# Recycling Subducted Organic Carbon as Diamonds: An Example From the New Caledonia Forearc Ophiolite

Jonathan C Aitchison, Daniel Patias, Dominique Cluzel, Trevor R Ireland, Renjie Zhou, Dongyang Lian, Jingsui Yang, Zhen Yan

► To cite this version:

Jonathan C Aitchison, Daniel Patias, Dominique Cluzel, Trevor R Ireland, Renjie Zhou, et al.. Recy- cling Subducted Organic Carbon as Diamonds: An Example From the New Caledonia Forearc Ophi- olite. Geochemistry, Geophysics, Geosystems, 2025, 26 (4), 10.1029/2025GC012250. hal-05045130

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Distributed under a Creative Commons Attribution 4.0 International License
# Geochemistry, Geophysics, Geosystems®

## RESEARCH ARTICLE

10.1029/2025GC012250

**Key Points:**

* Strongly negative δ¹³C values of New Caledonian ophiolitic diamonds and moissanite suggest an organically fractionated carbon source

* Localized super-reducing conditions associated with abundant subducting terrestrial C3 plant material likely favored mineral formation

* Global significance of supra-subduction zone ophiolites as an unconventional setting for diamond genesis is recognized

**Supporting Information:**

Supporting Information may be found in the online version of this article.

**Correspondence to:**

J. C. Aitchison,
jona@uq.edu.au

**Citation:**

Aitchison, J. C., Patias, D., Cluzel, D., Ireland, T. R., Zhou, R., Lian, D., et al. (2025). Recycling subducted organic carbon as diamonds: An example from the New Caledonia forearc ophiolite. *Geochemistry, Geophysics, Geosystems*, 26, e2025GC012250. https://doi.org/10.1029/2025GC012250

Received 19 FEB 2025
Accepted 8 APR 2025

**Author Contributions:**

**Conceptualization:** Jonathan C. Aitchison, Jingsui Yang

**Formal analysis:** Jonathan C. Aitchison, Daniel Patias, Trevor R. Ireland

**Funding acquisition:** Jonathan C. Aitchison

**Investigation:** Jonathan C. Aitchison, Daniel Patias, Dominique Cluzel, Trevor R. Ireland, Renjie Zhou, Dongyang Lian, Zhen Yan

**Methodology:** Jonathan C. Aitchison, Daniel Patias, Dominique Cluzel, Trevor R. Ireland, Dongyang Lian

© 2025 The Author(s). Geochemistry, Geophysics, Geosystems published by Wiley Periodicals LLC on behalf of American Geophysical Union. This is an open access article under the terms of the Creative Commons Attribution License, which permits use, distribution and reproduction in any medium, provided the original work is properly cited.

<figure><img src="image_2.png" /></figure>

## Recycling Subducted Organic Carbon as Diamonds: An Example From the New Caledonia Forearc Ophiolite

Jonathan C. Aitchison¹ ID, Daniel Patias¹ ID, Dominique Cluzel², Trevor R. Ireland¹, Renjie Zhou¹ ID, Dongyang Lian³,⁴ ID, Jingsui Yang³,⁴ ID, and Zhen Yan⁵ ID

¹School of the Environment, The University of Queensland, St Lucia, QLD, Australia, ²Institut des Sciences Exactes et Appliquées, Université de la Nouvelle-Calédonie, Noumea, New Caledonia, ³State Key Laboratory for Mineral Deposits Research, School of Earth Sciences and Engineering, Nanjing University, Nanjing, China, ⁴Key Laboratory of Deep-Earth Dynamics of Ministry of Natural Resources, Center for Advanced Research on Mantle (CARMA), Institute of Geology, Chinese Academy of Geological Sciences (CAGS), Beijing, China, ⁵Institute of Geology, Chinese Academy of Geological Sciences, Beijing, China

**Abstract** Micro-diamonds and moissanite (SiC) have been identified in ophiolitic mantle harzburgites and chromitites of the New Caledonian Peridotite Nappe. The pale yellow (100–250 µm) micro-diamonds and light blue moissanite (mean –26.5‰, range –33.5‰ to –23.8‰ and mean –26.9‰, range –31.8‰ to –25.6‰ respectively) exhibit consistently strong negative δ¹³C values consistent with vegetal (C3) photosynthesis. Preservation of U-Pb ages amongst co-occurring rutile xenocrysts, with a closure temperature of 620 ± 20°C, constrains the maximum thermal conditions experienced by these rocks. These temperatures indicate that the New Caledonian diamonds did not form under the deep mantle conditions typical of conventional diamond genesis but instead within a distinct supra-subduction zone (SSZ) forearc setting. The association with moissanite suggests formation within anoxic, organic carbon-rich sediments at the top of the subducting slab or within the subduction channel under localized super-reducing conditions. In light of mantle heterogeneity, extension of the known distribution of ophiolitic diamonds to the Southern Hemisphere supports interpretation of their formation in relation to an SSZ process rather than a deep mantle source. It also highlights a previously unrecognized aspect of the global carbon cycle, underscoring the significance of SSZ forearc ophiolites in deep carbon transport and transformation.

**Plain Language Summary** Geologists have discovered tiny diamonds and moissanite (a rare mineral made of silicon carbide) in mantle rocks from New Caledonia, marking the first such finding in the Southern Hemisphere. This supports the global presence of these minerals in special rocks called ophiolites that form in regions in front of an ocean island arc where one tectonic plate is subducted beneath another. The carbon in these micro-diamonds likely comes from organic materials, possibly originating from plants at Earth's surface. The presence of moissanite (SiC) suggests that they formed in low-oxygen, carbon-rich environments. Interestingly, these micro-diamonds do not appear to have formed deep in the Earth's mantle, as is usually the case, but rather amongst carbon-rich sediments on top of a tectonic plate as it was subducted deep beneath an island arc. This discovery points to an important but previously overlooked aspect of the global carbon cycle in such regions.

