# Forms and fluxes of carbon: Surface to deep

Alberto Vitale Brovarone, Kevin Wong, Donato Giovannelli, Benoit de Pins,

Fabrice Gaillard, Malcolm Massuyeau, Fabrizio Nestola, Martha Giovanna

Pamato, Isabelle Daniel

<figure><img src="image_1.png" /></figure>

## To cite this version:

Alberto Vitale Brovarone, Kevin Wong, Donato Giovannelli, Benoit de Pins, Fabrice Gaillard, et al. "Forms and fluxes of carbon: Surface to deep. Treatise on Geochemistry, 2, Elsevier, pp.647-698, 2025, 10.1016/B978-0-323-99762-1.00142-X. insu-04742757

HAL Id: insu-04742757

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Submitted on 18 Oct 2024

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<figure><img src="image_2.png" /></figure>
Forms and fluxes of carbon: Surface to deep

Alberto Vitale Brovarone<sup>a,b,c</sup>, Kevin Wong<sup>a</sup>, Donato Giovannelli<sup>d,e,f,g,h</sup>, Benoit de Pins<sup>d</sup>, Fabrice Gaillard<sup>i</sup>, Malcolm Massuyeau<sup>j</sup>, Fabrizio Nestola<sup>k</sup>, Martha Giovanna Pamato<sup>k</sup>, and Isabelle Daniel<sup>i</sup>,<sup>a</sup> Dipartimento di Scienze Biologiche, Geologiche e Ambientali (BiGeA), Alma Mater Studiorum Università di Bologna, Bologna, Italy;<sup>b</sup> Institut de Minéralogie, de Physique des Matériaux et de Cosmochimie (IMPMC), Sorbonne Université, Muséum National d'Histoire Naturelle, UMR CNRS 7590, IRD UR206, Paris, France;<sup>c</sup> Institute of Geosciences and Earth Resources, National Research Council of Italy, Pisa, Italy;<sup>d</sup> Department of Biology, University of Naples Federico II, Naples, Italy;<sup>e</sup> National Research Council—Institute of Marine Biological Resources and Biotechnologies (CNR-IRBIM), Ancona, Italy;<sup>f</sup> Department of Marine and Coastal Science, Rutgers University, New Brunswick, NJ, United States;<sup>g</sup> Marine Chemistry, Geochemistry Department—Woods Hole Oceanographic Institution, Falmouth, MA, United States;<sup>h</sup> Earth-Life Science Institute, Tokyo Institute of Technology, Tokyo, Japan;<sup>i</sup> CNRS, BRGM, ISTO, UMR 7327, Université d'Orléans, Institut des Sciences de la Terre (ISTO), Orléans, France;<sup>j</sup> Institut für Mineralogie, Westfälische Wilhelms-Universität Münster, Münster, Germany;<sup>k</sup> Department of Geosciences, University of Padova, Padova, Italy;<sup>l</sup> Universite Claude Bernard Lyon1, LGL-TPE, UMR 5276, CNRS, Ens de Lyon, Universite Jean Monnet Saint-Etienne, Villeurbanne, France

© 2024 Elsevier Inc. All rights are reserved, including those for text and data mining, AI training, and similar technologies.

<table><tr><td>Introduction</td><td>2</td></tr><tr><td>Origins of deep carbon</td><td>3</td></tr><tr><td>Carbon at the interface between the geosphere and biosphere</td><td>5</td></tr><tr><td>Carbon's central role in life</td><td>6</td></tr><tr><td>The deep subsurface biosphere</td><td>6</td></tr><tr><td>Impact of life and organic carbon on the deep carbon cycle</td><td>7</td></tr><tr><td>Biogenic carbon: From the surface to the subsurface of the planet</td><td>7</td></tr><tr><td>Effect of life on carbon fluxes to and from the deep Earth</td><td>8</td></tr><tr><td>Deep carbon forms</td><td>9</td></tr><tr><td>Carbon in the Earth's core</td><td>9</td></tr><tr><td>Carbon at the interface between the core and the lowermost mantle</td><td>11</td></tr><tr><td>Carbon throughout the Earth's mantle</td><td>12</td></tr><tr><td>Reduced lower mantle and transition zone</td><td>13</td></tr><tr><td>The carbon-depleted oxidized upper mantle</td><td>14</td></tr><tr><td>Sub-continental lithospheric mantle</td><td>14</td></tr><tr><td>Sub-oceanic lithospheric mantle</td><td>16</td></tr><tr><td>Deep carbon in the crust</td><td>16</td></tr><tr><td>Continental crust</td><td>16</td></tr><tr><td>Oceanic crust and upper lithospheric mantle</td><td>18</td></tr><tr><td>Deep abiotic organic matter</td><td>19</td></tr><tr><td>Carbon movement within and between deep reservoirs</td><td>21</td></tr><tr><td>Movements of carbon from the crust to the deep Earth</td><td>21</td></tr><tr><td>Subduction of crustal carbon</td><td>21</td></tr><tr><td>Subduction erosion</td><td>23</td></tr><tr><td>Movements of carbon from depth to the crust</td><td>23</td></tr><tr><td>Tectonic movements of solid carbon from the deep Earth to the surface</td><td>23</td></tr><tr><td>Metamorphic and magmatic processes of carbon mobilization</td><td>24</td></tr><tr><td>Carbon reprecipitation at depth</td><td>28</td></tr><tr><td>Movements of deep carbon in melts</td><td>28</td></tr><tr><td>Carbon in intraplate magmatism</td><td>28</td></tr><tr><td>Carbon in arc basalts</td><td>30</td></tr><tr><td>Diamonds as clues on deep and super-deep carbon movements</td><td>31</td></tr><tr><td>Diamonds over (geological) time</td><td>31</td></tr><tr><td>Fluxes of deep carbon in diamonds</td><td>32</td></tr><tr><td>Diamonds and their solid carbon inclusions</td><td>33</td></tr><tr><td>Diamonds as trackers of global carbon recycling and geobiological evolution</td><td>33</td></tr><tr><td>Secular variation of surficial versus deep carbon reservoirs or steady state?</td><td>34</td></tr><tr><td>Concluding remarks</td><td>35</td></tr><tr><td>Acknowledgments</td><td>36</td></tr><tr><td>References</td><td>36</td></tr></table>
**Abstract**

Carbon is an essential element for the coevolution of Earth and life, and its largest fraction is stored in the deep Earth. The availability of carbon at the Earth's surface or shallow subsurface over the past 3.8 billion years has been modulated by this reservoir of deep carbon, which has played a fundamental role in the emergence and diversification of life. In turn, major geobiological changes at the Earth's surface have profoundly affected the chemical fingerprints and morphological forms of carbon entering the deep Earth through subduction. This chapter presents an overview of the geology of deep carbon, from its origins and forms throughout the Earth's history, to its movements and fluxes between shallow and deep reservoirs.

**Keywords**

Abiotic organics; Biotic and abiotic carbon; Carbon budgets; Carbon cycle; Carbon fluxes; Carbon in melts; Deep carbon; Coevolution of Earth and life; Origin of carbon; Metamorphic carbon degassing

**Key points**

* We present the origins and forms of deep carbon throughout the Earth's history

* We summarize the reservoirs of deep carbon from the core to the crust

* We review the available estimates of carbon fluxes from the surface to the deep Earth and vice-versa

**Introduction**

Carbon (C) is an essential element for life and the functioning of our planet. Its presence and quantity in the biosphere and atmosphere has largely been controlled by exchange with the large carbon reservoir in the deep Earth throughout the Earth's history (Berner, 1994; Hazen and Schiffries, 2013). Deep carbon is recycled through tectonic (or mechanical), metamorphic, magmatic, and fluid-rock interaction processes, and — within the parameter space for life — through biologically-mediated processes. However, the largest fraction of Earth's carbon is stored in deep Earth reservoirs such as the mantle and the core (Dasgupta and Hirschmann, 2010). Because of that, the distribution, movements, and fluxes of deep carbon within and between geological reservoirs are still largely unconstrained and model-dependent (Dasgupta, 2013; Wood et al., 2013). The long-term carbon cycle on our planet is primarily controlled by plate tectonics, with convergent plate margins and subduction driving the input of surface carbon into the deep Earth and volcanism and diffuse degassing acting as the main pathway for the return of carbon to the Earth's surface (Halama and Bebout, 2021; Jarrard, 2003; Kelemen and Manning, 2015; Lee et al., 2019; Plank and Manning, 2019; Werner et al., 2019).

Defining the boundary between surface and deep carbon is not straightforward, as the two reservoirs are intimately connected. Deep carbon represents the largest fraction of carbon on Earth and controls its long-term, or geological, cycle (~$10^{23}$ mol; Lee et al., 2019). On the other hand, surface carbon represents a much smaller reservoir ($10^{21}$–$10^{22}$ mol; Hayes and Waldbauer, 2006; Javoy et al., 1982), but is central for — and profoundly affected by — biological activity with important implications on the types, sizes and chemical fingerprints of both solid carbon being transferred from the surface to the deep Earth at subduction zones, and deep carbon degassing into the atmosphere. The emergence and diversification of life over the Earth's history has substantially affected the diversity, distribution, and signatures of carbon-bearing geological materials (Hazen et al., 2019), thereby affecting deep carbon processes and chemical fingerprints. This includes subsurface life that acts on deep carbon to alter its composition.

Given the diverse interconnections among processes acting at various depths within the crust and mantle, and the complex interplay between geological and biological processes, the potential solution to identify a boundary between surface and deep carbon can be found in the respective sizes and response times of carbon in the various reservoirs. Lee et al. (2019) defined the limit between endogenic (deep; Earth's lithosphere and mantle) and exogenic (surface; oceans, atmosphere, biosphere, reactive marine sediments) carbon at $10^7$ Gt mass of carbon (where $1 \text{ Gt} = 10^{15} \text{ g}$) and 10,000–100,000 years of residence time, with endogenic carbon being richer in carbon and slower in response time relative to exogenic carbon. In the lithosphere and mantle reservoir, the residence time of carbon is estimated to be in the range of 1 million to 1 billion years, in comparison to the <100,000 year timescales typical of the exogenic carbon cycles (Lee et al., 2019).

Starting from the origin of deep carbon on Earth, this chapter presents a summary of the most common forms of deep carbon, their movements — from solids to fluid phases — and fluxes. Although carbon in the biosphere is considered exogenic (response time <10,000 years; Lee et al., 2019), this chapter also presents some key aspects of carbon in the biosphere. This is because of the central role of life in controlling the amounts, forms, and chemical signature of carbon not only at the Earth's surface, but also in the subsurface where endogenic and exogenic processes together act to modulate the fluxes of carbon between the surface and the deep Earth.
## Origins of deep carbon

Isotopic constraints on Earth's bulk carbon are consistent with a chondritic origin (Broadley et al., 2022). The carbon content of non-carbonaceous enstatite chondrites and carbonaceous chondrites could explain the concentration of carbon and the ¹³C/¹²C ratio of the bulk silicate Earth (BSE; Hirschmann, 2018; Marty et al., 2016). A similar chondritic provenance of most life-forming elements, i.e., hydrogen and nitrogen, is also consistent with the elemental and isotopic abundances for these elements (Broadley et al., 2022). It therefore seems likely that carbon and other life-forming elements were delivered to Earth during planetary accretion. The geological history of carbon thus started with the earliest planetary events that have shaped planet Earth (Fig. 1).

Carbon was first delivered by chondrite-like materials on small planetary bodies that formed within the first 1–3 millions of years of the solar system, via a mechanism termed ‘pebble accretion’ (Johansen and Lambrechts, 2017; Li et al., 2016). During this early stage, radioactive heating was intense enough to induce melting of these bodies, enabling magma ocean and core differentiation. Upon this early differentiation, carbon was greatly partitioned into the core (Li et al., 2016), but due to the combination of low gravity and high temperature, outgassing and loss of gaseous carbon was certainly prevailing. This implies that the carbon content of these small planetary bodies was low. It was only when the size of these planet-forming bodies reached diameters of ~1000 km that gravity was strong enough to limit gas escape. Carbon incorporation in the planet interior then became considerable, with the particularity that under the prevailing reducing conditions, a significant amount of carbon must have been dissolved in the forming core (Li et al., 2016; see Section “Carbon in the Earth’s core”) and outgassed into the atmosphere (Gaillard et al., 2022a). The proportion of carbon being dissolved in the magma ocean (the Hadean BSE) was in comparison small (Gaillard et al., 2022a,b).

Once pebble accretion ended, planetary growth occurred via collisions, a highly energetic process capable of partly or totally ejecting the atmosphere into space (Chen and Jacobson, 2022). Since the magma ocean atmosphere was primarily composed of carbon species, collisions mostly caused carbon loss. Secondary atmospheres were subsequently formed by the degassing of the planetary interior after these impacts. The primordial carbon delivered by chondrites has therefore been mostly lost to space or partitioned into the core, which, in both cases, constitutes an irreversible mechanism, subtracting some carbon from the BSE (Gaillard et al., 2021). Accordingly, non-carbonaceous (>3000 ppm C) and carbonaceous (several wt% C) chondrites actually contain more carbon than BSE (Gaillard et al., 2021), which has a carbon concentration ranging from 140 ppm (Hirschmann, 2018) to >500 ppm (Marty et al., 2013, 2016; see also Section “Carbon throughout the Earth’s mantle”). It has also been hypothesized that large impacts, such as that forming the Moon, could have contributed a substantial amount of carbon and other volatiles to the BSE, as opposed to facilitating overall carbon loss (Grewal et al., 2019).

<figure><img src="image_3.png" /><figcaption>**Fig. 1** The geological history of deep carbon summarized in a clockwise time line sketch – starting bottom left – (see Section “Origins of deep carbon” for details). It shows the evolution of deep carbon from the early accretion of planet Earth, where magma ocean processes prevailed and drove core-mantle-atmosphere exchanges, to the present-day solid-state convection regime wherein mantle-crust-atmosphere exchanges are dominated by plate tectonics.</figcaption></figure>
On large bodies such as the Earth, a thick magma ocean could develop with high pressures prevailing at its base. When pressure exceeds 15–20 GPa, ferric-iron-rich silicate melts (with Fe³⁺/Feₜot ~ 0.1–0.2) are stable in equilibrium with molten Fe-rich alloys (Armstrong et al., 2019). At lower pressure, under otherwise similar conditions, such melts would contain only ~5% ferric iron. Vigorous convection from the base to the top of the magma ocean implies that such ferric-iron-rich silicate melts would buffer a high oxygen fugacity when exposed to surficial conditions. This implicitly assumes that the Fe-rich molten alloys remain at the base of the magma ocean due to negative buoyancy forces. Consequently, while early and thin magma oceans outgassed mostly CO₂, late and deep magma oceans outgassed mostly CO₂. The early redistribution of oxygen from CO₂ at shallow conditions to Fe₂O₃ (i.e., Fe³⁺) in the deep magma ocean may have caused the first diamond precipitation in the interior of Earth as illustrated in Fig. 1 (Armstrong et al., 2019).

The solidification of the magma ocean is a complicated stage, blending thermomechanical and thermochemical processes (Ballmer et al., 2017; Hier-Majumder and Hirschmann, 2017; Korenaga, 2021). It is yet unclear how magma ocean carbon could behave during this solidification. If a deep basal magma ocean survived this process (Labrosse et al., 2007), we could expect substantial quantities of dissolved carbon species to be sequestered therein.

Most models suggest that the atmosphere forming at the end of the magma ocean was CO₂-dominated, with pressures ranging from 90 to 200 bar depending on the size of the degassing magma ocean (Gaillard et al., 2022a; Sossi et al., 2020). Assuming water was efficiently degassed from the solidifying mantle during the early Hadean, it has been proposed that its condensation in the form of early oceans could have efficiently pumped in atmospheric carbon into the deep Earth via carbonate mineral precipitation in a context of early plate tectonics and vigorous rock alteration (Sleep et al., 2011; Korenaga, 2021). This implies an early sequestration of deep carbon in the mantle via subduction of surface carbon.

A comparison between the mass of carbon outgassed by the magma ocean in the early Hadean and the present-day mass of surface carbon (atmosphere + oceans + crust) indicates surprisingly similar numbers of $110 \pm 50 \times 10^{21}$ g C (Gaillard et al., 2022a). This mass of surficial carbon is also indistinguishable from the mass of carbon constituting the dense Venusian atmosphere. A similar conclusion has been reached for nitrogen (Gaillard et al., 2022a), which can be explained by the very low fluxes of nitrogen involved in both present-day planetary outgassing and burying via subduction with respect to the size of the atmospheric reservoir (Gaillard et al., 2022a; Hirschmann, 2018). This may indicate that a surficial reservoir of carbon formed early in the history of planet Earth and has been only moderately affected by the subsequent 4.5 billion years of geological history. The balance of subduction versus outgassing might have maintained a steady state regime, although there are still many uncertainties in the actual fluxes of carbon involved in such geodynamic cycles (see Section "Carbon movement within and between deep reservoirs"; Dasgupta and Hirschmann, 2010; Iacono-Marziano et al., 2009; Kelemen and Manning, 2015) and it is difficult to discuss such processes further.

Toward the end of Earth's formation, a meteoritic bombardment — the so-called 'late veneer' — delivered significant portions of highly siderophile elements to the Earth in chondritic proportions, thereby elevating their BSE concentrations despite core differentiation (e.g., Walker, 2009; Wang and Becker, 2013). A volatile-rich late veneer, as suggested by Wang and Becker (2013), could have supplied a substantial proportion (20%–100%) of the carbon now comprising BSE. However, subsequent studies have cast doubt on the possibility of a solely late veneer contribution to BSE volatiles, including carbon. A simple mass balance calculation has shown that the mass of material required to account for the BSE carbon inventory (~2 ± 1 wt%) is equal to, or greater than, the mass delivered by the late veneer (0.1–2.0 wt%; Mikhail and Füri, 2019). Ruthenium, a highly siderophile element replenished in BSE by the late veneer, has an isotopic composition in Earth's mantle distinct from carbonaceous chondrite-derived material; this suggests that the late veneer was unlikely to be composed of the volatile-rich carbonaceous chondrites from which Earth's carbon is derived (Fischer-Gödde and Kleine, 2017). Instead, it is reasonable to suggest that the bulk of Earth's carbon was obtained during the principal stages of planetary accretion, as discussed above, with a minor fraction of carbon subsequently added by a late veneer (e.g., Hirschmann, 2016). This is supported by isotope fractionation modeling of sulfur and chalcophile elements, which suggests that planetary differentiation is sufficient to account for volatile element concentrations in the BSE with less than 40% delivered by the late veneer (Wang et al., 2023).

