifted portions of the shallow mantle. Diamond formation is also © The Author(s), under exclusive license to Springer Nature Switzerland AG 2023 S. Simakov et al., Nano and Micro Diamond Formation in Nature, SpringerBriefs in Earth Sciences, https://doi.org/10.1007/978-3-031-43278-1_7

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Fig. 7.1 Lower boundaries of nano, micro-, and macrosized formations in the geosphere in accordance with the models of Mitchell (1991) and Tissot and Welte (1978). 1—diamondoids, 2—nanosized diamond space particles, 3—nanodiamond, 4—microdiamond, 5—macrodiamond, 6—mantle peridotite xenoliths, 7—mantle eclogite xenoliths, 8—subducted slab (reproduced from Fig. 8 of Simakov 2018)

possible at the late stage of kimberlitic eruptive events on the shallow upper mantle and the overlaying crust. Such diamond formation is associated with secondary mineral formations by primary kimberlite and xenolith minerals. Hydrocarbons and nitrogen play an important role in the processes of diamond formation, especially for micron-sized and nanosized types. Regarding the first point, there is an increase in depletion in heavy C isotopes from mantle macrodiamonds, corresponding mainly to carbonate δ13 C values, to space nanodiamond particles, corresponding to oil-like δ13 C values (Fig. 7.2a). Accordingly, there is an increase in nitrogen content in the same direction (Fig. 7.2b). Moreover, the introduction of nitrogen atoms to the diamond structure leads to the stabilization of micron- and nanosized diamonds in the graphite stability field. Currently, the creation of new methods of nanocarbon synthesis is one of the urgent tasks of modern technologies. The possibilities of hydrothermal diamond growth are known (Szymanski et al. 1995; Roy et al. 1996). The results presented in our Monograph could be developed for the new approach of hydrothermal nanodiamond synthesis at low pressures and temperatures. Chapters 2–4 show nanodiamond formation in the experiments and natural processes at low P–T under hydrothermal

7 Conclusions

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Fig. 7.2 Variations in prominent modes of nitrogen content (in ppm) (a) and carbon isotopic composition (b) for the space nanosized diamond particles (I) metamorphic microdiamonds (II), secondary alteration micro- and nanodiamonds (III), and mantle micro- and macrosized (IV) diamonds (reproduced from Fig. 1 of Simakov 2018)

conditions from organic matter. Water–oil solutions are formed in the range of 300– 400 °C at pressures greater than 220 bar (Chekaluk and Filyas 1977; Price 1976) and could lead to nanodiamond synthesis (Simakov 2010). The formation of nanodiamonds in the “petroleum asphaltenes” of serpentinites described in Chap. 4 confirms this. As a result, our study could lead to the creation of a low-energy and ecologically sustainable industrial method of nanodiamond formation for the utilization of oil rests at P–T parameters corresponding to the processes of oil and oil rest refinement. Diamond dimensions generally increase with the growth of pressures and temperatures of the parent rocks. On the basis of geological, mineralogical, geochemical, and experimental data, it is possible to distinguish four main groups of natural diamonds. Each type has specific characteristics (Fig. 7.2). The first one comprises the interstellar nanodiamond particles whose formation is associated with the diamondoids in space objects. The second group includes nanodiamonds and microndiamonds associated with secondary (hydrothermal) alteration and replacement reactions in mafic rocks. The third type includes crustal micron- and nanometer-scale diamonds associated with coals, sediments, and metamorphic rocks. The fourth type includes macro, micron- and nanosized diamonds that are associated with kimberlites, lamproites and their peridotite and eclogite xenoliths. Kimberlitic and xenolitic macrodiamonds can have multiple origins. Their main part was formed from the C–H–O–N gaseous system. The petrologic model of diamond formation from this system presented in Chap. 6 of Monograph leads to the creation of new mineralogical-petrologic methods for the estimation of the diamond potential of kimberlite and lamproite deposits (Simakov and Stegnitskiy 2021; Simakov and Ustinov 2023). It is therefore possible to conclude that diamonds can form in nature in a wide range of pressures and temperatures both at equilibrium and disequilibrium conditions in

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the fields of diamond and graphite stability. Nanodiamonds formed in a carbonbearing magmatic or metamorphic fluid at low P–T may be nucleuses for larger diamond crystals, especially under transient high-pressure events (Simakov 2018). Acknowledgements We wish to thank M. Baidakova, E. Bagrii, G. Bulanova, D. DolivoDobrovolsky, V. Dubinchuk (deceased), F. Kaminsky, G. Khachatryan, K. Kiseeva, A. Kouchi, M. Magomedov, I. Mahotkin, F. Manuella, A. Moore, F. Pirajno, S. Sablukov, O. Shestakova, A. Shiryaev, C. Smith, N. Sobolev (deceased), B. Spitsyn, Y. Stegnitskiy, and A. Sukhanov (deceased) for their stimulating discussions, comments, and help.

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