## 1. Introduction

Diamonds, the most structurally ordered form of carbon, are typically associated with kimberlites and form under extreme pressure and temperature conditions (>950°C) at mantle depths >150 km (Cartigny, 2005; Shirey et al., 2013; Smit et al., 2022; Stachel & Luth, 2015; Yang et al., 2021). However, micro-diamonds hosted in ophiolite-hosted chromitites were first reported three decades ago from Luobusa in Tibet (Bai et al., 1993) along the Yarlung Tsangpo suture zone—an important structural feature marking the India-Asia collision. Since this initial discovery, numerous occurrences of micro-diamonds in oceanic mantle rocks, including chromitites and peridotites across various suture zones in the Northern Hemisphere including Albania, Turkey, Ladakh, Tibet, Myanmar, the Urals, the Altaids, Mexico and Cuba (Yang et al., 2021 and references therein) (see Figure S1 in Supporting Information S1) have been reported. These diamonds are often associated with ultra-high-pressure (UHP) minerals and super-reducing (SuR) phases such as Ni–Mn–Co alloys, Fe–Si and Fe–C phases, and moissanite (SiC) (Yang et al., 2014).
**Project administration:** Jonathan
C. Aitchison

**Resources:** Jonathan C. Aitchison,
Dongyang Lian, Jingsui Yang, Zhen Yan

**Supervision:** Jonathan C. Aitchison,
Renjie Zhou

**Validation:** Jonathan C. Aitchison,
Daniel Patias, Dominique Cluzel

**Writing – original draft:** Jonathan
C. Aitchison

**Writing – review & editing:**

Daniel Patias, Dominique Cluzel, Trevor
R. Ireland, Renjie Zhou, Dongyang Lian,
Jingsui Yang

Despite the growing number of reports on ophiolitic diamonds, their origin remains highly debated with many studies favoring a UHP genesis (Yang et al., 2021). However, their occurrence within supra-subduction zone (SSZ) forearc rocks, which have been emplaced onto passive continental margins via low-angle thrusting, suggests that they may have formed under conditions very different from the deep mantle environments typically associated with diamonds from subcontinental lithosphere. Well-characterized peridotitic (P-type) and eclogitic (E-type) diamonds, which are more commonly studied, generally exhibit carbon isotopic compositions dominated by primordial carbon (mean δ¹³C: P-type –4.9‰, E-type –5‰) (Stachel et al., 2022). In contrast, ophiolitic diamonds display significantly lighter δ¹³C values (~–25‰ (Lian & Yang, 2019)) although some eclogitic diamonds also show evidence of deep carbon cycling (Cartigny et al., 2014; Stachel et al., 2022). In ophiolitic settings such as Mexico and Cuba, smaller nano-diamonds have been inferred to have metastably precipitated from reduced C-O-H fluids in association with serpentinization (Farré-de-Pablo et al., 2019; Pujol-Solà et al., 2020).

To advance our understanding of the nature, origin, and global distribution of ophiolitic diamonds, we initiated an investigation on SW Pacific ophiolites in New Caledonia, New Zealand, and eastern Australia. This study presents our early findings, providing new insights into the formation and occurrence of diamonds in these regions.

## 2. Geological Setting

New Caledonia records a Paleogene interval of subduction in the Southwest Pacific, which culminated in an arc-continent collision (Aitchison et al., 1995; Cluzel et al., 2001; Maurizot, Cluzel, Patriat, et al., 2020). The main island, Grande Terre, stretches 450 km and is dominated by large ophiolite massifs (Figure 1).

These massifs, erosional remnants of mantle rocks (Peridotite Nappe), are interpreted as part of the Loyalty forearc mantle wedge (Cluzel et al., 2016). They were thrust southwestward during the latest Eocene over a basement of pre-Late Cretaceous arc-related terranes and their Upper Cretaceous–Eocene sedimentary cover (Aitchison et al., 1995; Cluzel et al., 2001).

Older basement elements formed in a Late Paleozoic to Mesozoic forearc setting, while younger elements represent parts of the Gondwana continental margin forearc and subduction complex (Maurizot, Cluzel, Meffre, et al., 2020). Overlying these rocks is a younger sedimentary cover (Maurizot, Bordenave, et al., 2020), including Upper Cretaceous to lower Eocene deposits that record continental break-up, rifting and thermal subsidence of northeastern Zealandia (Cluzel et al., 2011; Maurizot & Mortimer, 2020; Maurizot, Cluzel, Patriat, et al., 2020). Middle to upper Eocene sediments composed of turbidites and olistostromes document the evolution of a foreland basin before ophiolite obduction. A related subduction complex includes high-pressure metamorphic rocks such as an eclogite-facies mafic-ultramafic mélange (Pouebo Terrane) and a blueschist-facies metasedimentary unit (Diahot Terrane) (Clarke et al., 1997; Maurizot, Cluzel, Patriat, et al., 2020). These units incorporate subducted distal equivalents of younger sedimentary cover units, including the Formation à Charbon (coal-bearing Fm.) and Black Chert (Maurizot, Bordenave, et al., 2020).