Once the magma ocean was solidified, volcanism on Earth was mostly mafic (i.e., forming basaltic lavas), accompanied by superhot lavas such as Mg-rich komatiites (Fig. 1, top). No constraint allows us to understand the carbon outgassed by these magmas. The oldest available constraints are probably encapsulated within diamonds; diamonds older than 3.0 Ga host peridotitic inclusions, whereas younger diamonds prevalently host eclogitic inclusions (Shirey and Richardson, 2011; see Section "Diamonds as clues on deep and super-deep carbon movements" for a full discussion on inclusions in diamond). This is possibly a signature of the initiation of plate tectonics at the end of the Archean eon.

Cratons are thick (>200 km) and cold lithospheric blocks which have experienced little, if any, geological activity since the Archean. They host kimberlites, which constitute the dominant form of extrusive rocks containing diamonds (see Section "Diamonds over (geological) time"). Mantle xenoliths found in kimberlites and in many other intraplate lavas have revealed that mantle metasomatism, i.e., melt-fluid-rock interactions taking place at the base of lithosphere (O'Reilly and Griffin, 2013), is an important process, modifying the physical and compositional properties of cratonic mantle lithosphere and possibly storing an important amount of carbon (see Section "Sub-continental lithospheric mantle"); Aulbach et al., 2017; Hirschmann, 2018). To some extent, lithospheric diamonds can be seen as resulting from such metasomatism, and Ca-Mg carbonate minerals have also been widely observed. A hidden carbon reservoir might thus have continuously formed over time (Hammouda et al., 2021) by melt-fluid-rock interactions taking place in cratonic, non-cratonic, and even oceanic lithosphere (Aulbach et al., 2017; Delpech et al., 2023; Foley and Fischer, 2017; Hammouda et al., 2021; Hirano and Machida, 2022; Hirschmann, 2018; O'Reilly and Griffin, 2013).
The final picture of the geological history of carbon corresponds to the present-day outgassing of volcanoes — mid-ocean ridges, petit spots, hot spots, volcanic arcs, etc. (Fig. 1) — which are sourced by various reservoirs built by 4.5 billion years of geodynamic processes with a key role played by subduction processes (Aiuppa et al., 2017; Dasgupta and Hirschmann, 2010; Hammouda et al., 2021; Hirano and Machida, 2022; Plank and Langmuir, 1998). Carbon isotopic measurements of volcanic gasses in arc settings suggest the major involvement of sedimentary carbonates in degassing, with the possibility of crustal assimilation processes short-circuiting the deep carbon cycle (Section “Movements of deep carbon in melts”; Iacono-Marziano et al., 2009; Mason et al., 2017).

Carbon at the interface between the geosphere and biosphere

Deep carbon movements and transformations on Earth result from complicated and interconnected geological and biological processes. A large fraction of solid carbon in the crust is formed through biological processes and their successive solid-state transformations. For example, carbonate minerals can be synthesized by various organisms that grow a carbonate shell such as coccolithophores and other phytoplankton, be precipitated as a result of microbial metabolism such as in chemosynthetic ecosystems on the seafloor, or can also be dissolved by endoliths and other microorganisms during their metabolic reactions. The organic component of biological matter, or its transformation, represents the largest fraction of organic carbon on the surface of our planet. The main biological processes interacting with carbon and their related fluxes are presented in Fig. 2.

While biology might appear as a thin veneer on the surface of our planet, its interactions with the solid Earth are far reaching (Vernadsky, 1926). Life originated on this planet between 3.8 and 4.0 Ga, and soon after its appearance it has taken control over several surface and near-surface chemical cycles, deeply impacting the movement, speciation, and quantity of elements on Earth (Falkowski et al., 2008). Subducted organic carbon material as old as 3.0 Ga, as reflected from the carbon isotopic signature of diamonds (Howell et al., 2020), reveals that the biosphere has extended its reach far beyond the habitable area of the near planetary surface, and it is possible that today no Earth carbon reservoir, with the exception perhaps of the core, is devoid of the fingerprints of life.

This section introduces the contribution of biology to deep carbon processes and fluxes, reflecting on the complex geo-biological interactions that underpin the interplay between the fast and slow carbon cycle.

<figure><img src="image_4.png" /><figcaption>**Fig. 2** Summary of the major biological processes interacting with carbon that have direct effects and interactions with the deep carbon cycle. Major processes are colored according to the division between terrestrial (green), marine (blue) and deep subsurface (red) processes. Global estimates of biomass are reported (Bar-On et al., 2018). Gross (photosynthetic) primary productivity on land: $110–150 \times 10^{15} \text{ g C yr}^{-1}$ (Jian et al., 2022; Lu et al., 2024). Organic carbon diagenesis: $64–120 \times 10^{12} \text{ g C yr}^{-1}$ (Kump and Arthur, 1999). Heterotrophy at the ocean surface: $143 \times 10^{9} \text{ g C yr}^{-1}$ (del Giorgio and Duarte, 2002). Gross (photosynthetic) primary productivity in the ocean: $103–150 \times 10^{15} \text{ g C yr}^{-1}$ (Huang et al., 2021a,b). Water column chemosynthesis (in the photic zone, dark ocean and surface marine sediments): $\sim 1 \times 10^{15} \text{ g C yr}^{-1}$ (Middelburg, 2011). Marine organic carbon burial and diagenesis: $20–200 \times 10^{12} \text{ g C yr}^{-1}$ (Cartapanis et al., 2018; Dunne et al., 2007; Hayes et al., 2021; Jahnke, 1996). Oceanic sediment methane production is $75–320 \times 10^{12} \text{ g C yr}^{-1}$ (Valentine, 2002). Utilization of deep volatiles in the continental crust (and oceanic crust) is still unconstrained (subscript a), but likely to be significant; see Barry et al. (2019) and Fullerton et al. (2021) for estimates at the Costa Rican convergent margin ($\sim 10^{8}–10^{10} \text{ mol C yr}^{-1}$). Both continental and oceanic subsurface chemosynthesis is currently unconstrained (subscript b).</figcaption></figure>
Carbon's central role in life

Carbon is not merely a ubiquitous element; it is the cornerstone of biochemistry on Earth. Its ability to form stable, complex molecules with a variety of elements, especially hydrogen, oxygen, nitrogen, phosphorus, sulfur and minute amounts of metals underpins the chemistry of life. From the complex carbohydrates that fuel cellular processes to the intricate information storage systems of nucleic acids, carbon forms the scaffolding upon which life is built. Additionally, different carbon species play crucial roles in energy generation and electron transfer reactions that sustain life. Reduced carbon compounds like methane (CH₄) and carbon monoxide (CO) serve as energy sources for chemolithotrophic microorganisms, while oxidized forms like carbon dioxide (CO₂) and bicarbonate (HCO₃⁻) act as electron acceptors in various metabolic pathways.

Carbon is incorporated into building blocks through the process of carbon fixation, either starting from CO₂, HCO₃⁻, CO, CH₃-group and CH₄ as carbon sources. These carbon species are taken up by cells and incorporated into organic molecules following diverse carbon fixation pathways. To date, six carbon fixation pathways are known, namely the Calvin-Benson Cycle, the Arnon-Buchanan cycle, the Wood-Ljungdahl pathway, the hydroxypropionate bicycle, the 4-hydroxybutyrate pathway, the hydroxypropionate/hydroxybutyrate cycle and the dicarboxylate/hydroxybutyrate cycle (Fuchs, 2011). Of the six pathways, the Calvin-Benson cycle (most frequently called the Calvin cycle) was the first one to be elucidated and it is the pathway that currently dominates carbon fixation on the surface of planet Earth, probably due to the robustness of its enzymes to atmospheric molecular oxygen. Used by phototrophs, including plants, algae and cyanobacteria, and a number of chemolithotrophic (that is non-photosynthetic autotrophs) bacteria, the Calvin cycle accounts for the large majority of the organic carbon produced under the influence of the sun on the surface of the planet. Organic carbon produced using this pathway fractionates relative to the source carbon (Δ¹³C) between -20‰ and -30‰, generally producing organic carbon with a distinctive δ¹³C (VPDB) isotopic signature between -25‰ and -28‰, providing a unique marker for tracking carbon fixed using oxygenic photosynthetic processes (Fuchs, 2011).

The other five carbon fixation pathways play a key role in anaerobic microorganisms inhabiting marine sediments today, a variety of extreme environments, and the continental and oceanic subsurface. The Arnon-Buchanan cycle, also known as the reverse tricarboxylic acid (TCA) cycle or reverse Krebs cycle, is abundant in anoxygenic phototrophs, and a large number of anaerobic and microaerophilic chemolithotrophic microorganisms. Its δ¹³C carbon isotopic fractionation (Δ¹³C) is much smaller than the Calvin cycle, and generally between -2‰ and -12‰. Recently, it has been shown to be dominant in subsurface ecosystems at convergent margins (Fullerton et al., 2021; Rogers et al., 2023), and it has been previously identified as a key carbon fixation pathway in deep-sea hydrothermal vents (Hügler and Sievert, 2011; Sievert and Vetriani, 2012). The Wood-Ljungdahl pathway, also known as the reductive acetyl-CoA pathway, has been originally described in acetogenic bacteria and it is active in a number of strictly anaerobic bacteria and archaea, including methanogens and sulfate reducing bacteria, both of which play a key role in linking the fast and slow carbon cycle. The Wood-Ljungdahl pathway is capable of producing some of the most ¹³C-depleted organic carbon on this planet, with carbon isotopic fractionation Δ¹³C < -30‰ and reported values as low as -65‰ Freude and Blaser, 2016. The methane produced using this pathway by methanogens is also distinctively depleted in ¹³C (δ¹³C < -45‰) compared to abiotic and thermogenic methane sources (Whiticar, 1999). This "light" carbon isotopic signature can be retained in authigenic carbonate minerals (that is carbonate minerals directly precipitated within the sedimentary sequences rather than exported from the surface) created in chemosynthetic environments by the oxidation of methane, and can reach extreme δ¹³C values below -70‰. This is a value markedly different from the carbonate minerals created by photosynthetic organisms with a δ¹³C ~0‰. The remaining carbon fixation pathways play a minor quantitative role in ecosystems, with the exception of the hydroxypropionate/hydroxybutyrate cycle that is used by the ammonia-oxidizing archaea that dominate carbon fixation in the dark oceanic water column (Könneke et al., 2014). These examples highlight the diversity of interactions that life has with carbon as a building block, and they suggest that the isotopic signature of organic carbon and carbonates subducted and recycled in the deep Earth might have important local variations depending on the microbial processes involved.

Beyond its structural role, carbon compounds play critical energetic roles in ecosystems, especially in environments where light and oxygen are scarce. Reduced carbon forms such as CO and CH₄ can serve as electron donors (i.e., energy sources) in microbial redox energy reactions and are important energy sources especially in subsurface environments. Oxidized carbon forms, such as CO₂ and HCO₃⁻ can instead be used as terminal electron acceptors (i.e., oxidants) in microbial redox metabolic reactions, and are quantitatively important in several ecosystems despite the low Gibbs free energy released. Diverse organisms, both aerobic and anaerobic, can utilize these carbon compounds as substrates in their redox metabolism, altering the quality and quantity of the carbon cycling through diverse geological reservoirs. While oxidized carbon molecules (i.e., CO₂ and HCO₃⁻) can only be used in conjunction with hydrogen (H₂) as electron donor to generate compounds such as CH₄ and acetate, the reduced carbon forms, and in particular methane can be coupled to a diverse array of electron acceptors by taxonomically and metabolically diverse microorganisms (Hay Mele et al., 2023). Methane can be oxidized aerobically at the expense of oxygen (O₂) by aerobic methanotrophs, as well as anaerobically coupled to diverse electron acceptors such as sulfate (SO₄²⁻), iron (Fe³⁺), manganese (Mn⁴⁺) and nitrite (NO₂⁻) (Guerrero-Cruz et al., 2021).

The deep subsurface biosphere

While the interaction of life with carbon is at the basis of the fast — and near-surface — carbon cycle (Lee et al., 2019), the discovery of the deep subsurface biosphere has revolutionized our understanding of the interaction of life with the deep carbon cycle. Since
first proposed in the early 1990s (Gold, 1992), the deep biosphere has constantly challenged our understanding of the limit and adaptation of life and its interaction with the solid Earth (Colman et al., 2017). After 30 years of research we know that up to 14% of Earth's biomass is stored in the subsurface (Bar-On et al., 2018), and microbial communities thrive kilometers below the surface, inhabiting diverse environments within the Earth's crust, oceanic sediments, and subseafloor (Kallmeyer et al., 2012; Magnabosco et al., 2018). This hidden biosphere exhibits remarkable metabolic diversity, with organisms adapted to extreme conditions of temperature, pressure, and nutrient availability, and it is capable of interacting with geological processes at extremely diverse spatial and temporal scales, matching scale processes such subduction, organic matter diagenesis and geological volatile cycling (Giovannelli et al., 2022).

Communities in the deep subsurface are largely composed of bacteria and archaea, along with a very small number of eukaryotic microbes (Bar-On et al., 2018). These microorganisms exhibit remarkable adaptations that allow them to exploit redox energy couples provided by the geosphere — such as hydrogen or methane — together with organic carbon produced in situ biotically, and potentially abiotically, or derived from the surface. Their distribution is controlled by diverse factors such as energy availability, geochemistry, temperature, pressure and available liquid water and pore space (Heuer et al., 2020). A key factor controlling their distribution appears to be temperature, with the current upper limit for life set at 122 °C in the laboratory (Takai et al., 2008) and believed to be around 135 °C in natural ecosystems (Heuer et al., 2020; Merino et al., 2019). Current data suggests that the deep subsurface biosphere extends from a few meters below the surface to 3–5 km in marine and continental settings respectively (Magnabosco et al., 2018), with theoretical studies suggesting that under favorable conditions, such as those found in cold subduction zones, the habitable zone could extend up to 15–20 km depth (Plümper et al., 2017). These depths, together with the diversity of metabolic processes and the ability to interact with geological processes over long timescales (Lloyd, 2020), allow subsurface microbial communities to impact carbon and other volatiles cycling at planetary scales (Barry et al., 2022; Giovannelli et al., 2021, 2022). The upper lithosphere can thus be seen as a giant biofilter (Giovannelli et al., 2021), harboring microbial communities capable of altering the influx (mainly in marine sediments) and the outflux (in the oceanic and continental crust) of carbon materials, affecting the quantities of carbon recycled, its redox state and its isotopic signature.

## Impact of life and organic carbon on the deep carbon cycle

The biological production of organic matter and its subsequent deposition and burial lead to significant transformations in carbon reservoirs, which can ultimately be subducted and recycled within the Earth's interior (Fig. 2). This process introduces vast quantities of carbon into Earth's lithosphere and beneath, contributing to the formation of fossil fuel deposits and shaping the composition of deep carbon reservoirs. The evolution of oxygenic photosynthesis and the subsequent rise of oxygen in the atmosphere led to a shift from carbonate-based to organic-carbon-based burial of carbon at the beginning of the Proterozoic eon (Galvez et al., 2020), highlighting the profound influence of biological evolution on deep carbon cycling.

The metabolic activities of deep life significantly impact deep carbon cycling. Microorganisms facilitate the transformation of various carbon species, contributing to processes such as methanogenesis, hydrocarbon oxidation, and carbonate mineral precipitation. These activities alter the composition and mobility of carbon in deep reservoirs, influencing long-term carbon storage and release. For instance, methanogenesis in deep subsurface environments contributes to the global methane budget, while microbial oxidation of hydrocarbons can impact the quality and recovery rates of oil and gas reservoirs. Additionally, microbial processes contribute to the precipitation and dissolution of carbonate minerals, playing a role in the long-term carbon cycle.

Despite significant advancements in our understanding of the deep biosphere and its interactions with deep carbon, many questions remain unanswered. Quantifying the total biomass and metabolic activity of deep life remains a challenge due to limitations in sampling and detection methods. Additionally, most deep subsurface organisms remain uncultured, hindering our understanding of their metabolic capabilities and their role in biogeochemical cycles. Furthermore, the impact of anthropogenic activities and natural environmental changes on deep life and deep carbon cycling requires further investigation. Addressing these open questions is crucial for developing a comprehensive understanding of the Earth system and its response to environmental change. Further research at the interface of geochemistry and microbiology promises to illuminate the intricate interplay between deep carbon and the biosphere, revealing the hidden secrets of Earth's deepest realms.

## Biogenic carbon: From the surface to the subsurface of the planet

The vast majority of inorganic carbon enters the biosphere through the Calvin cycle enzyme rubisco (Raven, 2009). Rubisco is an ancient enzyme that appeared in archaea from ~4.0 Ga. Over time, it evolved into the primary carbon-fixing enzyme, notably in phototrophs like cyanobacteria and terrestrial plants, facilitating the historical accumulation of biogenic carbon on Earth (i.e., carbon that is either part of, or derived from, living organisms).

At the Earth's surface — on the land and the marine photic zone — biogenic carbon fixation is mostly ensured by plants and microalgae respectively through photosynthesis (Bar-On et al., 2018). It is estimated that gross primary productivity, a metric quantifying the amount of CO₂ removed from the atmosphere every year to fuel photosynthesis, is evenly distributed between land and ocean, each contributing ≈100–150 Gt C yr⁻¹ (Fig. 2; Huang et al., 2021a, Huang et al., 2021b; Jian et al., 2022; Lu et al., 2024). These primary producers form the basis of terrestrial and marine food webs, as they are consumed by heterotrophs, which build their biomass out of this fixed biogenic carbon. When these organisms die, their organic matter — when not decomposed — can accumulate and become buried under layers of sediment. On land, biogenic carbon-rich organic matter can for instance be
preserved from decomposition by mud or acidic water, eventually forming peat deposits that, over millions of years, transform into coal. Overall, ≈0.2 Gt C are buried in the ground each year (Kandasamy and Nagender Nath, 2016). Alternatively, this terrestrial biogenic carbon can be transported to the oceans by water motion and wind (with an estimated terrestrial carbon input to the ocean of 1.4 ± 0.5 Gt C yr⁻¹; Kwon et al., 2021). There, it joins a pool of biogenic carbon that is fixed in situ by photosynthetic phytoplankton. While most biogenic carbon at the ocean surface is recycled by local consumption (respiration by heterotrophic organisms), a small fraction of it (≈10 Gt C yr⁻¹) sinks or is transported down (by migrating animals or ocean circulation), and eventually joins the deep sea organic carbon stock (Siegel et al., 2023).

Variations in global and regional climate, tectonic activity, and the development of large depositional basins have played critical roles in determining the rates of carbon accumulation and the formation of these different deposits. These carbon reservoirs are therefore not uniformly distributed across the planet. Large coal deposits, for instance, are often found in regions that were historically forested and subject to specific geological conditions favoring the preservation of organic material. Oil and natural gas fields are frequently associated with marine sedimentary basins that accumulated vast amounts of organic matter. In the deep sea, most biogenic carbon comes from the surface. However, primary productivity performed by chemolithotrophs, which obtain energy by the oxidation of environmental electron donors (e.g., ammonium, nitrite, ferrous iron, sulfide), can significantly contribute to the organic carbon pool (≈0.1 Gt C yr⁻¹ in the deep ocean; Middelburg, 2011). Nitrification, the oxidation of ammonia to nitrite and nitrate, is understood to be the dominant chemolithotrophic metabolism in the deep ocean, while diverse ecosystems (e.g., hydrothermal vent, cold seeps, mud volcanoes, whale falls, oxygen minimum zones) host diverse communities of chemolithotrophs, further contributing to the total rate of non-phototrophic biogenic carbon fixation in the ocean (0.77 Gt C yr⁻¹; Middelburg, 2011).

All the biogenic carbon that is not recycled back into inorganic carbon (by respiration or fermentation) can eventually be buried under layers of sediment. There, the fate of biogenic carbon is an important component of the process of diagenesis, the process by which sedimentary deposits are transformed into a sedimentary rock over the years (Fig. 2). This is largely influenced by local biotic processes (LaRowe et al., 2020). While, in the upper layers of shallow sediments, aerobic respiration remineralizes biogenic carbon, in the deeper layers, and, by extension in all deep-ocean sediments, anaerobic respiration and fermentation processes dominate. The most prevalent one is sulfate reduction (Canfield and Des Marais, 1991). Sulfate-reducing microorganisms use the abundant oceanic sulfate as an electron acceptor to oxidize biogenic carbon (Jørgensen et al., 2019). Interestingly, sulfate is not the most favorable oxidizer thermodynamically (as compared to nitrate or ferric oxide), however its greater abundance in seawater makes it the dominant electron acceptor in the anaerobic oxidation of biogenic carbon in marine sediments. Deeper below, where sulfate is scarce, another metabolism, methanogenesis, is prevalent (e.g., Bojanova et al., 2023). Methanogens reduce organic matter or CO₂ into methane, releasing biogenic carbon that can seep up. Between the sulfate and methane zones lies the sulfate-methane transition zone (SMTZ), where both species cohabit. Within the SMTZ, a third metabolism, the anaerobic oxidation of methane, is dominant. Anaerobic oxidation of methane uses the sulfate to oxidize methane produced below by methanogens, thereby releasing CO₂ and hydrogen sulfide. Stable isotope analysis can track these stratified metabolisms, as each reaction leaves a characteristic isotopic signature, yielding a range of signatures in diagenetic samples (Meister et al., 2019).

A part of biogenic carbon in sediment remains unalterable. Chemically, intermolecular cross-links (Kunhi Mouvenchery et al., 2012) or abiotic modifications, such as sulfurization or peptide deamination (Abdulla et al., 2018; Burdige, 2005), decrease the bioavailability of the organic matter in the sediments. Physically, the phenomenon of sorption of organic matter to mineral surfaces also protects biogenic carbon remineralization by encapsulating it, denaturing extracellular enzymes, and limiting molecular diffusion. In the long run, the quantity of biogenic carbon buried therefore depends on the balance between a continuous input from sedimenting biological matter and its oxidation by local microbial life, in parallel with the ongoing process of sorption and the abiotic alteration of the matter, protecting this carbon from further oxidation.

Finally, another component of the deep carbon cycle is the 'carbonate pump' which contributes to carbon sequestration in the subsurface. Calcifying organisms use dissolved inorganic carbon in seawater to build calcite platelets called coccoliths. Upon their death, these organisms contribute to the accumulation of biogenic carbonate sediments on the ocean floor. Over geological timescales, these sediments can undergo lithification, forming carbonate rocks such as limestone (Folk, 1980).

Of the 10 Gt of biogenic organic carbon transported from the surface to the deep ocean annually, only 0.02–0.2 Gt C yr⁻¹ is eventually incorporated in ocean sediments (Cartapanis et al., 2018; Dunne et al., 2007; Hayes et al., 2021; Jahnke, 1996). Most of this biogenic carbon is therefore accumulating at various depths until it is eventually reoxidized, effectively sequestering it away from surface ocean and atmospheric exchange over timescales ranging from months to millennia, depending on the depth.

## Effect of life on carbon fluxes to and from the deep Earth

Over geological timescales, organic-matter-rich sediments can be subducted into the Earth's mantle along convergent plate boundaries (Fig. 2; Section "Subduction of crustal carbon"). Microbial communities can survive in sediments that are subducted to depths of several kilometers, where they continue to influence the transformation of carbon (Barry et al., 2019; Debret et al., 2022; Fryer et al., 2020; Plümper et al., 2017). In particular, these communities can rely on the interaction with fluids associated with serpininizing forearc mantle: reduced volatile species from these fluids (e.g., H₂, CH₄) encounter more oxidized redox couples from the subducted crust, thereby supporting deep chemolithotrophic life (Mottl et al., 2003; Ohara et al., 2012; Vitale Brovarone et al., 2020a). Chemolithotrophy leads to the formation of partially reduced organic carbon compounds (e.g., lipids, carbohydrates,
proteins, hydrocarbons), which may either be subducted further into the deep mantle or, heated under high-pressure conditions, can decompose and release remineralized carbon (CO₂, CH₄).

The subduction of biogenic carbon has significant implications for volcanic activity and the global carbon cycle. Organic carbon and carbonates carried into subduction zones can decrease the melting point of the mantle, leading to enhanced magma generation (Plank and Manning, 2019; see Section “Movements of deep carbon in melts”). This magma, when it reaches the Earth’s surface through volcanic eruptions, releases substantial amounts of CO₂ back into the atmosphere together with variable amounts of reduced carbon compounds (e.g., Giggenbach, 1996; Symonds et al., 1994). The amount and speciation of carbonic fluids released at volcanic arcs may be directly influenced by surface biology. For example, recent examination of melt inclusions from a global collection of Cenozoic arc magmas has shown that latitudinal variations in phytoplankton production may directly influence the redox state of the resulting arc magmas (Hu et al., 2024).

Degassing and rising fluids containing reduced carbon species (and other reduced compounds, e.g., H₂, NH₃, H₂S) have the potential to fuel subsurface life on their way to the surface, as evidenced by the abundance of chemolithotrophic microbial communities thriving in various volcanic and gas-seeping ecosystems. Information regarding the quantitative impact of subsurface life on the carbon recycled to the surface are limited, but suggests that potentially large amounts of the carbon compounds (and other volatiles) recycled to the surface might be altered and sequestered by the subsurface biosphere in the overriding crust. For example, recent work suggests that subsurface microbial communities present in the forearc region of convergent margins can sequester into biomass up to 90% of the carbon recycled to the surface, reducing the carbon going to the mantle by up to 40% (Barry et al., 2019; Fullerton et al., 2021). Dissolved organic carbon in the sampled forearc springs has δ¹³C signatures indicative of chemolithotrophic activity utilizing CO₂ originating from the subducting plate (Fullerton et al., 2021). Additionally, near-surface microbial communities associated with areas where reduced fluids and gas are recycled to the surface utilize these compounds for their metabolism, further influencing the diversity and fate of carbon molecules and impacting the carbon cycle. In this context, the interactions between subsurface life and the solid Earth play a significant role in the deep carbon cycle, impacting and altering the flux between the Earth's interior and surface. Constraining the contribution of life in altering the quality and quantity of carbon recycled to the surface will be of key importance to understand the interactions and feedbacks between biology and the solid Earth.

## Deep carbon forms

The deep forms of carbon are diverse and include multiple redox states and coordination numbers depending on their origin and location within the deep Earth (Fig. 3). The most common solid carbon forms are carbonate minerals, organic matter, graphite, diamond, and carbides. Their movement is largely controlled by redox changes and solubility. This section covers the deep reservoirs of the Earth from the core to the crust, passing through the mantle.

## Carbon in the Earth's core

Earth's core is located 2890 km below the surface and is primarily made of iron (85%) and nickel (5%). Small amounts of lighter siderophile elements, including hydrogen, oxygen, sulfur, silicon, and carbon are mandatory to match the density deficit and velocity excess of the Preliminary Reference Earth Model (PREM; Dziewonski and Anderson, 1981). As the Earth's core constitutes 32% of the mass of our planet, even minute amounts of carbon in the core (between 0.1 and 0.5 wt% C depending on estimates) make it the largest planetary carbon reservoir, containing approximately 90% of the bulk Earth carbon budget (see Fischer et al., 2020, and references therein).

As the Earth's core is not directly accessible, estimates of its carbon content rely on our interpretation of various observations. The seismic velocity and elastic parameters of the Earth's core indicate that it is ~10% less dense than iron-nickel alloy under the core pressure-temperature range (136–364 GPa and 4000 K at the core-mantle boundary, CMB) primarily due to the presence of light elements. Comparison of elemental concentrations in BSE relative to primitive meteorites and the volatility trend of elements indicates that the core is enriched in sulfur, hydrogen, and carbon as far as light volatile elements are concerned. Isotopic fractionation during core formation leads to one of the highest estimates of core carbon concentrations (0.5–2.0 wt%; Wood et al., 2013).

The relative enrichment of the core in carbon relative to the BSE was acquired when the core segregated from the mantle silicate reservoir during the collisional accretion stage of the Earth, which extended 10–100 million years after the initial solar nebula processes (see Section “Origins of deep carbon”; Barboni et al., 2017; Jacobson et al., 2014). The metal delivered to the growing Earth that ultimately formed the core existed as droplets that equilibrated at least partially with the magma ocean, while themselves sinking due to the higher density of the metal phase relative to silicate melt. The environment, depth, and redox conditions under which carbon partitioned between silicate and metal phases largely dictated the carbon content of the Earth's core. At low to moderate pressure and high temperature, carbon is siderophile. Our modern societies have largely taken advantage of this behavior, expanding our understanding of carbon partitioning behavior in metals through the production and intensive use of steel (i.e., iron carbide), and developing adjustable properties depending on the addition of other siderophile elements (e.g., chromium and often nickel to obtain stainless steel). The high cosmochemical abundance of carbon, coupled with its low BSE abundance (0.01%), has led to the suggestion that carbon dissolved extensively into the core, and that its concentration in the core is on the order of 2–4 wt% C. This opens the idea that carbon could be a major element in the inner core, which may contain high-pressure carbides such as
<figure><img src="image_5.png" /><figcaption>**Fig. 3** Schematic diagram illustrating the distribution, forms, and budgets of deep carbon in the Earth. Carbon in the core is distributed between solid iron carbides in the inner core and dissolved phase in the liquid outer core (Fischer et al., 2020). In the lowermost mantle, carbon is present as sp³-hybridized carbonate, diamond, iron carbides, and fluids and melts; carbonate is sp²-hybridized at the mantle transition zone and in the upper mantle (Stagno et al., 2019). Redox freezing at these pressures will convert carbonate into diamond as it descends into the lower mantle; likewise, redox melting will convert diamond into carbonate during decompression (e.g., Foley, 2021; Rohrbach and Schmidt, 2011). In the lithosphere, carbon takes multiple forms, including graphite (or diamond in thick sub-continental lithospheric mantle), carbonate minerals, and organic carbon (e.g., Hazen et al., 2013a). During mantle melting, carbon is incompatible in mantle minerals and will join the melt phase as carbonatite melts (e.g., Jones et al., 2013) or dissolved CO₂ in silicate melts, which exsolves into a vapor at low pressures (e.g., Dixon et al., 1995). Volcanic gas speciation is controlled by the redox conditions of the magma and exsolving gas (e.g., Giggenbach, 1996). Fluid-rock reactions, such as those at sub-seafloor conditions or during subduction, and metamorphism can generate other carbon forms, including methane (CH₄) (e.g., Connolly and Cesare, 1993; Holloway, 1984; Kelley et al., 2005; Vitale Brovarone et al., 2020a), organic acids and polycyclic aromatic hydrocarbons (PAHs; Andreani and Ménez, 2019, Debret et al., 2022).</figcaption></figure>

Fe₃C or Fe₇C₃ (Wood, 1993; Dasgupta et al., 2013a). Therefore, Fe-C alloy binaries have been extensively investigated under extreme pressure-temperature conditions. Their density and elastic properties were measured, as a viable composition model of the core must account for the density deficit and seismic characteristics of the core. For the inner core, both experiments and theory suggest that carbon and hydrogen are key impurity elements to explain the Poisson ratio (Vp/Vs) observed in the inner core as compared to pure iron. Among candidate Fe-C alloys and carbides, Fe₇C₃ is a good candidate from its ability to match the anomalously low Vs and high Poisson ratio of the inner core, in addition to reproducing the density deficit and having the highest electrical resistivity among all Fe-C alloys (Zhang et al., 2018a). A caveat is its density, which may be too low to match the inner core density, and would more generally imply a high carbon concentration in the core. This contradicts experiments performed at the pressures and temperatures expected in a deep magma ocean during formation of Earth's core which demonstrate that carbon becomes much less siderophile at depth (Fischer et al., 2020). Fischer et al. estimated that the core likely contains a maximum of
0.09 ± 0.04 to 0.20 ± 0.10 wt% C, which is in agreement with most of the theoretical atomistic and experimental mineral physics studies that were performed at or close to core pressures.

As the composition of the core is central to our understanding of the Earth's accretion and the terrestrial magnetic field, hundreds of studies focusing on mineral physics, cosmochemical constraints, or geochemical experiments have attempted to decipher the composition of the Earth's inner or outer core and provide a match with the seismic properties of the core. Depending on the chosen theoretical or experimental parameters and strategies, the pressure-temperature range covered by these studies, and the cocktail of light elements considered, results vary. Recent results still propose that the carbon concentration in the outer core spans one order of magnitude between 0.1 and 2.0 wt% (Bajgain et al., 2021). Existing data are yet insufficient to resolve the competing models of core composition because of limited data coverage in the relevant challenging pressure-temperature-composition space and uncertainties in experimental measurements and theoretical simulations. In their recent review, Hirose et al. (2021) propose a likely range of composition for the outer and inner core, with approximately 0.2 wt% and 0–1.3 wt% C respectively, depending on the concentration of other light elements.

Future progress and better constraints on the core composition and carbon budget could arise from the investigation of the interactions between melt and solid at the inner core boundary, as techniques and methods progressively allow us to reach inner core boundary pressure (330 GPa); inner core boundary temperatures, which may be as high as 7000 K, have not yet been attained. This would inform us on the evolution of Earth's magnetic field through time. The thermal and compositional evolution of Earth's core is related to the growth of the inner core. Since light elements control the liquidus temperature of Fe alloy and therefore the crystallization of the core during secular cooling, they could substantially increase the thermal conductivity of the outer core with profound implications for core thermal evolution and geodynamics. Inner core viscosity and dynamics are also certainly dependent on the light element composition through diffusion rate, grain growth rate and shear modulus, with virtually nothing known at this stage.

Moving upwards through the Earth to the outermost part of the core, seismology has revealed a ~400 km thick low-velocity and low-density layer named the E' layer (Fig. 3), which is characterized by a steep $V_p$ gradient (Brodholt and Badro, 2017; Helffrich and Kaneshima, 2010; Kaneshima, 2018; Kaneshima and Helffrich, 2013; Kaneshima and Matsuzawa, 2015). The E' layer is interpreted as a remnant of the chemical exchange between the early formed core and the whole-Earth magma ocean. The stagnation of such a layer over time necessitates buoyancy relative to the underlying convecting outer core, which could be explained by an enrichment in oxygen and depletion in silicon, but probably not at the 100-km scale. Presence of a significant amount of carbon in the E' layer is unlikely, as adding carbon slightly decreases the density of liquid iron and increases its $V_p$ (Morard et al., 2017; Nakajima et al., 2015).

Despite the challenges to reach the pressure and temperature conditions characteristic of the core-mantle boundary and deeper, the knowledge of the physical properties of Fe alloys and melts including carbon and other light elements has incredibly improved over the last decades. Even for the lowest estimated carbon contents (~0.2 wt%), the largest carbon reservoir of planet Earth is its core, which hosts at least 90% of the whole Earth carbon budget and is where carbon has been sequestered since the final stages of core formation.

## Carbon at the interface between the core and the lowermost mantle

The lowermost mantle above the core-mantle boundary, referred to as the D" layer, displays multiple seismic features (Fig. 3). It is also a strong thermal boundary layer, with temperatures increasing from an average mantle adiabat of about 2500 K at 2600 km depth (e.g., Stixrude et al., 2009) to a core-mantle boundary temperature of about 4000 K. The D" layer is both vertically and laterally highly heterogeneous, and it is layered on a kilometer scale in the lower 50 km (Sidorin et al., 1999; Thybo et al., 2003). Because of diffraction of seismic waves by the core, the resolution in this layer is lower than in the overlying mantle and hinders access to fine details. Velocity profiles show large lateral heterogeneities in this region. Estimates of the thickness of the D" layer suggest that it ranges from 100 km to about 400 km. Within this layer, unusually low wavespeeds have been measured beneath the African continent and the Pacific Ocean, and these regions are known as large low-shear-velocity provinces (LLSVPs) and exhibit a density excess of about 1.25% compared to the surrounding mantle. Most active volcanic hotspots on Earth's surface are located above and at the boundaries of LLSVPs.