A disrupted oceanic crustal unit, the Poya terrane, separates the basement and its cover from the overlying Peridotite Nappe (Cluzel et al., 2001, 2018). The Peridotite Nappe is ≤2,500 m thick and consists mainly of strongly depleted harzburgites with minor dunites (Pirard et al., 2013; Secchiari et al., 2020). Rare spinel and plagioclase lherzolites in the northern massifs (Tiébaghi, Poum) may be remnants of the original oceanic lithosphere (Cai et al., 2025; Secchiari et al., 2016). Podiform chromite, commonly associated with dunite, is also present (Maurizot, Sevin, et al., 2020). New Caledonian harzburgites are highly refractory and lack primary clinopyroxene. They contain 75%–85% olivine (Fo 0.9–92.9 mol.%), 15%–25% orthopyroxene (100*Mg/Mg + Fe²⁺ = Mg# of 89.8–94.2), and ≤1% chromium spinel (100*Cr/Cr + Al = Cr# of 0.44–0.71) (Secchiari et al., 2020). These rocks exhibit extremely low rare earth element (REE) concentrations (<0.1 chondritic values) with U-shaped profiles. Heavy REE compositions suggest an initial phase of dry fractional melting (15%), followed by remelting in a subduction zone (15%–18%) (Secchiari et al., 2016). Enrichment in Ba, Sr, Pb, Light–Medium REEs and Zr-Hf, along with Al₂O₃⁻ and CaO-poor orthopyroxene and low Na₂O, Al₂O₃, and TiO₂ clinopyroxene, indicate fluid-assisted melting and later metasomatism by subduction-related melts (Secchiari et al., 2020).

Isotopic data reveal complex fluid and melt interactions. Neodymium (eNdi 0.80–13.32) and strontium (⁸⁷Sr/⁸⁶Sr 0.70257–0.70770) isotopes suggest mantle depletion with enriched sediment-like contributions (Secchiari et al., 2020). Lead isotopes show a wide range, reflecting multiple sources. These geochemical signatures confirm the harzburgites as part of a forearc mantle wedge modified by subduction during the Eocene.
<figure><img src="image_3.png" /><figcaption>**Figure 1.** Geological map of New Caledonia in the South Pacific (redrawn and simplified from the online geological map of New Caledonia (DIMENCSGNC-BRGM, 2010)) showing the distribution of the Peridotite Nappe and locations of diamondiferous samples. Inset map of SW Pacific shows the relative locations of major crustal blocks including LHR (Lord Howe Rise), NR (Norfolk Ridge), LR (Loyalty Ridge), and HP (Hikurangi Plateau).</figcaption></figure>

Two generations of high-temperature minerals indicate superimposed melting and enrichment episodes (Secchiari et al., 2016, 2018, 2020). Early Mg-rich olivine and orthopyroxene porphyroclasts are preserved, while secondary neoblasts align within the rock's foliation (Pirard et al., 2013). Orthopyroxene cores yield Ca-in-orthopyroxene temperatures of 950–1150°C, with core-to rim variations reflecting localized cooling of ca. 50–100°C (Secchiari et al., 2020). Additional temperature estimates from Ca-in-olivine and olivine-spinel geothermometers indicate cooling to 800–965°C. Secondary orthopyroxene films yield slightly lower temperatures of 810–850°C. The pristine nature of these rocks suggests high-temperature processing without low-temperature overprints.

Above the harzburgites lies a 300–500 m thick dunitic transition zone (DTZ) capped by ultramafic-mafic cumulates. These cumulates, including gabbronorite, crystallized from forearc magmas intermediate between tholeiites and boninites (Secchiari et al., 2018). Eocene (55–50 Ma) intrusions range from ultramafic (pyroxenite) to felsic (granitoid) compositions with supra-subduction zone (SSZ) affinities, generated by slab melting and hydrous mantle melting. The magmatic activity ended with the intrusion of island-arc tholeiites (50–47 Ma) as the
forearc cooled (Cluzel et al., 2006, 2024). Slab melts injected into the fractured forearc along with melt/fluid circulation led to late-stage high-temperature peridotite metasomatism (Secchiari et al., 2020) and fluid-induced mantle wedge melting (Cluzel et al., 2016).

At lower temperatures, the Peridotite Nappe underwent serpentinization (lizardite ± chrysotile, magnetite) and crack-sealing precipitation (tremolite, antigorite, chlorite) driven by dehydration of the subducting slab (Cluzel et al., 2020). The timing and nature of these processes remain unclear. Subduction initiation, generating the New Caledonia ophiolite, began around 56 Ma, recorded by high-temperature amphibolite lenses at the nappe's base (Cluzel et al., 2012; Soret et al., 2016). Forearc magmatism and ophiolite formation ended by 50–47 Ma supported by magmatic zircon (46.84 ± 0.22 Ma) and rutile (47.13 ± 0.64 Ma) ages from Tiébaghi chromite ore (Patias et al., 2024). Finally, Peridotite Nappe emplacement, HP-LT complex exhumation, and isostatic re-equilibration were concluded by 34 Ma (Maurizot, Cluzel, Patriat, et al., 2020 and references therein).

**3. Sample Locations, Materials, and Processing**

To investigate the occurrence of diamonds and associated ultrahigh-pressure (UHP) and super-reduced (SuR) mineral phases akin to those reported from Tethyan and other Northern Hemisphere ophiolites (Yang et al., 2021 and references therein), two samples (~500 kg each) were collected from the New Caledonian Peridotite Nappe. One sample consisted of chromitite, and the other of fresh harzburgite.

The fresh harzburgite (unweathered and unserpentinized), representative of rocks found in the nearby Kopeto and Poya massifs (Secchiari et al., 2020) (Figure S2 in Supporting Information S1), was sourced from massive boulders in the Baraoua River bed. This river exits a canyon over 1 km deep, draining the southern flanks of the Me Maoya massif northwest of Bourail (GPS coordinates: 21.466378°S, 165.394475°E).