Adjacent to the LLSVPs, seismic observations show ultra-low velocity zones (ULVZs), often in contact with the core-mantle boundary. ULVZs are not ubiquitous. They also have a density excess and a low seismic velocity feature as indicated by their name. Their $V_p$ and $V_s$ reduction is about 10% and 30%, respectively (e.g., Lay, 2005). Although the presence of ULVZs and LLSVPs is not questioned, their origin and composition remain uncertain and there are multiple conceptual models that are consistent with seismic tomography results and the geochemistry of the hotspot lavas of oceanic islands, which are commonly assumed to result from the partial melting of large rising mantle plumes deeply rooted in the LLSVPs and ULVZs (see reviews by Garnero et al., 2016; Weis et al., 2023).

The ULVZs may represent partially leaky 'windows' between the D" layer and the outer core. Such localized core-mantle exchanges are supported by geochemical investigations of plume-related volcanic rocks. Recent isotopic measurements have revealed a negative correlation between the $^{182}W/^{184}W$ and $^{3}He/^{4}He$ ratios in major plume-related suites of oceanic island basalts (see review by Weis et al., 2023). Plumes giving rise to ocean island basalts with low $^{182}W/^{184}W$ ratios might have sampled core metal, most likely via some ULVZs located in the root zones of plumes (Rizo et al., 2019).
The location of LLSVPs and ULVZs between regions of sinking tectonic plates is intimately associated with the long-term dynamics of the convecting mantle, although it is unclear whether they are inherited from the early evolution of the mantle and related with remnants of the deep magma ocean, or they developed over time (e.g., McNamara, 2019; Flament et al., 2022). As long-lived structures, stable or slowly mobile at the core-mantle boundary over hundreds of millions of years, they will be negatively buoyant relative to the convecting mantle, and their composition is likely correlated to the observations of their increased intrinsic density and seismic velocity reduction.

Among the multiple models of LLSVPs and ULVZs, we focus here on the potential role of carbon-bearing phases that may explain the density excess and low shear wave pattern of LLSVPs. Enhanced iron content in deep mantle mineral assemblages is the easiest way to account for the density excess and low velocity of LLSVPs and ULVZs, and low partial melting could also explain the low-velocity features. Liu et al. (2016) proposed that metallic Fe-C produced in subducted slabs beyond a depth of 250 km may partially melt at core-mantle boundary conditions, since the eutectic melting curve of the Fe-C system crosses the geotherm near Earth's core-mantle boundary. At the eutectic, the Fe-C melt could contain 2 wt% C at the pressure and temperature conditions of the core-mantle boundary (Fei and Brosh, 2014; Lord et al., 2009). If concentrated into isolated patches, such a metallic melt could induce the seismically observed density and velocity features of ULVZs as it is twice as dense as the surrounding mantle. These hypotheses connect peculiar features near Earth's core-mantle boundary to subduction of the oceanic lithosphere through the deep carbon cycle. Depending on the melt fraction and wetting angle, Liu et al. (2016) distinguished three types of dynamics of the Fe-C melt in ULVZs that could have different seismic signatures, and may explain the laterally varying electrical conductivity in the D'' layer: (a) ULVZs containing solid phases that have become iron-rich through reaction with Fe-C melt; (b) ULVZs containing 9-16 vol% non-wetting Fe-C melt coexisting with a small degree of silicate melt in a solid silicate matrix; (c) ULVZs containing 5-11 vol% Fe-C melt that wets the solid silicate matrix and will ultimately merge with the core, explaining some of the W isotopic signature sampled by ocean island basalts (Rizo et al., 2019).

The seismic features of the hot and dense LLSVPs, including seismic anisotropy, are generally consistent with the presence of bridgmanite, CaSiO₃-perovskite (or Davemaoite; Thomson et al., 2019). They have recently appeared to be also consistent with the high-pressure polymorphs of CaC₂O₅ (Wang et al., 2024), with a structure that is based on CO₄ groups (tetracarbonates) with sp³-hybridized carbon. In particular the symmetry change from orthorhombic Fdd2 to monoclinic C2 leads to a density increase of 5.8% and a Vₛ decrease of 7.4%, and to elastic parameters of CaC₂O₅-C2 consistent with seismic observations within LLSVPs. The orthorhombic to monoclinic transition in CaC₂O₅ occurs at 80 GPa, broadly consistent with the top of LLSVP piles. More generally, the elastic properties of the polymorphs of CaC₂O₅ that may form by the reaction between CaCO₃ and CO₂ in deep subducting slabs from transition zone conditions and deeper are in good agreement with the Primary Reference Earth Model. This would connect the deep carbon cycle from the oceanic crust to the lowermost mantle and back to ocean island basalts. Given similarities in elastic properties between CaSiO₃-perovskite and CaC₂O₅-C2, it may be challenging to evaluate the actual amount of oxidized carbon present in LLSVPs. Their fate may be better estimated through their isotopic signature in suites of ocean island basalts.

Should there be remnants of the basal magma ocean, sp³ hybridized tetracarbonates are less siderophile, and thus large amounts of it may be stranded instead in the deep mantle. These tetracarbonates may therefore provide a possible source for carbon-rich emissions registered at the surface in suites of ocean island basalts (Cerantola et al., 2023).

Carbon throughout the Earth's mantle

Carbon is a trace element in the Earth's convecting mantle. Estimates of the carbon content in the bulk silicate Earth generally reconstructed from noble gas content are estimated to vary between 500 and 1000 ppm (Marty et al., 2013, 2016). Ratios of C to elements that partition in a similar manner in the mantle, such as Ba or Nb, yield a BSE estimate of 110 ± 40 ppm for the mantle, and 140 ± 40 ppm for the BSE (Hirschmann, 2018). When estimated from elements that behave geochemically like carbon but are refractory in a cosmochemical sense, concentrations of ~192 ppm and ~350 ppm C are derived from U and Ba respectively (Lee et al., 2019).

The concentration of carbon in the lower and upper mantle is substantially different. Measured carbon fluxes near hot spots are much higher than those for the depleted mid-ocean ridge basalt (MORB)-source mantle indicative of a higher concentration of carbon in the lower mantle. The most comprehensive and detailed analysis of the carbon content in MORBs showed very-fine-scale variations in their mantle source, with carbon fluxes from the mantle to the surface varying by over three orders of magnitude at the segment scale along mid-ocean ridges (Hauri et al., 2019). Integration over the entire length of ridges yields a global carbon flux in the form of CO₂ of 1.3–1.5 × 10¹² mol yr⁻¹ (16–18 Mt. C yr⁻¹) and a mean content of 100 ppm C in the depleted mantle source of MORB (Cottrell et al., 2019, Hauri et al., 2019; Le Voyer et al., 2019).

Measurement of carbon solubility in the major minerals of the upper mantle, transition zone and lower mantle shows that carbon is virtually insoluble in the silicate minerals, at the level of a ppm or less (Keppler et al., 2003; Shcheka et al., 2006). This implicitly requires that carbon is stored in accessory, nominally carbon-rich phases such as elemental carbon in the form of graphite or diamond, carbides, hydrocarbons in regions rich in hydrogen or carbonates, with oxidation numbers ranging between -4 and +4 depending on pressure, temperature and oxygen fugacity (fO₂). As such, the diversity of carbon-rich phases is diagnostic of the vertical and lateral heterogeneities in the convecting mantle. Over geological time, the forms of carbon-rich phases are controlled by the degree of oxygenation of the mantle and are diagnostic of its evolution through time.
## Reduced lower mantle and transition zone

Rock samples from depths of 300 km and below are scarce; the deeper, the scarcer. Observations rely almost exclusively on inclusions within the few super-deep diamonds that have been collected through time in cratons (see Section "Carbon movement within and between deep reservoirs"). Observations of natural samples are complemented by experimental and theoretical studies to assess the redox state and carbon-hosting phases in the deep mantle.

The deepest natural messengers from the mantle are the 'CLIPPIR' diamonds (Cullinan-like, Large, Inclusion Poor, Pure, Irregular, and Resorbed) that formed at lower mantle conditions and were brought back to the surface (see Shirey et al., 2013, 2019, and references therein). Their most common inclusion is a composite metallic Fe-Ni-C-S mixture. The metallic inclusions are comprised of cohenite (Fe,Ni)₃C, interstitial Fe-Ni alloy, segregations of Fe-sulfide, and minor occasional Fe-Cr-oxide, Fe-oxide, and Fe-phosphate. A thin fluid layer of CH₄ and lesser H₂ is trapped at the interface between the inclusion and surrounding host diamond. The mixture is commonly interpreted to have been trapped as a molten metallic liquid. Their depth is bracketed to depth between 360 and 750 km by inclusions containing high-pressure silicate minerals such as retrogressed CaSiO₃-perovskite (Davemaoïte) or majoritic garnet.

The Fe-Ni-C-S inclusions in CLIPPIR diamonds are unique physical samples of an Fe-rich metallic liquid from the deep mantle, which confirm the fundamental process of Fe²⁺ disproportionation at depth, also called 'redox freezing'. Indeed, the Earth's mantle oxygen fugacity is buffered by ferric-ferrous (Fe³⁺-Fe²⁺) equilibria and decreases with increasing depth. This has been measured in natural garnet peridotites (Woodland and Koch, 2003), and confirmed by experiments that tested garnet oxythermobarometry equilibria under relevant high-pressure conditions (Stagno et al., 2013, and references therein). Theoretical analysis of the mantle fO₂ evolution with increasing depth (Ballhaus, 1995; Ballhaus and Ronald Frost, 1994) predicts that an isochemical mantle composition becomes reduced by ~0.6 log units in fO₂ per GPa as a result of molar volume changes of redox equilibria and changing mantle mineralogy. At lower mantle conditions, experimental studies have shown that aluminous bridgmanite and majoritic garnets are sinks for ferric iron, which in turns forces ferrous iron disproportionation into ferric silicates and Fe-metal (e.g., McCammon, 1997; Frost and McCammon, 2008). This is thought to generate up to about 1 wt% metal in the lower mantle and establishes a carbon-reducing environment in much of the deep mantle, where the forms of carbon would be diamond, metal and carbides (Fig. 3).

Experimental work has also assessed the relevant carbon phases in the lower mantle. At the base of the upper mantle, the Fe-Ni-C solidus at 10 GPa is low (1150–1250 °C) compared to mantle geotherms, which may indicate that Fe-Ni-C alloys are stable as molten phases, potentially limited in size to the grain scale. Cohenite [Fe,Ni]₃C observed in the CLIPPIR diamonds likely forms together with metal and diamond at ~10 GPa or 350 km depth at intermediate Ni# between 0.11 and 0.24 (Rohrbach et al., 2014). In the deeper mantle, where carbon is at least locally more abundant, carbides in the form of molten Fe₂C₃ along with diamond are expected to be the phases hosting carbon (Dasgupta and Hirschmann, 2010). The presence of 100–200 ppm of sulfur prevents the formation of carbides and favors the precipitation of diamonds, thus explaining the Fe-Ni-S inclusions in the CLIPPIR diamonds. Whether such small fractions of melt and/or carbides affect seismic attenuation and Vp and Vs remains an open question.

Some super-deep diamonds show a wide range of carbon isotope compositions extending to very low values (e.g., δ¹³C ~0‰ to -25‰), which are interpreted as the result of the contribution of crustal material (see Section "Carbon movement within and between deep reservoirs"; Howell et al., 2020). Others, like the super-deep diamonds from Kankan (Guinea) and their inclusions, have very narrow multi-isotope signatures that are typical of a mantle that has experienced little or no crustal interaction. The latter isotopic signature could be explained by a slight hydration of the uppermost lower mantle by a hydrated subducted oceanic lithosphere that destabilizes carbides to form diamond and would be consistent with a lower-mantle barrier for carbon subduction (Regier et al., 2020). Upon infiltration into the deep mantle, oxidized carbonate minerals are submitted to 'redox freezing' and reduced to diamond ± carbides as described above; whether carbonate rocks are preserved and enter the deep mantle through subduction therefore remains controversial and might actually display some significant lateral variability within the mantle.

A few inclusions hosted in super-deep diamonds have also shown the presence of solid carbonate minerals, including calcite, magnesite and ferromagnesite (Brenker et al., 2007; Kaminsky et al., 2016; see Section "Diamonds and their solid carbon inclusions" for additional details), suggesting that oxidized carbon in the form of carbonate minerals could be stable, at least locally. A recent estimate of redox state at the conditions of the transition zone was proposed based on the ferric iron content of majorite inclusions in diamonds (Kiseeva et al., 2018), which increased linearly as a function of pressure and is likely due to the oxidizing effect of CO₂-rich fluids rather than Fe²⁺ disproportionation. This would correspond to an fO₂ value about 2 log units above the iron-wüstite buffer.

Observations of carbonate inclusions and their potential local stability in the deep mantle have been complemented by a wealth of experimental and theoretical studies investigating the stability of mineral endmembers (i.e., CaCO₃, MgCO₃, FeCO₃) and their solid solutions at high pressures and temperatures, with the aim of assessing the fate of subducted oxidized carbon that may be potentially transferred deep in the mantle (see reviews by Boulard et al., 2020; Merlini et al., 2020). At Earth's surface, carbonate minerals mainly occur as calcite (CaCO₃) and dolomite (CaMg(CO₃)₂). In the lower mantle, they react with silicate minerals such as pyroxenes and bridgmanite to form ferromagnesite ([Mg,Fe]CO₃), which is considered the dominant carbonate phase in the deep mantle. Independent of their reactivity toward silicate minerals, oxides, or reduced carbon, all carbonate minerals are stabilized at the conditions of Earth's lower mantle through a series of high-pressure phase transitions; no fewer than 20 new carbonate structures have been discovered so far (Merlini et al., 2020). The main feature occurs at mid-lower mantle conditions, where new crystal structures including tetrahedrally coordinated sp³-hybridized carbon form a systematic class of carbonates, nesocarbonates, cyclocarbonates, and inocarbonates, with complex crystal chemistries analogous to silicate structures.
In the mid-lower mantle, the spin crossover of Fe²⁺ from high spin to low spin changes the physical and chemical properties of iron in minerals (Badro, 2014, and references therein), including Fe-bearing carbonates (e.g., Fu et al., 2017; Liu et al., 2014; Mattila et al., 2007). As the expected changes in the elastic parameters across the transition are large, there has been hope that this would help to detect seismically ferromagnesite in the deep lower mantle. Unfortunately, both the iron content and the fraction of magnesioferrite that would produce a detectable 1% shear velocity decrease compared to non-carbonated lithologies is exceedingly high (Chariton et al., 2020).

The carbonate phase that is most likely to survive in the mid-lower mantle and deeper (>1800 km, 80–90 GPa) is CaCO₃. Carbonate minerals entrained in cold slabs in the form of magnesite or dolomite react with silicate CaSiO₃ to form CaCO₃ with sp³-hybridized carbon (CaCO₃-P21/c-h; Lv et al., 2021). The presence of such carbonates is unlikely to be directly detected, as the change in cation within a carbonate mineral would not be observable seismically. However, CaCO₃ produced in the deep mantle would carry a higher calcium isotopic composition (δ⁴⁴/⁴⁰Ca), distinct from the much lower Ca isotope signature of surface-derived CaCO₃. If preserved in diamond inclusions and returned to the surface, ⁴⁴Ca-enriched CaCO₃ could be used to trace the presence of oxidized carbon in the lowermost mantle.

## The carbon-depleted oxidized upper mantle

The upper mantle gradually became more oxidized by 1.5 log units between ~3.0 Ga and 1.9 Ga when fO₂ values reached levels near the FMQ buffer, as recorded by mid-ocean ridge basalts and metabasalts over geological time (Aulbach and Stagno, 2016). Their redox is largely controlled by the initial ferric iron content of the mantle and the geotherm, as elemental carbon turns into carbonates while ferric iron is reduced to ferrous iron during redox melting (see Stagno and Fei, 2020).

Therefore, below approximately 90 km in the asthenosphere, the oxidized form of carbon is carbonate, either as dolomite or melt depending on the thermal regime. Evidence of the carbonate forms present in the mantle is provided by solid carbonate inclusions within diamond (see Section 5.5.3) with Mg-rich and Ca-rich dolomite stable in peridotitic and eclogitic assemblages, respectively (Falloon and Green, 1989; Hammouda, 2003). As silicate minerals react with this carbonate, the latter tend to be mostly in the melt phase. At depths greater than approximately 150–180 km, the asthenospheric mantle is too reducing for carbon to stay in its oxidized form and only graphite or diamond is present, as discussed above. Hence, the region located in the depth range of 90 km to 150–180 km is where carbonatitic melts are most likely to be produced and subsequently impregnate the surrounding mantle through metasomatism, fostering movement of carbon in the upper mantle (see Section “Movements of deep carbon in melts”, and review by Hammouda and Keshav, 2015). The presence and composition of dolomite in the upper mantle are critical for the composition of carbonatitic melts.

When in its elemental zero oxidation state, the upper stability limit of diamond is at 150–170 km depth for the present-day mantle geotherm, and diamond is replaced by graphite upon ascent and vice-versa upon descent, though with a slow kinetics. If sufficient hydrogen is present, C-O-H fluids containing CO₂, CH₄, and more complex ionic species in variable proportions (Sverjensky and Huang, 2015; Sverjensky et al., 2020) may also be present and facilitate graphite or diamond crystallization (e.g., Holloway, 1984; Luque del Villar et al., 1998; Stachel et al., 2017).

## Sub-continental lithospheric mantle

The sub-continental lithospheric mantle (SCLM) acts as a sluice gate between the Earth's interior and exosphere. Substantial volumes of carbon and other volatile elements have accumulated in the SCLM over long periods of geological time before they are liberated during continental rifting and magmatism (Gibson et al., 2020; Gibson and McKenzie, 2023). Direct evidence for carbon accumulation in the SCLM exists in the form of carbon-rich mantle xenoliths, which are observed at numerous sites around the globe representative of thick continental crust and lithosphere (Chen et al., 2018; Ducea et al., 2005), and at continental rifts where thick continental lithosphere is tectonically ruptured (Casagli et al., 2017; Rooney et al., 2017; Trestrail et al., 2017). Such xenoliths appear to be neither compromised in their carbon concentrations during their ascent, nor reflective of preservation bias (e.g., Hirschmann, 2018). Additionally, alkaline low-SiO₂ silicate melts encountered in juvenile continental rift settings require low degrees of melting in environments with abundant CO₂ concentrations, further betraying the nature of their sub-lithospheric source region (Foley et al., 2009).