The chromitite sample was collected from the Tiébaghi Massif, which constitutes less than 2% of the Peridotite Nappe's surface area and once hosted one of the world's largest and richest chromium mines (Moutte, 1982). The chromium content of the ore is exceptionally high, ranging from 45 to 56 wt% Cr₂O₃ (Maurizot et al., 2023). Over a period spanning more than a century (1887–1992), chromite was intermittently extracted from podiform ore-bodies enclosed in a 10–50 m thick dunite envelope situated at the boundary between spinel herzolite and harzburgite. A chromite ore sample from the mine's deepest levels was gathered from spoil heaps at the decommissioned (1992) chromite mine's wash plant located approximately 15 km north of Koumac (GPS coordinates: 20.437636°S, 164.229450°E).

To ensure the integrity of the samples and prevent anthropogenic contamination (Ballhaus et al., 2021; Howell et al., 2015; Litasov et al., 2019), rigorous protocols were followed during mineral separation (methodology detailed in Supporting Information S1). The hypothesis of in situ diamond formation via isochoric shock heating from lightning strikes on mountaintop ophiolites (Ballhaus et al., 2017) was considered but deemed unlikely for both samples. The Tiébaghi chromitite sample was sourced from a depth of at least 300 m below the Tiébaghi mountain top, making lightning strikes improbable. Similarly, the peridotite boulders from Me Maoya, collected from the Baraoua River bed, exhibited no joints or fissures and had undergone extensive physical weathering and milling due to river transport and frequent tropical cyclone events.

Both samples were shipped to China for mineral separation at the Institute of Multipurpose Utilization of Mineral Resources (IMUMR), Chinese Academy of Geological Sciences (CAGS), Zhengzhou. The laboratory workflow, developed over the past two decades, is designed to eliminate any possibility of contamination and has been extensively validated (Aitchison et al., 2022). As a further check against contamination, several smaller ophiolitic samples from New Caledonia and New Zealand were separately processed using high-voltage electronic fragmentation (SelFrag) at the University of Queensland in Brisbane and at Macquarie University in Sydney. Following comminution, non-magnetic heavy minerals were concentrated using magnetic and electrostatic separators and a Wilfley table, followed by final magnetic separation with a Frantz magnetic separator and hand-picking of residues under a binocular microscope.

**4. Results**

The non-magnetic heavy minerals separated from the comminuted residues of the original bulk samples include zircon, rutile, titanite, apatite, sulfides, metal alloys and corundum. Detailed analyses of the zircons, rutiles and titanites are presented elsewhere (Aitchison et al., 2022; Patias et al., 2024).
<figure><img src="image_4.png" /><figcaption>**Figure 2.** Binocular microscope images of yellow micro-diamonds from the Peridotite Nappe of New Caledonia. (a) From Me Maoya harzburgite and (b) from Tiébaghi chromitite. (c and d) Scanning Electron Microscope (SEM) image of diamond extracted from Me Maoya harzburgite and associated Raman spectrum. Binocular microscope image of light blue to greenish moissanites from (e) Me Maoya harzburgite and (f) Tiébaghi chromitite. (g and h) SEM image of a moissanite grain from Tiébaghi chromitite sample and associated Raman spectrum.</figcaption></figure>

Rare pale yellow, translucent monocrystalline micro-diamonds were recovered from both samples. Eighteen micro-diamonds were recovered from the Me Maoya harzburgite (Figure 2a) and 14 from the Tiébaghi chromitite (Figure 2b). These micro-diamonds range in size from 100 to 250 µm. Scanning Electron Microscope imagery illustrates the octahedral to cubo-octahedral morphologies (Figure 2c).

Accompanying the diamonds are numerous moissanite grains, primarily light blue but also dark green and greenish blue in color. These grains exhibit hexagonal to rhombohedral shapes and range in size from 75 to 250 µm (Figures 2e–2g).

## 4.1. Diamonds and Moissanites

A Renishaw Raman microscope and spectrometer were used to identify mineral phases including diamond and moissanite (see SI File S2 for details). The characteristic Raman shift diamond peak at 1,332 cm⁻¹ is generated by vibrations of the two interpenetrating cubic sublattices (Frezzotti et al., 2014) and was distinctly observed in micro-diamonds from both the Me Maoya (Figure 2d) and Tiébaghi samples. Moissanites grains from both the Me Maoya harzburgite (Figure 2e) and Tiébaghi chromitite (Figures 2f and 2g) displayed Raman spectra with wavenumber shifts at 969, 788, 767 cm⁻¹, confirming their identification as moissanite (Nakashima & Harima, 1997) (Figure 2h).

## 4.2. Isotopic Composition

Isotopic compositions of the micro-diamond and moissanite grains were determined using a Sensitive High Resolution Ion Microprobe (SHRIMP) at Australian National University (ANU) (for full details of the SHRIMP methodology see SI File S2; for data see SI Datafile S1). The δ¹³C values (relative to Vienna Pee Dee belemnite (VPDB)) for micro-diamonds from both the harzburgite and chromitite samples consistently fall within a narrow highly negative range (Table 1). Micro-diamonds from the Me Maoya harzburgite display δ¹³C values ranging from -33.5 to -23.9‰ (n = 9, mean δ¹³C = -27.3‰, 1σ = 3.0), while those from the Tiébaghi chromitite range from δ¹³C -27.4 to -23.8‰ (n = 10, mean δ¹³C = -25.8‰, 1σ = 1.3). No zoning was observed, as consistent values were recorded across individual grains where multiple analyses were possible.
<caption>**Table 1**</caption>

δ¹³C<sub>VPDB</sub> Measurements for Diamonds and Moissanite Grains From Me Maoya Harzburgite and Tiébaghi Chromitite Samples From New Caledonian Peridotite Nappe