Foley and Fischer (2017) outline three principal means of sequestering carbon in the SCLM. Firstly, some carbon is emplaced during the initial formation of the SCLM. In this case, it may be assumed that the cratonic lithospheric mantle formed through the accretion of island arcs (e.g., Lee, 2006); the initial cratonic lithosphere is therefore reflective of sub-arc mantle wedge lithosphere and will therefore host a similar concentration of carbon.

Secondly, gradual addition of small fraction carbon-rich asthenospheric melts that ascend into the lithosphere from the underlying asthenosphere will passively contribute toward the SCLM carbon budget. At the base of the lithosphere, the temperature is generally lower than in the convective asthenosphere. Small-fraction low-SiO₂, high-CO₂ mantle melts are therefore stable at these conditions, and can rise and freeze, causing so-called mantle metasomatism (Aulbach et al., 2017, 2020; Grégoire et al., 2000; Hirschmann, 2010; Massuyeau et al., 2021; McKenzie, 1989; Yaxley et al., 1998; O'Reilly and Griffin, 2010). Small fractions of carbonate melt derived from the convecting mantle may end up frozen within layers near the bottom of the lithosphere as a result of their low volumes and the reduced conditions in the SCLM (McKenzie, 1989; Sleep, 2009), or metasomatize the SCLM itself. Metasomatism therefore re-enriches depleted lithosphere by reintroducing volatiles and incompatible elements. This process is broad enough that it is documented in both cratonic and oceanic lithosphere, and it has been shown to produce H₂O- and CO₂-bearing phases such as carbonates. Mantle xenoliths from Ethiopia and Tanzania appear to have been altered by carbonate
fluids in this manner (Rudnick et al., 1993; Trestrail et al., 2017). Sometimes, lithospheric metasomatism results in mantle rejuvenation; in other words, the base of the lithospheric lid is transformed into asthenospheric convective domains (Aulbach et al., 2017; Foley, 2008, 2011; Liu et al., 2019; Tappe et al., 2007).

A third, more active means of sequestering mantle carbon is through the episodic passage of mantle plumes in the underlying asthenosphere. In contrast to gradual accumulation of carbon in the SCLM, plume-related enrichment would deliver carbon derived from enriched mantle sources during lithospheric metasomatism (Foley et al., 2009). However, the high temperatures of convective plumes may partially destabilize fusible metasomes in the lithosphere, thereby remobilizing volatiles within the SCLM (Rooney et al., 2014) as opposed to emplacing them. Further carbon may be deposited in the SCLM from the diapirism of carbonates delivered into the mantle through subduction (Ducea et al., 2022). Xenoliths from the West Antarctic Rift demonstrate that multiple discrete periods of metasomatism may be recorded and correlated with different SCLM refertilization events, including subduction (Broadley et al., 2016). Similarly, the metasomatism of xenoliths from the Kaapvaal craton has been attributed to subduction-derived fluids (e.g., Jackson and Gibson, 2023).

Because of these multiple contributing factors that may sequester carbon in the SCLM, literature estimates of SCLM carbon content are both broad and uncertain. Further uncertainty is introduced through the necessity to estimate the mass of the SCLM, which is combined with SCLM carbon concentrations to assess total SCLM carbon budgets. The widest range in the literature is provided by assuming fixed bulk mantle volumes and carbon concentrations representative of mantle lithosphere (3.5–350 × 10²¹ g C, corresponding to 100–1000 ppm C; Kelemen and Manning, 2015). Following the three principal carbon addition methods listed in the previous paragraphs (initial sequestering at formation; gradual addition; plume addition), Foley and Fischer (2017) suggest an initial SCLM carbon budget of 3.2 × 10²¹ g C, which has increased by an order of magnitude to 37.5 × 10²¹ g C over geological time through the gradual addition of carbon from small-fraction melts and plumes, corresponding to an SCLM mantle carbon concentration of 860 ppm. However, Hirschmann (2018) argues that this value is unrealistic owing to the high carbon concentrations required by the infiltrating plume magmas of Foley and Fischer (>10 wt% C in silicate melts, necessitating a convecting mantle of ~2 wt% CO₂; c.f. 100–1000 ppm, Section “Carbon throughout the Earth’s mantle”).

The largest single estimate of 60 × 10²¹ g C is derived from assumptions made about the frequency of kimberlite emplacement at depth within the SCLM (Sleep, 2009). In contrast, bulk rock analyses of SCLM xenoliths suggest a value closer to 10 ± 5 × 10²¹ g C (100 ± 20 ppm; Hirschmann, 2018), which is similar to a recent combined volumetric-geochemical estimate of 7 ± 3 × 10²¹ g C at a concentration of 140 ± 56 ppm (Gibson and McKenzie, 2023). From further subdivision of the carbon within lithospheric mantle, Gibson and McKenzie estimate that cratonic mantle contains ≥45% of the total lithospheric carbon budget under continents and oceans; off-craton mantle, which is more prone to rifting, will therefore have substantially lower carbon concentrations (93 ± 61 ppm versus 158 ± 74 ppm) and lower carbon residence time than cratonic mantle. This vast amount of sub-lithospheric carbon would be sequestered in the lower cratonic lithosphere as diamonds in sub-continental peridotites and eclogites owing to their highly depleted and refractory compositions (Foley, 2011; Section “Diamonds as clues on deep and super-deep carbon movements”), and/or in hydrous dike assemblages comprising carbonate, phlogopite, and clinopyroxene (Foley and Fischer, 2017). Amphibole-bearing lithospheric metasomes may also host substantial carbon, as evidenced by xenoliths (e.g., Rooney et al., 2017). Such carbonated lithologies may contribute toward seismic discontinuities observed within cratonic lithosphere (Saha et al., 2021).

As carbon is an incompatible element in the mantle, the SCLM will be devolatilized during continental rifting (Gonzalez and Gorczyk, 2017; Gorczyk and Gonzalez, 2019). Changes in temperature and oxidation state, and reactions with infiltrating fluids and melts will result in the destabilization of carbon-bearing phases in the SCLM and the mobilization of carbon, which will migrate upwards toward the rift surface (Foley and Fischer, 2017). Redox melting is likely to dominate in the earliest stages of continental rifting. Significant oxidation of the SCLM, possible through the infiltration of small-fraction asthenospheric melts, will result in the oxidation of diamond into carbonate, thereby mobilizing carbon into low-SiO₂ carbonate-rich melts such as carbonatites or nephelinites (Foley, 2011). At later stages of rifting, carbon will depress the peridotite solidus and facilitate mantle melting (Section “Movements of deep carbon in melts”; e.g., Dasgupta and Hirschmann, 2006). Decompression of the upwelling asthenosphere under thinning rifting lithosphere, especially in the presence of a mantle plume, will result in the formation of silicate melts which will overshadow the low-volume, low-fraction, volatile-rich melts generated in the earlier stages of rifting. The heat supplied from lithospheric thinning and asthenospheric melts will destabilize carbon-bearing phases present in the SCLM (Gorczyk and Gonzalez, 2019), forcing an upwards-migration of carbon into the continental crust. The gradual focusing of melts under rift zones over time into a narrow volcano-tectonic locus (Brune et al., 2023) will concentrate carbon derived from a broad area into a narrow region under the rift.

During the mechanical process of rifting, advection of unstable cratonic keel material may concentrate the carbon mobilized toward continental rifts, further increasing fluxes of CO₂ seen at rift zones (Muirhead et al., 2020). Cratonic cores themselves do not degas CO₂ significantly, in contrast to the rift zones that border them (Muirhead et al., 2020). As cratonic mantle can contain ≥45% of the overall lithospheric carbon budget (Gibson and McKenzie, 2023), mobilization of cratonic carbon in this manner may substantially enhance the quantity of carbon feeding into rift zones. Through this mechanism, carbon-rich magmas such as carbonatites may be generated without necessitating a carbon-enriched mantle source (Fischer et al., 2009).

Silicate melts generated during decompression melting at continental rifts will ascend through the SCLM and pond at or within the crust (Section “Movements of deep carbon in melts”). At this stage, melts may interact with crustal carbonate minerals and sources of biogenic carbon, resulting in further CO₂ release through thermal decarbonation (e.g., Mason et al., 2017). The solubility of CO₂ in magmas will decrease with decreasing pressure (Dixon et al., 1995), thereby resulting in extensive degassing of magmatic
bodies within the continental crust. At active continental rift zones and rift volcanoes, vast degrees of diffuse and active carbon degassing are characterized worldwide (Brune et al., 2017; Cappelli et al., 2023; Foley and Fischer, 2017; Hunt et al., 2017; Lee et al., 2016; Tamburello et al., 2018; Wong et al., 2019), with carbon isotope ratios marking a strong affinity with a mantle-derived magmatic source (e.g., Blomgren et al., 2019; Bräuer et al., 2018; Correale et al., 2019; Frondini et al., 2008; Muirhead et al., 2020). The difference in CO₂ concentrations between undegassed crustal melt inclusions and parental mantle-derived magmas (determined through CO₂-incompatible trace element ratios) is sufficient to account for the considerable fluxes of diffuse CO₂ degassing at continental rifts (Iddon and Edmonds, 2020; Wong et al., 2023).

## Sub-oceanic lithospheric mantle

Carbon may also be sequestered within sub-oceanic lithospheric mantle. At mid-ocean ridges, asthenospheric carbon mobilized by mantle melting is focused toward the ridge axis, where a significant quantity degasses into the water column. In addition, a portion of the mobilized carbon is directed toward the flanks of the spreading ridge, where it is sequestered through the metasomatism of the juvenile sub-oceanic lithospheric mantle (Keller et al., 2017). As oceanic plates age, their underlying mantle lithosphere will thicken, and by doing so it will accumulate additional small-fraction carbon-rich asthenospheric melts in the developing lithosphere. Furthermore, xenoliths collected from the base of the Pacific plate suggest that later extensional stresses in the lifetime of the oceanic plate such as plate flexure may induce further metasomatism (Pilet et al., 2016). Carbon present in the oceanic lithosphere in this manner is believed to contribute to observed seismic low-velocity zones and velocity contrasts (Keller et al., 2017). The controls on CO₂ sequestration within oceanic lithosphere include mantle temperature, mantle CO₂ content, and melt productivity/ fertility (Keller et al., 2017), which are known to vary substantially across the present-day ridge system (e.g., Le Voyer et al., 2019). In combination with the difficulty in accessing the lower oceanic lithosphere, there is likely to be substantial uncertainty in quantifying the carbon budget of the lithosphere underlying oceanic plates. Following geodynamic models and tomographic estimates of lithospheric mass, recent first order estimates suggest that 0.5–2.2 × 10²¹ g C may be stored within present-day unaltered sub-oceanic lithospheric mantle (Gibson and McKenzie, 2023; Müller et al., 2022).

## Deep carbon in the crust

In the continental crust, solid carbon will take multiple different forms depending on the pressure, geothermal gradient, and oxidation state of the local environment (e.g., Hazen et al., 2013a; Oganov et al., 2013). Voluminous solid continental carbon phases can be broadly subdivided into two principal forms. The largest fraction of crustal carbon is represented by inorganic carbon in carbonate minerals, in which oxidized carbon is present within the bicarbonate anion (CO₃²⁻) bonded in mineral structures to a wide selection of cations, which account for about 80% of total crustal carbon (Berner, 2004; Hayes and Waldbauer, 2006). The second largest form of solid crustal carbon, accounting for about 20% of total carbon, is organic matter, in which carbon is present in its reduced form (Hazen et al., 2013a; see chapter in this book by Derry, 2024). It includes biogenic organic matter, petrogenic or detrital graphitic carbon in sedimentary rocks, and abiotic forms of condensed reduced carbon (Beyssac and Rumble, 2014; Galy et al., 2008; Luque del Villar et al., 1998).

A large fraction of crustal solid carbon, both organic and inorganic, in the crust formed through biological processes and their successive solid-state transformations. Carbonate minerals can be synthesized by various organisms that grow a carbonate shell, or result from biology-mediated chemical reactions, as discussed in Section “Origins of deep carbon” above. Over the past 150 million years, deposition of calcareous plankton in deep-sea sediments, mainly coccoliths and foraminifera, has increased and currently accounts alone for more than half of global carbonate sedimentation (Berner, 2004; Milliman, 1974). Organic carbon accumulates in rocks during sedimentary processes and is progressively transformed to carbonaceous materials and graphitic carbon by diagenesis and metamorphic processes (see discussion above). Carbonate minerals and organic matter can also form abiotically in the crust, through multiple processes ranging from silicate weathering at the Earth’s surface, to deep crustal fluid-rock interactions. The following sections describe the forms and distribution of carbon in the main crustal reservoirs.

### Continental crust

The continental crust is an important reservoir of carbon. Recent estimates derived from statistical analysis of lithological distributions suggest that ~20% of present-day continental crust is composed of carbonate rocks (Walton and Shorttle, 2024), which operate as the largest store of crustal carbon at the present-day (Hirschmann, 2018). However, the distributions and forms of carbon on and within present-day and geologically historic continental crust is highly heterogeneous, complicating our capacity to estimate the carbon budget of the crust. Here, we summarize the current understanding of carbon reservoirs in the continental crust.

Crustal carbonate and organic matter are distributed within both the sedimentary and crystalline materials that comprise the bulk continental crust (Hirschmann, 2018). Of these two carbon forms, carbonate is perhaps more commonly analyzed, while data on reduced carbon is sparse in comparison (Derry, 2014; Hartmann et al., 2012; see chapter in the current Treatise, Derry (2024)). Carbon in the continental crust dates back to 4.0 Ga, coinciding with some of the earliest records of life on the planet (Schielowski, 1988). Primordial Earth had a primarily silicate crust, which has been gradually carbonated over time as a consequence of continuous deep-Earth degassing (Alcott et al., 2024). From present-day records of lithological preservation, it is estimated that ~5% of total preserved Precambrian crustal volume comprises carbonate rocks, which are principally biogenic stromatolitic and abiotic precipitated carbonate rocks deposited in oceanic environments and preserved on continental crust (Walton and Shorttle, 2024). After the Cambrian and Ordovician radiations of biomineralizing marine life (Gilbert et al., 2022), there is a substantial
increase in the proportion of carbonate preserved in continental crust (~20–25%); similarly, a decrease in the fraction of accumulated carbonate minerals is recorded within crustal sections in the aftermath of the Permian–Triassic extinction (Walton and Shorttle, 2024). This ongoing carbonate accumulation within Earth’s continental crust has had substantial consequences for the co-evolution of life and the geosphere and atmosphere (e.g., Alcott et al., 2024; Duncan and Dasgupta, 2017, Isson et al., 2020). The abundance of carbonate minerals present with the bulk Earth crust is therefore both affected by and in turn affects surface biogeochemical cycles (see Section “Carbon at the interface between the geosphere and biosphere”).

In the continental lithosphere (crust and uppermost non-convecting mantle), the principal form is also carbonate minerals (Hazen et al., 2013a; Oganov et al., 2013). Of these minerals, the most abundant carbonate minerals in the Earth's crust are calcite (β-CaCO₃) and dolomite (CaMg(CO₃)₂), which are present within massive sedimentary and metamorphic carbonate formations that account for 90% of crustal carbonate minerals (Reeder, 1983). Calcite is additionally one of the most common cement minerals found within siliciclastic sedimentary rocks (Manning et al., 2013). Aragonite (λ-CaCO₃), the orthorhombic polymorph of the trigonal calcite, is a metastable form of carbonate which principally originates from biomineralization, and is therefore found in substantial proportions on continental shelves where it is sourced from the dissolved bicarbonate anion in seawater by marine biomineralizers (e.g., Gilbert et al., 2022). Other key but minor carbonate minerals include magnesite (MgCO₃), which forms through the carbonation of Mg-rich igneous and metamorphic rocks or through authigenic precipitation (e.g., Scheller et al., 2021), and siderite (FeCO₃), a principal component of Precambrian banded iron formations (e.g., Klein, 2005). At the time of writing, 614 carbonate minerals are registered by the online mineral database mindat.org, n.d. (i.e., filtering by 'CO₃²⁻'; retrieved June 2024), illustrating the broad and diverse nature of carbonate minerals on Earth at the present day (Hazen et al., 2013b). Carbonate minerals are also generated through chemical weathering of silicate rocks, which consumes atmospheric CO₂ (Gaillardet et al., 1999; Hartmann et al., 2012; Stewart et al., 2019; Urey, 1952) and has hence been correlated to long-term climate modulation (e.g., Gernon et al., 2021; Johansson et al., 2018).

Oxidized carbon in crystalline continental basement is present in igneous and metamorphic rocks as carbonate minerals or as fluid inclusions (e.g., Lowenstern, 2001; Skelton, 2011). In magmatic systems, mobile carbon, present as a vapor phase or dissolved within silicate magmas, typically takes the form of CO₂ (typically 1–50 mol%), with its more reduced gaseous forms CO and CH₄ much lower in concentration (typically ≪1 mol%; Giggenbach, 1996; Symonds et al., 1994). While carbon concentrations are enhanced in peralkaline/alkaline compositions (Ni and Keppler, 2013), CO₂ is highly insoluble in silicate magmas (e.g., Dixon et al., 1995), and will degas deep within the Earth's crust to form fluid phases (Section “Carbon reprecipitation at depth”). The majority of this CO₂ is expected to be degassed into the atmosphere or oceans, although magmatic carbon-bearing fluids, among other carbon sources, may contribute to the formation of hydrothermal solids (e.g., Bénézeth et al., 2013). An additional minor reservoir of crustal carbon is in rare low-melt-fraction volcanic rocks, such as carbonatites, nephelinites, and kimberlites (e.g., Jones et al., 2013).