<table><thead><tr><th></th><th></th><th>Grain</th><th>δ¹³C</th><th>±</th></tr></thead><tbody><tr><td rowspan="8">Tiébaghi</td><td>Diamond</td><td>C1D-1.1</td><td>-26.95</td><td>0.14</td></tr><tr><td>Diamond</td><td>C1D-1.2</td><td>-27.16</td><td>0.10</td></tr><tr><td>Diamond</td><td>C1D-1.3</td><td>-27.41</td><td>0.13</td></tr><tr><td>Diamond</td><td>C1D-1.4</td><td>-27.09</td><td>0.08</td></tr><tr><td>Diamond</td><td>C1D-2.1</td><td>-23.77</td><td>0.04</td></tr><tr><td>Diamond</td><td>C1D-3.1</td><td>-25.94</td><td>0.09</td></tr><tr><td>Diamond</td><td>C1D-4.1</td><td>-25.40</td><td>0.03</td></tr><tr><td>Diamond</td><td>C1D-5.1</td><td>-25.67</td><td>0.07</td></tr><tr><td rowspan="8">Tiébaghi</td><td>Diamond</td><td>C1D-6.1</td><td>-23.78</td><td>0.10</td></tr><tr><td>Diamond</td><td>C1D-7.1</td><td>-24.99</td><td>0.11</td></tr><tr><td>Moissanite</td><td>C1M-1.1</td><td>-26.38</td><td>0.14</td></tr><tr><td>Moissanite</td><td>C1M-2.1</td><td>-27.14</td><td>0.32</td></tr><tr><td>Moissanite</td><td>C1M-3.1</td><td>-25.57</td><td>0.11</td></tr><tr><td>Moissanite</td><td>C1M-4.1</td><td>-26.42</td><td>0.24</td></tr><tr><td>Moissanite</td><td>C1M-5.1</td><td>-26.65</td><td>0.17</td></tr><tr><td>Moissanite</td><td>C1M-6.1</td><td>-27.98</td><td>0.27</td></tr><tr><td rowspan="8">Me Maoya</td><td>Moissanite</td><td>C1M-7.1</td><td>-27.11</td><td>0.13</td></tr><tr><td>Diamond</td><td>H1D-1.1</td><td>-23.85</td><td>0.04</td></tr><tr><td>Diamond</td><td>H1D-2.1</td><td>-27.71</td><td>0.11</td></tr><tr><td>Diamond</td><td>H1D-3.1</td><td>-33.51</td><td>0.11</td></tr><tr><td>Diamond</td><td>H1D-3.2</td><td>-27.44</td><td>0.11</td></tr><tr><td>Diamond</td><td>H1D-3.2</td><td>-31.35</td><td>0.15</td></tr><tr><td>Diamond</td><td>H1D-5.1</td><td>-25.33</td><td>0.05</td></tr><tr><td>Diamond</td><td>H1D-6.1</td><td>-24.47</td><td>0.06</td></tr><tr><td rowspan="8">Me Maoya</td><td>Diamond</td><td>H1D-7.1</td><td>-26.76</td><td>0.05</td></tr><tr><td>Diamond</td><td>H1D-8.1</td><td>-25.46</td><td>0.06</td></tr><tr><td>Moissanite</td><td>H1M-1.1</td><td>-26.46</td><td>0.11</td></tr><tr><td>Moissanite</td><td>H1M-2.1</td><td>-26.25</td><td>0.09</td></tr><tr><td>Moissanite</td><td>H1M-3.1</td><td>-26.68</td><td>0.21</td></tr><tr><td>Moissanite</td><td>H1M-4.1</td><td>-26.20</td><td>0.10</td></tr><tr><td>Moissanite</td><td>H1M-5.1</td><td>-31.77</td><td>0.09</td></tr><tr><td>Moissanite</td><td>H1M-6.1</td><td>-25.97</td><td>0.20</td></tr><tr><td rowspan="4">Moissanite</td><td>Moissanite</td><td>H1M-7.1</td><td>-28.03</td><td>0.16</td></tr><tr><td>Moissanite</td><td>H1M-8.1</td><td>-25.94</td><td>0.08</td></tr><tr><td>Moissanite</td><td>H1M-10.1</td><td>-25.67</td><td>0.14</td></tr></tbody></table>

Moissanite grains exhibit similarly restricted δ¹³C values. In the Me Maoya, harzburgite values range from -31.8 to -25.7‰ (n = 9, mean δ¹³C = -27.0‰, 1σ = 1.8), while in the Tiébaghi chromitite values range from -28.0 to -25.6‰ (n = 7, mean δ¹³C = -26.7‰, 1σ = 0.7).

## 5. Discussion

The occurrence of rare micro-diamonds in ophiolites contrasts with the conditions associated with those of more common kimberlitic and peridotitic diamonds, which are typically formed at considerable depths (150—>>250 km) in the subcontinental asthenosphere or mantle transition zone (MTZ) (Cartigny, 2005; Shirey et al., 2013; Stachel & Luth, 2015; Yang et al., 2021). The characteristics of the New Caledonian micro-diamonds align with similar occurrences reported in other ophiolitic rocks from the Northern Hemisphere (Yang et al., 2021). The narrow range of light δ¹³C in diamonds (n = 27 mean = -28.2 ± 3.9) and associated SiC, along with the co-occurrence of other SuR-UHP minerals amongst SSZ ophiolites world-wide (Yang et al., 2014), suggests a consistent geologic process rather than a random occurrence. Given their global distribution, it is pertinent to explore whether ophiolitic diamonds represent a unique formation pathway distinct from that of diamonds from deeper origins.

The diamonds in this study were recovered from residues of large rock samples, with total diamond yields from each ~500 kg sample amounting to ≪1 carat (i.e. ≪0.2 g). The combined yield of zircons, rutiles, diamonds and moissanites totaled only a few grams, representing approximately 4.0 × 10⁻⁷ of the total sample mass. This yield is consistent with reports from Northern Hemisphere ophiolite occurrences (Yang et al., 2014). Although in situ ophiolitic diamonds from New Caledonia have not been observed, they have been reported from Ray Iz in the Urals (see Figure 2 of Yang et al. (2014)) and for moissanite within chromian spinel from Luobusa (Liang et al., 2014). Other ophiolites, such as the Tehuitzingo serpentinite in southern Mexico (Farré-de-Pablo et al., 2019), Moa-Baracoa Ophiolitic Massif in eastern Cuba (Pujol-Solà et al., 2020) and Nidar ophiolite in Ladakh, India (Das et al., 2017), contain in situ nano-diamonds often associated with serpentinite, but differ from the New Caledonian micro-diamonds in fresh harzburgite from Me Maoya.

Deep subduction of low-pressure rocks to the MTZ or lower mantle followed by rapid exhumation due to slab-rollback in a collisional zone is the prevailing explanation for ophiolitic diamond formation (McGowan et al., 2015; Yang et al., 2021). However, this model fails to account for the consistently light carbon isotopic composition of ophiolitic diamonds in contrast to the broad range of δ¹³C values typically seen in diamonds from deeper origins. Zhou et al. (1996) proposed diamond formation during chromite crystallization from boninitic melts reacting with older MORB-type peridotites during SSZ ophiolite formation in an intra-oceanic island arc setting. To explain the

presence of SuR and UHP minerals, they invoked a slab tearing or break-off, allowing the asthenospheric mantle to intrude (Robinson et al., 2015; Zhou et al., 2014).

An association with a deep-seated mantle plume, whereby UHP and SuR minerals formed in the MTZ and mixed with deeply subducted surface material are returned to shallow mantle levels, has also been considered (Doucet et al., 2021; Robinson et al., 2004). However, both the above models are not consistent with spatial and temporal distributions among ophiolitic diamonds. The Tethyan ophiolites developed above lithospheric slab(s) subducting at a high angle rather than parallel to the associated plate boundary, which, to be affected by slab tear or a mantle plume along its linear entirety, would have had to track coincident with any such feature.
The ubiquitous isotopically light δ¹³C compositions are a defining feature that any model must address. New Caledonian micro-diamonds (n = 19, mean δ¹³C = -26.5‰, 1σ = 1.8, range -33.5 to -23.77‰; Figure 3) and other ophiolitic diamonds (n = 189, mean δ¹³C = -25.5‰, 1σ = 1.6) display a narrow, normally distributed range of δ¹³C values unique to those from SSZ ophiolites (Lian et al., 2018). In contrast, peridotitic (P-type) diamonds also exhibit a narrow range but at heavier δ¹³C values (n = 1,906, median = -4.9‰, mode δ¹³C = -5.2‰, 1σ = 1.9 ‰, range -34.5-2.3‰) (Cartigny et al., 2014; Stachel et al., 2022). Although some eclogitic (E-type) diamonds exhibit light isotope fractionation, they show a wider range skewed toward heavier mantle values (n = 1,204, mode = -5.0‰, mean δ¹³C = -6.5‰, 1σ = 6.7‰, range -40.7-2.5‰) (Cartigny et al., 2014; Stachel et al., 2022). These differences suggest that MTZ-origin models followed by recycling into the upper mantle are unlikely for ophiolitic micro-diamonds. Published δ¹³C ranges for synthetic diamonds vary widely and do not match the tight isotopic range of ophiolitic diamonds (H. Wang et al., 2018; Xu et al., 2018), ruling out contamination by synthetic materials.

Moisanites from the New Caledonian ophiolite samples also exhibit light δ¹³C compositions (n = 16, mean = -26.9‰, 1σ = 1.4‰) comparable to moisanites from other ophiolites (n = 25, mean δ¹³C = -28.4‰) (e.g., Tibet, Oman, Urals - Trumbull et al., 2009). The known range of carbon isotopic compositions for natural moisanites is light (δ¹³C = -20 to -37‰) and is consistent with fluids originating from metamorphosed organic carbon (Mathez et al., 1995; Schmidt et al., 2014; Trumbull et al., 2009).

The isotopic distinctiveness of ophiolitic micro-diamonds suggests a formation mechanism related to their occurrence in subduction zones. Subduction of sediment rich in organic carbon is a likely pathway (Clift, 2017; Galvez et al., 2020; Kelemen & Manning, 2015; Müller et al., 2022). Fractionation of carbon isotopes during photosynthesis in C3 plants, dominant in pre-Oligocene times, produces isotopically light δ¹³C values that mirror those of ophiolitic micro-diamonds (O'Leary, 1988; Kohn, 2010). In New Caledonia, terrestrial to marginal marine sediments from the Formation à Charbon, subducted during the Eocene, present a potential source of C3 carbon (Garrouste et al., 2021; Maurizot, Bordenave, et al., 2020). The ages of zircon and rutile xenocrysts in our samples support this origin (Aitchison et al., 2022; Patias et al., 2024).

Subduction-related processes often lead to significant CO₂ loss beneath the forearc as a result of devolatilization in the subducting slab and its sedimentary cover, yet initial recycling of sedimentary carbon begins early at depths of <50–150 km (Tumiati et al., 2022). Light carbon isotope enrichment in forearc regions, as observed in Costa Rica (Barry et al., 2019) and the Marianas forearc (Eickenbusch et al., 2019), suggests preferential mobilization of isotopically light biogenic carbon during early subduction.

The abundance of light δ¹³C moissanite indicates redox conditions controlled by carbon-buffering (Shiryaev & Gaillard, 2014). Natural moissanite forms under super-reducing C-O-H-fluid conditions at extremely low fO₂, possibly below the iron-wustite (IW) buffer (Golubkova et al., 2016; Schmidt et al., 2014). The presence of awaruite (FeNi₃) in New Caledonia SSZ peridotites suggests equilibrium with hydrogen-bearing fluids below FMQ-3 (fayalite-magnetite-quartz) in the SSZ mantle (Evans & Frost, 2020; Evans et al., 2017; Frost, 1985).