In near-surface environments, reduced forms of carbon (i.e., C⁰ and hydrocarbons) are attributed to biological processes, and are hence commonly referred to as simply 'organic carbon' (Derry, 2014). This carbon is present almost exclusively within sediments, such as deep-sea/lacustrine shales or coals, and is derived from organic material resistant to microbial processing (e.g., microbial membranes or plant lignin; Manning et al., 2013). Burial of organic material in marine environments is the second largest sink of atmospheric CO₂ after silicate weathering and carbonate mineral precipitation; the content of organic carbon in marine sediments is enhanced by fluvial transport of organic carbon from continents to oceans (Galy et al., 2008). At higher crustal pressures, reduced and dehydrated carbon in sediments takes the form of graphite (Hazen et al., 2013a), which is a key contributor to the electrical conductivity of continental crust (Glover, 1996). Graphite may also form within some igneous and metamorphic rocks as an accessory phase precipitated from reduced C-O-H fluids (e.g., Luque del Villar et al., 1998). Under specific circumstances, organic carbon in the crust will undergo thermal maturation to generate biogenically-derived hydrocarbons (e.g., Sephton and Hazen, 2013). Further thermal maturation will result in the decarboxylation and dehydrogenation of organic compounds, resulting in the gradual oxidation and aromatisation of organic hydrocarbons to conclude in graphite (see Section “Deep abiotic organic matter” for a full discussion on deep abiotic organics).

Carbon in the continental crust will be remobilized during metamorphism in various forms such as CO₂, CH₄, or more complex ionic species (e.g., Manning et al., 2013; Sverjensky et al., 2020; Section “Carbon movement within and between deep reservoirs”). Destabilization of solid carbon during metamorphism may result from a number of tectonic or volcanic processes that result in disequilibrium pressure-temperature conditions, including volcanism (e.g., Black and Gibson, 2019; Ganino and Arndt, 2009; Johnston et al., 2011; Mason et al., 2017), localized microbial and thermogenic breakdown to form CH₄ (Section “Metamorphic and magmatic processes of carbon mobilization”; e.g., Etiope et al., 2019; Etiope and Sherwood Lollar, 2013), and regional and contact metamorphism and metasomatism (Section “Metamorphic and magmatic processes of carbon mobilization”; Aiuppa et al., 2017; Ague, 2000; Ague and Nicolescu, 2014; Becker et al., 2008; Capriolo et al., 2021; Groppo et al., 2013, 2017; Halama and Bebout, 2021; Kerrick and Caldeira, 1998; Stewart and Ague, 2018; Svensen et al., 2004). As a result, the lengths of tectonic settings, in particular those of arcs and subduction zones, are believed to correlate with periods of icehouse-greenhouse activity on Earth (McKenzie et al., 2016; Gernon et al., 2022; Pall et al., 2018). Retrograde metamorphism may in turn consume mobile carbon in fluids to precipitate minerals such as carbonates and/or graphite (Section “Carbon reprecipitation at depth”; e.g., Hu et al., 2021; Piccoli et al., 2016, 2018, 2021; Scambelluri et al., 2016). Finally, solid carbon in carbonates will be mobilized by processes such as oceanic and atmospheric acidification (e.g., Doney et al., 2009; Eyssautier-Chuine et al., 2016) and reverse weathering (Isson and Planavsky, 2018).
Much like other elements in the continental crust, the concentration of carbon is assessed by determining the volumetric quantities of each lithology in the crust, and then further assigning each lithology with a carbon concentration characteristic of representative rocks exposed at the Earth's surface (see current Treatise chapter Derry, 2024, for a full review). In the context of continental carbon, sampling biases may introduce additional uncertainty: the uppermost crust is easiest to access, yet is also most prone to chemical weathering, which will increase carbon content over time as a result of secondary carbonate formation (Hartmann et al., 2012). In addition, both oxidized and reduced forms of carbon must be considered to provide a complete picture of continental carbon budgets. In the bulk continental crust, the proportion of reduced carbon relative to the total carbon budget is considered to range from 0.10 to 0.28, depending on the choice and proportions of characteristic crustal lithologies selected (Derry, 2014; Gao et al., 1998; Hartmann et al., 2012; Wedepohl, 1995; Yaroshevsky, 2006); for example, sediments typically have a greater proportion of reduced carbon than crystalline basement (e.g., Yaroshevsky, 2006).

A recent summary of the present-day carbon budget of the continental crust is provided by Hirschmann (2018). Literature values assessed by Hirschmann (2018) converge on values of $76 \pm 16 \times 10^{21}$ g C hosted in sedimentary rocks and $15 \pm 8 \times 10^{21}$ g C within crystalline rocks for a total of $91 \times 10^{21}$ g C, and is based on previous estimates of crustal carbon budgets (e.g., Gao et al., 1998; Hayes and Waldbauer, 2006; Hunt, 1972; Wedepohl, 1995; Yaroshevsky, 2006). This budget is distributed across the solid exosphere, including carbon in oceanic domains (see following section), at a concentration of $27 \pm 5$ ppm (Hirschmann, 2018).

## Oceanic crust and upper lithospheric mantle

At the present day, the total mass of carbon in oceanic crust and lithosphere is on the order of $\sim 1.0–1.5 \times 10^{22}$ g C (Hirschmann, 2018; Müller et al., 2022). The majority of this carbon ($\sim 60\%$) is stored within pelagic sediments, with the remainder comprising carbonate minerals in the oceanic basement and sub-oceanic lithospheric mantle (Hirschmann, 2018; Müller et al., 2022). Because seawater-rock interactions in the ocean do not only affect crustal sequences but also the uppermost part of the oceanic lithospheric mantle, this section presents an overview of the upper oceanic lithosphere affected by magmatic and hydrothermal processes at mid-ocean ridges.

### Altered lithospheric mantle

Seawater percolates through permeable uppermost oceanic crust and lithospheric mantle where it may be exposed, e.g., at mid-ocean ridges or through faulting (e.g., Fisher, 2005). Carbonate minerals can form during serpentinization through the coupled hydration and carbonation of olivine to form serpentine minerals and carbonates (e.g., Power et al., 2013; Kelemen et al., 2011; Menzel et al., 2018). Ultramafic rocks can therefore have significant potential for rapid carbonate formation upon reaction with $CO_2$-bearing fluids (e.g., Kelemen et al., 2011; Kelemen and Matter, 2008; Matter and Kelemen, 2009). Significant natural carbon sequestration is expected where mantle peridotite is exposed to seawater at oceanic settings and subject to carbonation (Alt et al., 2013). Serpentinization and coupled carbonation of the oceanic lithosphere may therefore occur via a number of different tectonic processes: (a) widespread exposure and faulting of peridotite at (ultra-)slow spreading centers (Alt et al., 2013; Cannat et al., 2010; Dick et al., 2003; Kelemen et al., 2011; Merdith et al., 2020; White et al., 2001) and locally at the ridge axis at faster spreading centers and oceanic transform zones (Alt et al., 2013; Klein et al., 2024); (b) exposure of continental lithospheric mantle within magma-poor continent-ocean transitions (Albers et al., 2021; Grevemeyer et al., 2022; Liu et al., 2023; Sawyer et al., 1994; Schwarzenbach et al., 2013); c) faulting of oceanic plates during bending at the outer rise as they enter subduction zones (van Avendonk et al., 2011; Christensen and Ruff, 1988; Faccenda, 2014; Grevemeyer et al., 2018; Lefeldt et al., 2012; Ranero et al., 2003).

There is a higher concentration of carbon near exposed serpentinite seafloor as a consequence of seawater accessibility. While uppermost oceanic serpentinite is primarily inorganic carbon, deeper serpentinite ($>50–100$ m) principally hosts organic carbon, which is sourced from methanogenic microorganisms within serpentinites that thrive on the hydrogen gas byproduct of serpentinization (Schwarzenbach et al., 2013) and abiotic forms of organic matter (e.g., Sforna et al., 2018; Ménez, 2020). Most data on carbon concentrations in altered peridotite are obtained from seafloor drill samples, with the remainder determined from ophiolites (Kelemen and Manning, 2015). These altered peridotites have an average bulk concentration of $681 \pm 45$ ppm C (Kelemen and Manning, 2015), although concentrations from some settings may be as high as 17.1 wt% $CO_2$ (Albers et al., 2021). Carbonated serpentinites therefore host a minor yet not insignificant proportion of the overall oceanic lithosphere carbon budget, estimated as $\sim 0.2–0.3 \times 10^{21}$ g C (Merdith et al., 2020; Müller et al., 2022).

### Altered oceanic crust

In the high-permeability upper crust, carbon, sourced from magmatic fluids and seawater, precipitates as the carbonate minerals calcite and aragonite within rocks and veins at low temperatures ($<100$ °C; Alt and Teagle, 1999). Most basaltic crust alteration will occur within the first few tens of millions of years after crust formation (Alt and Teagle, 1999; Coogan and Dosso, 2015; Coogan and Gillis, 2018; Coogan et al., 2016). The combination of pore infilling through carbonate mineral precipitation and gradual accumulation of sediment cover will limit subsequent circulation of water through oceanic crust over geological time, thereby decreasing carbon sequestration as the crust ages (Grevemeyer et al., 1999); overlying carbonate sediments may themselves be incorporated into the basaltic crust as recrystallized carbonate minerals to facilitate this process (Gillis and Coogan, 2011). High degrees of basalt alteration will also limit the availability of exposed fresh material required to precipitate carbonate minerals, thereby limiting further carbonate precipitation in aged crust (Albers et al., 2023). Another control on crustal carbonate formation is
bottom-water temperature, which promotes faster fluid-rock reactions resulting in carbonate precipitation (Coogan and Gillis, 2013; Gillis and Coogan, 2011).

Estimates of bulk oceanic crustal CO₂ concentration suggest a range of 0.18–0.22 wt% CO₂ (Alt and Teagle, 1999; Kelemen and Manning, 2015). Most CO₂ is concentrated in the uppermost 300 m of oceanic crust at an estimated mean concentration of 1.85 wt% (Müller and Dutkiewicz, 2018), which represents two-thirds of the total CO₂ in typically ~7 km-thick oceanic crust (Gillis and Coogan, 2011; Staudigel et al., 1989). Estimates of present-day oceanic crust carbon storage using these concentration estimates range from 2 to 4 × 10²¹ g C (Hirschmann, 2018; Kelemen and Manning, 2015; Müller and Dutkiewicz, 2018). At least a fraction of this carbon is present in the form of organic compounds (Shilobreeva et al., 2011).

## Sediments

The key parameters to consider when assessing the carbon budget of pelagic oceanic sediments are the mass of present-day sediment and the amount of carbon present within said sediment. Both parameters are highly heterogeneous in the modern oceans (Conrad, 2013; Dutkiewicz et al., 2016; Hayes et al., 2021), and are assessed respectively through seafloor seismic reflection surveys and direct sampling through the ocean research drilling programs (Plank, 2014; Plank and Langmuir, 1998; Rea and Ruff, 1996). The principal controls on the thickness of sediment at any given point on the seafloor are (a) the distance of the point from the nearest passive continental margin, a measure of the sedimentary input to that point, and (b) the age of the oceanic crust, a measure of time over which sediment accumulation has occurred (Dutkiewicz et al., 2017; Olson et al., 2016). Large rivers will accelerate local oceanic sedimentation rates; for example, the Ganges river fan possesses a column of sediment of 7 km thickness at its maximum (Curray, 2014).

Carbon-bearing sediments can be subdivided into carbonate (predominantly CaCO₃) and organic carbon (e.g., Clift, 2017). Accumulation of significant thicknesses of oceanic sedimentary carbonate is a geological novelty on Earth (e.g., Gilbert et al., 2022). While metazoan carbonate biomineralization is known from the Ediacaran onwards (e.g., Grant, 1990; Wood, 2011), the onset of extensive neritic biomineralization during the Marine Mesozoic Resolution was the principal herald of widespread and significant marine carbonate production representative of present-day carbonate sedimentation in the oceans (Erba, 2006; Falkowski et al., 2004; Ridgwell, 2005).

Sedimentary carbon accumulation on oceanic crust is limited by the carbonate/calcite compensation depth (CCD), which marks the depth below which the rate of inorganic carbonate dissolution exceeds the rate of deposition (e.g., Broecker, 2008; Derry, 2022; Dutkiewicz et al., 2019; Ridgwell and Zeebe, 2005; van Andel, 1975). Within present-day oceans, the CCD corresponds to 4.5–5.0 km below sea level (Dutkiewicz et al., 2016; Zeebe, 2012). Terrestrial sediment fluxes (Derry, 2022; Walling and Fang, 2003), biological productivity (Kumar et al., 1995; Sarmiento et al., 2004), and sea-surface temperature and salinity (Dutkiewicz et al., 2016), all act in concert to affect carbonate deposition or dissolution, thereby controlling the position of the CCD. The parameterization of these factors permits the extrapolation of carbonate sediment thicknesses to all present-day ocean basins (Dutkiewicz et al., 2019). Furthermore, the thermal subsidence of ocean basins with age will mean that old oceanic crust will carry negligible sedimentary carbonate minerals, in contrast to shallow seafloor in regions of high biological productivity (Dutkiewicz et al., 2016, 2019; Plank, 2014). On the other hand, organic carbon is rapidly consumed by biological processes such as respiration (e.g., Higgins et al., 2009). Preservation of organic carbon in sediments therefore requires rapid deposition and burial, for example, beneath regions of high biological productivity or at deep sea fans where sedimentation rates are higher (Galy et al., 2007; Plank and Manning, 2019). Temperature during sediment burial additionally has a strong effect on organic carbon sediment preservation (Malinverno and Martinez, 2015). Finally, the weathering of silicate minerals and organic carbon in pelagic sediments will fix substantial CO₂ into carbonate minerals over geological time (Wallmann et al., 2008). At the present day, the mass of carbon within sediments is estimated to be on the order of 10 × 10²¹ g C (Dutkiewicz et al., 2019; Hirschmann, 2018; Yaroshevsky, 2006); pelagic sediments are therefore the largest reservoir of carbon on oceanic plates at the present day.

## Deep abiotic organic matter

Carbon dioxide is generally considered as the dominant carbon species in the modern Earth (Fig. 1). However, as conditions become more reducing at depth, CH₄ and H₂, together with smaller hydrocarbons, become stable at the expense of CO₂ or other more oxidized carbon species within the stability field of graphite or diamond (Sverjensky et al., 2020). In the mantle, the transition from CO₂-dominated to CH₄-dominated regimes occurs at depths of 125–140 km (~4 GPa) in sub-continental domains, and at shallower depths of ~90 km (~2.5 GPa) in the sub-oceanic mantle. The chemical complexity of the fluid organics that could be present is exemplified by the wealth of organic matter found in the fluid inclusions of diamonds from the Urals placers from ~200 km depth, though their exact origin remains unknown (see Table 4 in Sobolev et al., 2019). They contain aliphatic hydrocarbons up to C17, cyclic hydrocarbons up to C15, oxygenated hydrocarbons, and heterocyclic compounds, together with nitrogenated and sulfonated compounds (plus CO₂ and H₂O). This is supported by experimental and theoretical work. Above 2 GPa, heating pure CH₄ at high temperature triggers its polymerization to higher alkanes (C3), with ethane the most abundant species (Kolesnikov et al., 2009). Theoretical simulations have shown that formic acid is thermodynamically more stable than the products of the water-gas shift reaction (CO₂ and H₂) above 3 GPa and at ~750–1150 °C (Stolte et al., 2021). This result is in good agreement with the spectroscopic signature of carboxylic functional groups coating diamond inclusions in the deeply subducted metamorphic rocks of the Lago di Cignana in the Western Alps at P ≥ 3.2 GPa and T ~ 600 °C (Frezzotti, 2019), and with thermodynamic models suggesting that organic acids play an important role in diamond nucleation and growth in deep Earth
(Sverjensky et al., 2014; Sverjensky and Huang, 2015). Similarly, but at shallower depths, the graphite or other condensed carbon forms present in many crustal and mantle rocks such as those described at mid-ocean ridges of within subduction contexts — and now emplaced and incorporated in the continental crust — may not be of biotic origin (i.e., transformation of subducted biogenic organic matter), but instead has precipitated from deep carbonic fluids (e.g., Boutier et al., 2024b; Debret et al., 2022; Evans et al., 2002; Luque del Villar et al., 1998; Ménez et al., 2018) or from the interaction between reduced fluids and carbonates (Galvez et al., 2013; Malvoisin et al., 2012; Peng et al., 2021; Tao et al., 2018; Vitale Brovarone et al., 2017).

Expanding on the carboxylic acids, experimental work has shown that pressure expands miscibility gaps in C-O-H fluids, and immiscible fluids are generated at pressures in excess of 1.5 GPa (600–700 °C), with formic acid yielding both aqueous and dry methane and ethane fluids ±CO₂ and ±H₂ depending on the oxygen fugacity (Li, 2017). At higher pressures (up to 4.6 GPa) and lower temperatures (300–350 °C), aqueous Na- or Ca-acetate yields aqueous and mineral carbonate (as expected from the speciation of C(IV) at high pressure; Facq et al., 2014), soluble light alkanes (C1–C3), and immiscible hydrocarbons; the higher the pressure, the greater the amount of hydrocarbons and methane generated (see Huang et al., 2017; Huang et al., 2023 for details). At 350 °C and above, immiscible hydrocarbons tend to become aromatic, and may transform into polycyclic aromatic hydrocarbons and ultimately into graphite, depending on temperature and composition. Natural fluid inclusions from high-pressure metamorphic rocks support the immiscibility of CH₄, and potentially the immiscibility of other hydrocarbons in aqueous fluids at high pressure conditions (Giuntoli et al., 2024; Sverjensky et al., 2020; Vitale Brovarone et al., 2017).

As of today, the complexity of the fluid organics observed in deep fluid inclusions has been partly reproduced in experiments (McCollom, 2013), emphasizing that highly mobile and likely highly reactive hydrocarbons are stable at sub-lithospheric conditions, provided that the temperature is low enough to kinetically prevent the formation of graphite or diamond, as their compositions are mainly controlled by the local water activity and oxygen fugacity.