Given that redox conditions prevailing in the mantle wedge seem incompatible with ophiolitic micro-diamond and moissanite generation, we consider whether appropriate redox conditions might instead occur on the subducting slab. Global subducted sediment averages (GLOSS and GLOSS II) (Plank, 2014) are widely considered as a proxy of starting conditions in considerations of subduction zone redox conditions (Cannaò & Malaspina, 2018; Evans, 2012). Such averages are not generally indicative of highly reducing conditions. However, Plank (2014) specifically comments that there is a major gap in characterization of the subduction budget in response to volatile species including N and organic carbon components. In particular, the mass specific redox capacities (δO₂) for lignite (>>50,000 µmol O₂/g), euxinic lake sediments, and black shales are many orders of magnitude greater than those for dunite (circa 175 µmol O₂/g) (Galvez & Jaccard, 2021). Extremely low fO₂ values associated with the high mass-specific redox capacity of subducting terrestrial organic matter would not only favor the formation of moissanite (Golubkova et al., 2016; Schmidt et al., 2014) but also that of diamonds. In this regard, it is relevant to recall that the black shales of the Diahot terrane subducted beneath the New Caledonia forearc mantle contain C3-rich sediments of appropriate isotopic composition. We posit here that these sediments hosted the carbon for the redox-buffering necessary for fO₂ values approaching IW-6 under which moissanite might form. Similar conditions have been reported for moissanite formation in Triassic amphibolite-facies metasedimentary rocks of Bulgaria associated with temperatures of at least 500–580°C (Machev et al., 2018). Interestingly, the reported metamorphic grades are consistent with the maximum temperature conditions
<figure><img src="image_5.png" /><figcaption>**Figure 3.** Histograms of relative abundances of δ¹³C values for diamonds and moissanites from Tiébaghi chromitite and Me Maoya harzburgite samples (this study) compared with δ¹³C values from diamonds and moissanites of known paragenesis worldwide (Cartigny et al., 2014; Trumbull et al., 2009).</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>**Figure 4.** Possible reaction pathway for subducting organic carbon-rich sediments (after Galvez et al. (2020)).</figcaption></figure>

experienced by xenocrystic rutile and zircon (U—Pb closure temperatures of 620 ± 20°C and 900°C respectively) within our samples (Aitchison et al., 2022; Patias et al., 2024). We note here that no micro-diamonds are known from heavy mineral residues of lithologies fingerprinted as potential sources of these zircons and rutiles (Adams et al., 2009; Aitchison et al., 1998; Cluzel et al., 2011). Furthermore, the unique isotopically light carbon signature of the ophiolitic diamonds compared to that of most other more common types of diamond strongly mitigates against the recycling of any pre-existing diamonds.

Localized super-reducing conditions in the sedimentary carapace atop the subducting slab could facilitate diamond and moissanite formation. Such conditions (Griffin et al., 2018) could support the spontaneous nucleation and precipitation of diamonds and moissanites in microenvironments within the subduction zone (Frezzotti et al., 2014).

Profound changes in the nature of fluids associated with subduction zones occur at depths associated with the maximum temperatures experienced by rutile xenocrysts in our samples (Manning & Frezzotti, 2020). Thermodynamic modeling demonstrates that the duration and rate of plate subduction together with the age of the subducting lithosphere all influence the temperature gradient that affects the devolatilizing slab top (Holt & Condit, 2021; Maunder et al., 2019). Organic sediments such as kerogen in the subduction channel can serve as CH₄ and H₂-bearing source rocks essential for diamond formation (Suzuki et al., 2024). Having formed in highly reducing shales from early diagenetic transformation of abundant biological organic carbon, type 1 kerogen experiences thermal transformation to type 2 kerogen followed by transformation to graphite/diamond (Figure 4). By association, such conditions are likely not simply coincidental to the genesis of ophiolitic diamonds.

## 6. A New Caledonian Model for the Formation of Ophiolitic Micro-Diamonds

The complex multi-collisional histories of Northern Hemisphere SSZ ophiolites, such as the Yarlung Tsangpo suture zone in Tibet (Aitchison et al., 2000), have hindered straightforward interpretations of micro-diamond formation. In contrast, the Eocene tectonic evolution of the New Caledonia subduction zone is well documented and remains largely unaltered by later tectonic events, providing clearer insights into the formation of ophiolitic micro-diamonds (Cluzel et al., 2018; Maurizot & Mortimer, 2020).

The ophiolitic mantle forearc rocks of New Caledonia, from which our samples were collected, were formed during intra-oceanic subduction and emplaced onto the edge of a Gondwanan continental fragment as it was drawn into the subduction zone (Aitchison et al., 1995; Cluzel et al., 2001; Maurizot, Cluzel, Patriat, et al., 2020). The presence of supra-subduction dikes intruding the New Caledonia ophiolite confirms a forearc setting for the diamond-hosting lithologies (Cluzel et al., 2016, 2024). Moreover, the fresh unserpentinized nature of the Me Maoya harzburgite sample constrains the timing of diamond formation before widespread serpentinization of SSZ mantle rocks. Temperature constraints for minerals in both the Tiébaghi and Me Maoya samples are provided by undisturbed xenocrystic Cretaceous and older rutiles (Aitchison et al., 2022), which must not have exceeded the U-Pb closure temperature of 620 ± 20°C (Kooijman et al., 2010).
<figure><img src="image_7.png" /><figcaption>**Figure 5.** Model of the geodynamic setting in which diamonds may have formed from organic carbon-rich, highly reducing sediments that were subducted beneath the supra-subduction zone forearc ophiolite of New Caledonia.</figcaption></figure>

During the Eocene, suitable carbon-rich source sediments were subducted beneath the New Caledonian forearc wedge (Maurizot, Cluzel, Patriat, et al., 2020). The age spectra of inherited zircon (90–2,700 Ma) and rutile (125–500 Ma) populations in both samples closely match those of potential source lithologies on the subducting slab. Alban-age xenocrystic zircons (113.2–100.5 Ma) are consistent with derivation from Upper Cretaceous sediments on the slab (Aitchison et al., 2022; Patias et al., 2024). Terrigenous plant material, abundant in the Diahot black shales (metamorphosed lateral equivalents of the Formation à Charbon), represents a potential light δ¹³C source essential to explaining the origins of New Caledonian ophiolitic diamonds (Cluzel et al., 2011; Garrouste et al., 2021). The reducing capacity of these sediments far exceeds that of mantle lithologies, providing the conditions necessary for diamond formation (Galvez & Jaccard, 2021).

The thermal structure of intra-oceanic subduction zones is such that pressures necessary to cross the graphite-diamond phase transition (Day, 2012) can be exceeded before a subducting lithospheric slab reaches 700°C (Gerya, 2011; Syracuse et al., 2010). Thus, diamond formation in New Caledonian likely occurred under localized super-reducing conditions at appropriate pressures and temperatures within the organic sediment veneer on the subducting slab (Figure 5). The subduction channel beneath the New Caledonian forearc facilitated introduction of anoxic carbon-rich sediments to depths conducive to the growth of diamond, moissanite and other SuR UHP
minerals. These minerals, along with pre-existing refractory zircons and rutiles, were introduced into the forearc mantle via solid-state mass transfer before rapid uplift and emplacement with the Peridotite Nappe.

While this model explains the New Caledonia ophiolitic micro-diamonds, its broader applicability warrants further consideration. Samples from many other ophiolitic micro-diamonds occurrences also contain xenocrystic zircon and rutile populations. However, in Tibetan Tethyan examples, while mid-Cretaceous ophiolite formation occurred in an intra-oceanic SSZ setting, Eocene ocean closure occurred much later, making coeval subduction of terrestrial-derived sedimentary units like the Formation à Charbon unlikely.

An alternative source of super-reducing sediments could be anoxic sediments from the thin deep marine sedimentary succession on the Neotethyan Ocean floor (Bomou et al., 2013; Wan et al., 2003; T. Wang et al., 2019). These sediments are associated with negative Carbon Isotope Excursions (CIEs) during Oceanic Anoxic Events (OAEs) such as the Turonian Jenkyns event (Fernández-Martínez et al., 2023; Jenkyns, 2010; Zhang et al., 2016). If subducted to sufficient depth, localized redox-buffering conducive to the formation of moissanite and micro-diamonds may also have occurred in these settings.

## 7. Conclusions

The discovery of micro-diamonds in the forearc rocks of the Eocene New Caledonian ophiolite expands the known distribution of these minerals, suggesting that the presence of micro-diamonds in SSZ ophiolites could be a general global phenomenon. In light of mantle heterogeneity (de Sampaio et al., 2024), it supports interpretation of their formation in relation to an SSZ process rather than a deep mantle source. This finding enhances our understanding of the recycling of biologically fractionated carbon in forearc settings (Barry et al., 2019; Galvez et al., 2020).

Traditionally, natural diamond formation is associated with well-defined pressure and temperature conditions (Cartigny, 2005; Shirey et al., 2013; Smit et al., 2022; Stachel & Luth, 2015). However, the unique characteristics of ophiolitic micro-diamonds and their geological context challenge these conventional models, indicating that alternative mechanisms such as formation amongst highly reducing carbon-rich sediments atop the subducting slab with later incorporation into the over-riding mantle wedge must be considered. Furthermore, the coexistence of moissanite in forearc mantle settings, where redox conditions are typically not suited to its formation (Golubkova et al., 2016), adds complexity to our understanding of these occurrences.

We propose that subduction of highly reducing organic carbon-rich anoxic sediments to depths conducive to the diagenetic transformation of organic carbon into kerogen followed by thermal transformation into graphite and micro-diamonds may explain these findings. Subsequent solid-state transfer of these micro-diamonds into the overriding forearc mantle provides a plausible, though unverified, pathway for the formation of ophiolitic micro-diamonds. The mechanism of this solid-phase transfer from the subducting slab to the SSZ mantle remains poorly understood, but Xiong et al. (2025) speculate that it must be fluid-assisted.

While this hypothesis remains speculative, it represents a significant step toward resolving key questions surrounding micro-diamond formation in SSZ ophiolites. It also has broader implications for understanding how subducted sediments influence oxygen fugacity within subduction zones. Although this process may be more common than previously thought, the volumetric transfer of material from the subducting slab to the mantle wedge appears to be very limited.

## Conflict of Interest

The authors declare no conflicts of interest relevant to this study.

## Data Availability Statement

All data sets generated during this study have been archived in Zenodo via Aitchison (2025).

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## Acknowledgments

This article is a contribution to the Australian Research Council-funded Discovery Project DP190100814 "Diamonds in Ophiolites: Recycling deep mantle into supra-subduction zones" and to the IGCP 649 Project "Diamonds and recycled mantle." D.C. and J.C.A. thank M. Henri Reuillard, M. Marc Akaro and the Tiébaghi Historical Society for arranging access to the mine wash-plant and to Aileen Davis, Elisabeth Mancel, Martine and Vincent Mardhel for assistance in sample collection. We thank the G-Cubed Editor David Hernández Uribe and anonymous reviewers for their constructive comments on earlier versions of this manuscript. Open access publishing facilitated by The University of Queensland, as part of the Wiley - The University of Queensland agreement via the Council of Australian University Librarians.
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