While lithospheric subsurface environments should be the exclusive realm of CO₂ as discussed in previous sections, there are numerous occurrences of CH₄ reported in vents, seeps, springs and aquifers at slow-spreading mid-ocean ridges (Abrajano et al., 1988; Andreani et al., 2023; Charlou et al., 2010; Etiope and Sherwood Lollar, 2013; Grozeva et al., 2020; Klein et al., 2019; Young et al., 2017), in ophiolites or peridotitic massifs and metamorphic complexes (Boutier et al., 2021, 2024b; Harada and Tsujimori, 2024; Herviou et al., 2021; Hu et al., 2021; Peng et al., 2021; Shi et al., 2005; Tao et al., 2018; Vitale Brovarone et al., 2017, 2020b), in alkaline igneous complexes (Potter and Longstaffe, 2007; Salvi and Williams-Jones, 1997), and in Precambrian basements (Etiope and Sherwood Lollar, 2013; Sherwood Lollar et al., 2008; Warr et al., 2021). Although CH₄ in these settings may be related to biotic processes such as thermogenic degassing, in many of these settings, reduced conditions are achieved through local fluid-rock interaction, in particular via aqueous alteration of Fe²⁺-rich rocks, the most famous examples being the flames of the Chimera vents in Turkey and 'Los Fuegos Eternos' in the Philippines on continents, or black and white smokers at mid-oceanic ridges related to the process of olivine hydroxylation (serpentinization). Other localized anomalous deviances from lithospheric oxidation state are related to radiolytic dissociation of water in Precambrian basements (Sherwood Lollar et al., 2014), or the composition of fluids associated with alkaline intrusions (Potter and Longstaffe, 2007; Salvi and Williams-Jones, 1997). Methane in these settings is thought to form abiotically, even though successive biological reworking is likely (Young, 2019).

Natural occurrences of recognized abiotic organic volatiles of hydrothermal origin also include short-chain alkanes, and formic acid, the smallest organic acid. They have mainly been observed in geothermal systems or continental seeps within ophiolites and Precambrian shields, and at hydrothermal vents near mid-ocean ridges and subduction forearms. These small abiotic organics have attracted a lot of attention over the last decades, in particular those formed at low temperature (below 400 °C), as they may provide a potential route to the emergence of life and/or support subsurface life (see Section "Carbon at the interface between the geosphere and biosphere").

The effectiveness of abiotic organic volatile generation at low temperatures remains a subject of controversy and debate. The mechanism of abiotic CH₄ synthesis is commonly assumed to be Fischer-Tropsch type reactions (FTT) between an oxidized carbon molecule (CO₂) and H₂, which is naturally produced by the serpentinization reaction occurring in ultrabasic rocks. As the ferrous iron of olivine (or pyroxene) is oxidized to form ferric hydrous minerals, water is reduced to form H₂, with catalytic support from Ni-, Fe-rich minerals. This H₂ in turn provides the ingredients for the FTT reaction. However, CH₄ formation is kinetically hindered at temperatures below 300–400 °C (McCollom, 2013). Understanding the detailed mechanism of the highly energetic H₂ and eventually CH₄ production by FTT reaction is of tremendous importance, as the serpentinization reaction is ubiquitous on Earth and likely on other planets, when hot water is in contact with ultramafic rocks, and might support life on other planetary bodies. The challenges are (i) the effectiveness of FTT reactions in an aqueous medium, as the Fischer-Tropsch reaction is restricted in its definition to the gaseous phase, (ii) the availability of CO that is the reactant form of C, (iii) the identification of a catalyst phase, and (iv) the potential distribution in mass of the expected alkane products according to the Shultz-Flory distribution.

Tens of experimental serpentinization studies, performed over decades, have therefore been attempted to identify the parameters that control the abiotic production of H₂, the subsequent FTT formation of CH₄, and eventually light hydrocarbons. A recent network analysis of >30 experimental studies (see Barbier et al., 2020; Huang et al., 2021a, Huang et al., 2021b for details) achieved an unbiased reading and interpretation of the literature and confirmed that H₂ is systematically and significantly formed in experiments, indicating that its production during serpentinization is well established. However, little or virtually no CH₄ is being produced in agreement with the kinetic/thermal control on CH₄ formation. The carbon budget of some experiments yielding high CH₄ production indicated some contamination, in total agreement with the very few studies that included labeled ¹³C in the carbon source (Grozeva et al., 2017; McCollom et al., 2010). The FTT reaction is very unlikely to be the route to produce the large methane emissions observed in near-surface ultramafic bodies, unless most of the CH₄ present at those settings is related to hotter
and longer water-rock interactions in fluid inclusions (Andreani et al., 2023; Klein et al., 2019; McDermott et al., 2015). Similarly, abiotic production of CH₄ and potentially other light hydrocarbons in serpentinizing ultramafic rocks may be kinetically favored at higher pressure and temperature conditions, such as at subduction zones (Boutier et al., 2021; Vitale Brovarone et al., 2020a).

In continental ultramafic systems (ophiolites and peridotite massifs), Etiope and Whiticar (2019) reevaluated the origin of methane by taking into account the latest theory, experiments and analytical methods, e.g., CH₄ isotopologues, applied to a larger set of field data on vents, seeps and springs. They converged on a model of abiotic methane generation based the Sabatier reaction, which consists in a catalytic CO₂ hydrogenation in a water-free (or unsaturated) rock, at relatively low temperature, with metal-rich catalysts such as the magnetite or chromite found in many ultramafic rocks.

For oceanic contexts, the reader is referred to the recent detailed evaluation of abiotic organics in hydrothermally altered mantle-derived rocks of Andreani and Ménez (2019). In short, CH₄ observed at mid-oceanic deep-sea vents likely results from the entrapment and re-speciation at high temperature (>400 °C) of mantle-derived CO₂ within fluid inclusions or vesicles in mantle and magmatic rocks (e.g., Andreani et al., 2023; Klein et al., 2019; McDermott et al., 2015), the fluids later being released during hydrothermal processing of the oceanic rocks. Moreover, in the large amount of total organic carbon measured in some serpentinized abyssal peridotites, a fraction is very likely present as solid carbonaceous matter, potentially including an abiotic contribution. As a consequence, understanding the formation and diversity of abiotic organic material produced during the hydrothermal alteration of ultramafic rocks requires a paradigm shift, no longer centered on FIT reactions. One includes the potential (and possibly metastable) formation of organic carbon compounds with intermediate oxidation states, mediated by rock-forming minerals during hydrothermal reactions (see Seewald et al., 2006; Fig. 15.4 in Andreani and Ménez, 2019); such compounds may include formic acid, formaldehyde, methanol, methane, ethane, ethene, benzene and graphite and all the polycyclic aromatic hydrocarbons in between, which form from successive reversible redox, dehydrogenation and hydrogenation reactions. At temperatures below 400 °C, where the formation of methane and graphite is kinetically inhibited, a large variety of highly volatile and mobile single-carbon organic species can then become available in addition to CO₂ species in fluids within natural serpentinizing systems of highly variable pH and local H₂ levels. As of today, the liquid and solid forms of abiotic organic matter have often received much less attention in both experimental and natural studies. Occurrence of organic minerals related to hydrothermal alteration in mafic and ultramafic rocks has been seldom reported, e.g., karpatite (C₂₄H₁₂), which has the highest molecular weight among natural organic minerals (see review of polycyclic aromatic hydrocarbon minerals by Echigo and Kimata, 2010). Recent high-resolution investigations of closed micro-environments void of biotic contamination in fresh or metamorphic hydrothermally altered ultramafic samples highlighted that the presence of hydrothermally derived abiotic organic carbon trapped within the uppermost lithosphere occurs often as heavy and aromatic compounds that are chemically and structurally diverse and closely associated to phyllosilicates (e.g., Andreani et al., 2023; Ménez et al., 2018; Sforna et al., 2018). Similarly, as anticipated above, graphite in many metamorphic rocks may be of abiotic origin and result from organic geochemical pathways deviating from conventional metamorphic petrology assumptions (Boutier et al., 2024b; Debret et al., 2022).

Carbon movement within and between deep reservoirs

The global carbon cycle results from the interplay and interconnections among deep and surface processes, and the response time of various forms of carbon to their timescales (Lee et al., 2019). Deep carbon on Earth is mobilized through multiple processes such as tectonics, metamorphism, magmatism, or a combination of them, and may involve movements of solid or fluid carbon species. This section introduces the main plate tectonics processes mobilizing solid carbon and the related solid carbon fluxes, and then summarizes the main processes of carbon mobilization from solids to fluids and the associated fluxes of deep carbon in fluids.

Movements of carbon from the crust to the deep Earth

Subduction of crustal carbon

Subduction, the tectonic mechanism of returning oceanic lithosphere into the Earth at convergent settings, is the principal input of exospheric carbon into the deep Earth (Fig. 4). Processes resulting in the sequestration of carbon in oceanic plates will determine the quantity of carbon returned into the Earth's interior, which at the present day is predicted to be 82 ± 14 Mt C yr⁻¹ (where 1 Mt. = 10¹² g; Plank and Manning, 2019), alternatively 68–140 Mt. C yr⁻¹ (Müller et al., 2022) or 107 ± 26 Mt. C yr⁻¹ (see Section “Secular variation of surficial versus deep carbon reservoirs or steady state?”). The largest fraction of this carbon is represented by carbonate minerals, which account for about 80% of total subducted carbon (Plank and Manning, 2019, and references therein). The second largest input of solid carbon into subduction, accounting for about 20% of total carbon inputs, is reduced carbon (organic matter; graphitic carbon; graphite). The contribution of surface/subsurface hydrocarbons to deep carbon recycling is less constrained owing to the question as to if these compounds are transferred to depth at convergent margins or mobilized at shallow depths. Nevertheless, some experimental studies have assessed the potential transformation and stability of subducted hydrocarbons at convergent margins (Serovaiskii et al., 2019).

The overarching subducting carbon budget may be evaluated by characterizing the carbon concentrations within the different individual lithological components comprising the subducting slab that can be sampled, and combining these concentrations with the kinematic parameters of subduction (e.g., Dasgupta and Hirschmann, 2010; Kelemen and Manning, 2015; Plank and Manning, 2019). An adjacent methodology is to parameterize the controls on global carbon subduction at the present-day, and apply these
<figure><img src="image_6.png" /><figcaption>**Fig. 4** Schematic diagram illustrating the budgets and fluxes of carbon within the near-surface deep carbon cycle. The continental lithosphere hosts ~100 × 10²¹ g carbon (Hirschmann, 2018; Gibson and McKenzie, 2023), which is obtained from enrichment of the sub-continental lithospheric mantle (Foley and Fischer, 2017) and weathering of continental crust (Gaillardet et al., 1999), and lost to the atmosphere through orogenic outgassing (Becker et al., 2008; Stewart and Ague, 2018; Stewart et al., 2019), reverse weathering (Isson and Planavsky, 2018; Isson et al., 2020), and magmatism. Volcanic outgassing at intraplate (rift and plume) and arc settings contributes to the mobilization of crustal and lithospheric carbon in continental settings (Bekaert et al., 2021; Brune et al., 2017; Hauri et al., 2019; Werner et al., 2019; Wong et al., 2019), which is enhanced through a number of additional unquantified lithospheric processes (Gorczyk and Gonzalez, 2019; Mason et al., 2017; Muirhead et al., 2020). In oceanic settings, the carbon budget is estimated to be ~10–15 × 10²¹ g (Hirschmann, 2018; Müller et al., 2022), which is subdivided between carbon-bearing sediments (Cartapanis et al., 2018; Clift, 2017; Dutkiewicz et al., 2019; Hayes et al., 2021; Hirschmann, 2018; Yaroshevsky, 2006), altered oceanic crust (Kelemen and Manning, 2015; Li et al., 2019; Merdith et al., 2019; Müller and Dutkiewicz, 2018; Müller et al., 2022), serpentinized peridotites (Alt et al., 2013; Merdith et al., 2020; Müller et al., 2022; Schwarzenbach et al., 2013), and sub-oceanic lithospheric mantle (Gibson and McKenzie, 2023; Keller et al., 2017; Müller et al., 2022). Carbon in oceanic lithosphere is obtained at mid-ocean ridges (Keller et al., 2017) and through weathering of pelagic sediments and oceanic crust (Coogan and Gillis, 2018; Wallmann et al., 2008); additional carbon is introduced into the oceans through mid-ocean ridge volcanism (Hauri et al., 2019; Le Voyer et al., 2019; Müller et al., 2022; Tucker et al., 2018). During subduction, the slab is decarbonated (Kelemen and Manning, 2015; Müller et al., 2022), dissolved (Ague and Nicolescu, 2014; Frezzotti et al., 2011; Piccoli et al., 2021), or reduced (Malvoisin et al., 2012; Galvez et al., 2013; Vitale Brovarone et al., 2017; Peng et al., 2021), and carbon is sequestered into the forearc (Barry et al., 2019; Debret et al., 2022; Fullerton et al., 2021) and in arc magmas (Johnston et al., 2011). Carbon exchange between continents and downgoing slabs may also occur as a result of subduction erosion and accretion (Clift and Vannucchi, 2004; Straub et al., 2020). Organic carbon burial across both oceans and continents is likely to balance organic carbon oxidation (Kump and Arthur, 1999). 10¹² g C is equivalent to 1 Mt C; 10²¹ g C is equivalent to 10⁹ Mt C, or 10⁶ Gt C.</figcaption></figure>

parameterizations to plate tectonic reconstruction models (e.g., Müller et al., 2022; Wong et al., 2019). As seafloor/subseafloor alteration processes account for a large fraction of carbon storage in the oceanic lithosphere, reviews of carbon ingassing at subduction zones traditionally categorize inputs into three principal lithological components, which are characterized by observations from ocean research drilling programs (e.g., Staudigel et al., 1989; Kelemen and Manning, 2015; Plank and Manning, 2019): carbon in serpentinized lithospheric mantle (Section “Altered lithospheric mantle”), basaltic oceanic crust (Section “Altered oceanic crust”), and carbon in the sediment overlying the basaltic crust (Section “Sediments”). The carbon sequestered within these subducting reservoirs in the lithosphere is sourced from the ocean itself, as the percolation of seawater into permeable oceanic lithosphere is the principal means of carbon deposition within the crust and uppermost mantle (Section 4.4.2). Contributions to the sedimentary component of subducting carbon come primarily from marine organisms in the form of carbonate shells and reduced organic carbon remains (Clift, 2017), although there is significant input from terrestrial sediments, predominantly at river estuaries and deltas where sedimentation rates are substantial (e.g., Dutkiewicz et al., 2016).

In addition to these three traditional reservoirs of subducting carbon, recent work has highlighted that a non-negligible fraction of carbon can be present in oceanic lithospheric mantle sections unaffected by seafloor/subseafloor alteration (see Section “Sub-oceanic lithospheric mantle”; Gibson and McKenzie, 2023; Keller et al., 2017; Müller et al., 2022). This sequestered carbon in unaltered oceanic lithosphere may be remobilized upon entering the Earth at subduction zones. However, as with quantifying the mass of the ocean lithospheric budget, there is substantial uncertainty as a result of the multiple controls on carbon emplacement at mid-ocean ridges (Keller et al., 2017). By selecting representative values for these multiple parameters, it is believed that sequestered carbon in the unaltered oceanic lithosphere may contribute 5–21 Mt C yr⁻¹ to the overall subducting carbon flux at the present day (Merdith et al., 2019; Müller et al., 2022).
Owing to the low production rate of oceanic peridotite, exposed oceanic lithosphere generated at magma-poor rifts and mid-ocean ridges has previously not been considered as a significant contributor toward the budget of subducting carbon (1-4% of all present-day oceanic crust, contributing ~1.3 Mt. C yr⁻¹; Alt et al., 2013; Kelemen and Manning, 2015). However, more recent estimates suggest that carbonated serpentinite from (ultra-)slow spreading ridges may contribute ~5-13 Mt. C yr⁻¹ at the present day (Müller et al., 2022). Serpentinites formed during plate flexure are also believed to be substantial contributors of carbon owing to the formation of lithosphere-scale faulting and serpentinization (e.g., Ranero et al., 2003), itself a function of slab dip (Faccenda, 2014). Serpentinite carbonation as a result of outer-rise bending is therefore expected to contribute a substantial amount to the subducting carbon budget, although this value remains very unconstrained (0-22 Mt. C yr⁻¹; Kelemen and Manning, 2015; Müller et al., 2022). Instead, carbon uptake by basaltic oceanic crust as a result of weathering and alteration is believed to be a greater store of carbon at the mid-ocean ridge environment, contributing ~20 Mt. C yr⁻¹ toward total fluxes of subducted carbon (Kelemen and Manning, 2015; Li et al., 2019; Merdith et al., 2019; Müller et al., 2022; Müller and Dutkiewicz, 2018).

Quantifying the subduction flux of carbon-bearing sediments is also complex. As established in Section 4.4.2.3, factors such as the age of subducting plate, distance from continents, and biological productivity will directly control the amount of sedimentary carbon accumulated on the plate (Dutkiewicz et al., 2016, 2019; Plank, 2014). Furthermore, sediment subduction may be hindered by accretion on the overriding plate, which appears to be substantial at slow subduction rates (<6-7 cm yr⁻¹; Clift and Vannucchi, 2004). If considering subduction fluxes over geological time, prior to the Marine Mesozoic Revolution, sedimentary carbonate is not considered to be a significant contributor to subduction zones (Edmond and Huh, 2003; Müller et al., 2022), although the radiative evolution of biomineralizing metazoans from the Cambrian onwards may have significantly affected the exospheric carbon cycle in the absence of carbonate sediment subduction (Walton and Shorttle, 2024). Similarly, the fluxes of subducted organic carbon have changed over geological time following major variations of surface organic carbon production such as the Cambrian explosion (Giuliani et al., 2022). In the modern oceans, organic carbon constitutes a mass fraction ranging from 0.5 to 1.5 wt% in carbonate-free oceanic sediments (Malinverno and Martinez, 2015). However, the quantities of subducted organic carbon through geological time are currently unknown and are assumed to constitute a very minor fraction of total subducting carbon through geological time, reflecting values typical of deep sea sediments distal to continents (<0.5 wt%; Müller et al., 2022). Nevertheless, recent work has shown that latitude-dependent variations in the compositions of subducted sediments, possibly related to latitudinal variations in the primary production of phytoplankton and organic matter burial, may impart a strong control on the redox state of arc magmas (Hu et al., 2024). The GLOSS-II composition, representative of 'average' subducted sediment, has 3.07 ± 0.23 wt% carbon, but in practice sedimentary compositions will be substantially variable depending on depth and environment of deposition (Plank, 2014). Furthermore, different subduction systems will subduct variable proportions of carbonate to organic carbon (Clift, 2017). At the present day, seafloor sediment is believed to be the greatest contributor to the overall subducted carbon budget, of which ~20% is organic carbon (~60 Mt. C yr⁻¹; Clift, 2017; Müller et al., 2022; Wong et al., 2019).

## Subduction erosion

In contrast (but not antithetical) to the process of sediment accretion on continental crust is the erosion of the upper plate forearc during subduction (e.g., Clift and Vannucchi, 2004; Stern, 2011, 2020; Straub et al., 2020), which is believed to be the most significant process destroying continental crust (e.g., Clift et al., 2009b,a). The principal evidence for continental erosion arises from the subsidence of some convergent margin forearc (e.g., Ranero and von Huene, 2000; Vannucchi et al., 2001). In addition, despite potential for blending with geochemical signatures for slab components and crustal contaminants (e.g., Straub et al., 2020), it is possible to identify signatures of eroded crust in arc magmas (e.g., Stern, 2020; Straub et al., 2015). Estimates of present-day subduction erosion typically range from 30 to 150 km³ Myr⁻¹ per km length of subduction zone (e.g., Clift, 2017; Straub et al., 2020; Vannucchi et al., 2016); it is therefore likely that erosion of the overriding plate can contribute toward the carbon budget of subduction. However, it is difficult to determine the contributions of this contribution beyond qualitative assessment. It is known that the rate of material transfer during erosion depends on several factors, including the topography and sediment cover of the subducting plate, high convergence rates, and subduction channel fluid dynamics (Clift and Vannucchi, 2004; Straub et al., 2020; Vannucchi et al., 2012). Furthermore, eroded convergent margins typically have overlying plates exposing igneous arc basement, which bears less carbon than other lithologies (Clift, 2017; Yaroshevsky, 2006); the lithology of the overriding plate will be a further contributor to the overall carbon budget of subduction erosion. Further characterization of the subduction erosion process, and therefore also of volume and mass fluxes of eroded overlying crust at subduction zones, will be required in order to completely capture the quantities of carbon contributed to subduction from this process.

## Movements of carbon from depth to the crust

### Tectonic movements of solid carbon from the deep Earth to the surface

At convergent margins, tectonic or mechanical movements of solid carbon from mantle depths to the crust or within the crust are primarily controlled by off-scraping or diapirism of downgoing slab material or eroded mantle wedge material (see reviews by Klein and Behn, 2021; Plank and Manning, 2019). Carbon-bearing subducted materials, especially subducted sediments but also large volumes of subducted oceanic igneous and mantle rocks, can be removed from the downgoing slab, reducing the amount of subducted carbon reaching the convective mantle. A dominant fraction of this material is removed from the downgoing slab at trenches to form accretionary prisms (Tewksbury-Christle et al., 2021), At greater depths within the forearc to arc region, the fate of
this material is uncertain; subducted sediments may be removed from the slab to form diapirs that partially melt or do not melt as they rise into the mantle wedge, and be emplaced at the base of the crust (Behn et al., 2011; Ducea et al., 2022; Kelemen and Behn, 2016), or be exhumed as tectonic slices or mélanges along the plate boundary to form orogenic complexes (Agard et al., 2009; Guillot et al., 2009). The fluxes of solid carbon associated with these tectonic or mechanical movements of deep rocks toward the surface are largely unknown. It has been suggested that off-scraping of sediments at trenches may remove a fraction of subducted sediments as high as 35% of global subducted sediments (Straub et al., 2020). Nevertheless, the ultimate amount of solid carbon being mobilized through tectonic or mechanical processes varies from one subduction zone to another, or even subduction zone sub-segments relative to adjacent sub-segments, based on the specific tectonic setting, the type and amounts of subducted sediments, and the thermal regimes controlling the contextual mobilization of carbon through metamorphic or magmatic degassing (Plank and Manning, 2019).

## Metamorphic and magmatic processes of carbon mobilization

The mobilization of carbon from solids to fluids requires changing conditions and is strongly dependent upon the closed or open system behavior of the considered systems. Rock chemistry, changes in pressure and/or temperature conditions, changing redox, and/or on the presence and nature of fluids are the main parameters controlling the mobilization of carbon from the main solid carbon phases, namely carbonate minerals, organic matter, diamond (Fig. 5), and fluids and melts. Such conditions can be achieved through multiple processes and at various geological contexts. These processes are summarized in the following sections.

### Mobilization of carbon from carbonate minerals

The three main processes mobilizing carbon from carbonate minerals are decarbonation reactions, carbonate dissolution, and melting. In the forearc to subarc regions of subduction zones as well as in the continental crust, decarbonation and dissolution are the most important processes, whereas carbonate melting is expected to take place at greater depths in the mantle, or during anatexis in the lower crust.

Decarbonation is the process of carbonate breakdown through carbonate-silicate reactions in response to rising temperature. Also referred to as 'Urey reactions' (Urey, 1952), decarbonation can be summarized by the generic reaction:

$$MCO_3 + SiO_2 \rightarrow MSiO_3 + CO_2$$

where M represents a divalent cation, most typically $Ca^{2+}$ or $Mg^{2+}$.

Examples of decarbonation reactions are common in metamorphic rocks in metasedimentary, metamafic, and ultramafic rocks, and involve a broad range of carbonate (calcite, dolomite, ankerite) and silicate (quartz, serpentine, and many others) reactants (Baumgartner and Valley, 2001; Epstein et al., 2020; Ferry, 1991; Cook-Kollars et al., 2014; Collins et al., 2015; Stewart et al., 2019; Vitale Brovarone et al., 2018; Fig. 6a). Decarbonation is effective in regional and contact metamorphic settings. Besides the characteristic calc-silicate products, decarbonation reactions are also identified by shifts in the carbon isotopic composition of the residual carbonate fraction; residual carbonate is $^{12}$C-enriched due to the positive $\Delta$Carb-CO$_2$ (e.g., Baumgartner and Valley, 2001; Bottinga, 1969; Epstein et al., 2020; Cook-Kollars et al., 2014; Collins et al., 2015). Decarbonation reactions do not require the presence of an aqueous fluid to produce carbon-bearing fluids. Nevertheless, the presence of aqueous fluids, for example during infiltration of external fluids or in response to local dehydration reactions, has been shown to enhance decarbonation reactions (Ferry, 1991; Gorman et al., 2006; Menzel et al., 2020; Vitale Brovarone et al., 2018). Decarbonation is believed to contribute substantially to metamorphic carbon degassing at convergent margins (Gorman et al., 2006; Kelemen and Manning, 2015; Kerrick and Connolly, 2001), with collisional metamorphism producing higher CO$_2$ fluxes relative to subduction metamorphism, and cumulative fluxes as high as about 6–84 Mt C yr$^{-1}$ (Stewart et al., 2019; Fig. 4). Experimental studies and thermodynamic modeling results predict decarbonation reactions and the formation of CO$_2$-enriched fluids to be particularly efficient at subarc depths in subducting slabs (Gorman et al., 2006; Gonzalez et al., 2016). At shallower depths, contact metamorphism at convergent margins may also substantially contribute to carbon degassing (Aiuppa et al., 2017).

In addition to decarbonation reactions, carbonate dissolution also contributes to the mobilization of deep carbon from solid to fluid phases. Carbonate dissolution has only recently been added to the inventory of processes contributing to deep carbon budgets and fluxes (Ague and Nicolescu, 2014; Facq et al., 2014; Frezzotti et al., 2011; Kelemen and Manning, 2015). The lack of previous assessments was related to the only recent establishment of values for the dielectric constant of water at high-pressure conditions (Sverjensky et al., 2014). Carbonate dissolution may be effective at mobilizing carbon into aqueous fluids in the absence of silicate minerals, i.e., when decarbonation reactions cannot take place. Carbonate solubility is dependent upon pressure, temperature, and redox conditions (Fig. 6b,c). Experimental studies on aragonite solubility demonstrated that, under oxidized conditions, the speciation of carbonic fluids produced by carbonate dissolution is strongly dependent upon the pressure and temperature conditions, with bicarbonate ions dominating at pressure < 3 GPa to at least 400 °C and carbonate ions at higher pressure and temperature conditions (Facq et al., 2014). CO$_2$ was found to be present in appreciable amounts below 2–3 GPa (Li, 2017). Reduced conditions, instead, have been shown to boost carbonate mineral solubility in aqueous fluids, with methane controlling the speciation of carbonic aqueous fluids resulting from carbonate dissolution (Lazar et al., 2014) (Fig. 6c). Natural examples of carbonate dissolution are documented in metamorphic terranes such as the Cyclades (Ague and Nicolescu, 2014) and Corsica (Piccoli et al., 2016, 2021).
<figure><img src="image_7.png" /><figcaption>**Fig. 5** Photomicrographs of natural deep carbon forms. (a) Olivine-bearing dolomite marble, Hida belt, Japan (Harada and Tsujimori, 2024). Olivine formation indicates the reaction between silicates and dolomite to produce olivine + CO₂ with increasing temperature. (b) Calcite crystals being partially converted to graphite during the infiltration of reduced fluids. Note the preferential reaction along grain boundaries and carbonate twin planes and cleavage. Samples from the Belvidere Mountain Complex, Northern Vermont (Boutier et al., 2024b). (c) Graphitic carbon in metapelitic rocks from the Wepawaug Schist, Connecticut, USA (Ague, 2003; Zhang et al., 2018a,b). Based on carbon stable isotopes, the graphitic carbon is interpreted to represent metamorphosed biogenic organic material (Zhang et al., 2018a,b). (d) Hydrocarbon oil and gas inclusions in hydrothermal quartz from Porretta Terme, Northern Apennines, Italy (Olivieri and Miglioli, 2021). The hydrocarbon oils and gases are interpreted to have formed through thermogenic processes during burial and heating of biogenic organic matter (Tassi et al., 2022). (e) Example of leaching of graphitic carbon during fluid infiltration in graphitic schists (Vitale Brovarone et al., 2020b). Note the presence of abundant graphitic carbon included in lawsonite, and the absence of it in the surrounding matrix. (f) Fluid-deposited graphite-quartz vein, Amitsoq graphite mine, Southwestern Greenland (Bondam, 1992; Fratelli, 2023). Sample provided by GreenRoc Strategic Materials plc. (g) Discordant graphite veins in high-pressure serpentinite, Balangero, Western Italian Alps (Vitale Brovarone et al., 2017). (h) Network of discordant carbonate veins in serpentinite, Alpine Corsica, France. (i) Dolomite formation in metasomatized garnet peridotite from the Ulten zone, Italy (Sapienza et al., 2009; Consuma et al., 2020). (j) Blue boron-bearing diamond, with dark inclusions of ferropericlase [(Mg,Fe)O], olivine [(Mg,Fe)₂SiO₄] and nyerereite [Na₂Ca(CO₃)₂] studied in Smith et al. (2018). This gem weighs 0.03 carats and its depth of formation resulted to be at least 430 km in the transition zone. (a) Photo credits Orlando Olivieri. (j) Photo credit Evan M. Smith/GIA.</figcaption></figure>
<figure><img src="image_8.png" /><figcaption>**Fig. 6** Processes of carbon degassing from carbonate minerals. In all subfigures, $[Ca_{TOT}]$ is total carbon concentration (dotted lines) and $[Ca_{TOT}]$ is total calcium concentration (solid lines). (a) Pressure-temperature diagram showing the topology of decarbonation reactions (carbonate-silicate reactions) for varying H₂O-CO₂ ratios, here shown as XCO₂. The classical reaction calcite + quartz = wollastonite + CO₂ is shown as an example. (b) High-pressure determinations of calcite solubility in H₂O. Solubilities in molality. Experimental data from Caciagli and Manning (2003), Fein and Walther (1989), and Walther and Long (1986). (c) Schematic diagram showing carbonate solubility variations as a function of fO₂, expressed as Δ log relative to the quartz-fayalite-magnetite buffer (QFM). Vertical dashed line: cc* calcium carbonate solubility transition point. Defined as the fO₂ at which $[Ca_{TOT}]$ is 1% greater than $[Ca_{TOT}]$ at the hematite-magnetite buffer relative to the $[Ca_{TOT}]$ plateau within the calcium carbonate + portlandite field. At fO₂ values greater than cc*, $[Ca_{TOT}]$ is approximately constant; below cc*, $[Ca_{TOT}]$ increases as fO₂ decreases. The schematic diagram illustrates the higher solubility ($[Ca_{TOT}]$) and carbon mobility of calcium carbonate under reducing conditions. Vertical shaded rectangle: calcite + portlandite + fluid stability field (metastable in C). Vertical solid line: limit of calculation, where X(H₂O) = 0.9. (a) Modified after Stewart EM, Ague JJ, Ferry JM, Schiffries CM, Tao R-B, Isson TT, and Planavsky NJ (2019) Carbonation and decarbonation reactions: Implications for planetary habitability. *American Mineralogist* 104(10): 1369–1380. https://doi.org/10.2138/am-2019-6884. (b) Modified after Manning CE, Shock EL, and Sverjensky DA (2013) The chemistry of carbon in aqueous fluids at crustal and upper-mantle conditions: Experimental and theoretical constraints. *Reviews in Mineralogy and Geochemistry*, 75(1): 109–148. https://doi.org/10.2138/rmg.2013.75.5. (c) Modified after Lazar C, Zhang C, Manning CE, and Mysen BO (2014) Redox effects on calcite-portlandite-fluid equilibria at forearc conditions: Carbon mobility, methanogenesis, and reduction melting of calcite†. *American Mineralogist* 99(8–9): 1604–1615. https://doi.org/10.2138/am.2014.4696.</figcaption></figure>

An increasing body of literature has documented examples of natural carbonate reduction at various metamorphic settings (Boutier et al., 2024b; Galvez et al., 2013; Tao et al., 2018; Peng et al., 2021; Vitale Brovarone et al., 2017; Wang et al., 2022) (Fig. 5b). Although at least some of these cases can be explained by carbonate dissolution in aqueous fluids under reducing conditions, the immiscibility of H₂ in metamorphic fluids under a vast pressure-temperature range may promote carbonate-H₂ reactions in the absence of an initial aqueous fluid. In this instance, the term carbonate hydrogenation should be preferred. A simple example of carbonate reduction is.

$$CaCO_3 + 4H_2 \rightarrow Ca(OH)_2 + CH_4 + H_2O$$

where the breakdown of carbonate is promoted by hydrogenation. Carbonate reduction and hydrogenation are documented in subduction zone metamorphic rocks as a consequence of fluid-rock interactions such as serpentinization (Boutier et al., 2024b; Peng et al., 2021; Vitale Brovarone et al., 2020a), or during prograde or early retrograde metamorphism of weakly altered oceanic crustal rocks (Wang et al., 2022). The same processes are also expected to occur at mantle pressures and temperatures (Scott et al., 2004). Fluxes of carbon through these processes are largely unquantified. Fluxes of CH₄ produced through metamorphic decarbonation reactions of ultramafic and mafic rocks are estimated at 1.5 Mt. CH₄ yr⁻¹ (Vitale Brovarone et al., 2017) and up to about 11 Mt. CH₄ yr⁻¹ (Zhang et al., 2023).

## Mobilization of carbon from organic matter, graphite, and diamond

During diagenesis and low-temperature (<300 °C) metamorphism, organic matter undergoes a series of chemical reactions that progressively release heteroatoms such as hydrogen, sulfur, nitrogen, and oxygen, in addition to carbon. Following laboratory experimental results and observations in natural settings, this process results in the carbonization and – at higher temperature – graphitization of organic matter (Vandenbroucke and Largeau, 2007) (Fig. 7a). With progressive loss of heteroatoms during carbonization (or kerogen transformation) and graphitization, carbon is released in the fluid in the form of hydrocarbon oils and gasses (Buseck and Beyssac, 2014; Tissot and Welte, 1984; Vandenbroucke and Largeau, 2007), promoting the formation of a progressively purer and more structured organic matter, or graphitic carbon (Fig. 5c) (Beyssac and Rumble, 2014). Light hydrocarbon gasses produced during this process are referred to as thermogenic gasses and, as well as for oils, are referred to as biotic in origin (Fig. 5d) (i.e., formed through the transformation of biogenic matter; Etiope, 2015). Studies conducted in low-grade metamorphic terranes suggest that thermogenic degassing is efficient up to about 300 °C, whereas at higher temperatures the contribution of this process is considered negligible (Mullis et al., 1994; Tarantola et al., 2007). Thermogenic degassing is primarily controlled by temperature and does not require an external agent such as aqueous fluids. Such conditions are expected during diagenesis in
<figure><img src="image_9.png" /><figcaption>**Fig. 7** Processes of carbon degassing from organic matter and graphitic carbon. (a) Biotic pathways of solid organic carbon mobilization to form hydrocarbon oils and gasses. Van Krevelen diagram showing the transformation pathways for kerogens and mobilization of carbon as biotic hydrocarbon oils and gasses during carbonization and successive graphitization. Kerogen types I, II, and III derive from lacustrine algae, marine micro-organisms, and terrestrial plants, respectively, from. The immature zone represents conditions prior to the release of large amounts of oil and gas hydrocarbons. The diagram also shows compositional data of graphitic carbon from selected metamorphic rocks from Grew (1974). (b-c) Abiotic pathways of solid organic carbon mobilization in aqueous fluids. Diagrams calculated with the Thermotopes-COH software (Boutier et al., 2024a). (b) C-O-H thermodynamic model at 500 °C and 1 GPa. The lower part of the diagram, below the carbon saturation curve, corresponds to carbon undersaturated conditions (no graphite/graphitic carbon/diamond stable). See Holloway (1984) and Connolly and Cesare (1993) for details. The subfigure also shows the predicted fluid compositions for carbon-saturated systems at water-maximum conditions (ΔFMQ = -0.64 at the model conditions), ΔFMQ = -2, and ΔFMQ = +0.5 for a carbon activity (aC) equal to one (graphite stable). Note the higher carbon concentrations in more reduced or more oxidized carbon-saturated fluids relative to the so-called water maximum. The carbon saturation curve for a less graphitic carbon form is also shown and calculated with aC equal to 10. Note the higher solubility for aC > 1, corresponding to less graphitic carbon forms less crystalline than graphite. (c) Pressure-temperature diagram showing total dissolved (Xc), calculated by summing the molar fractions of CH4 and CO2 (XCH4 and XCO2, respectively), for a graphite/diamond-saturated fluid at water maximum. At these conditions, CH4 and CO4 have the same concentration. Other C-bearing fluid species are in negligible concentrations. It can be seen that the solubility of graphite/diamond increases with temperature and decreases with pressure. (a) Modified from Buseck PR and Beyssac O (2014) From organic matter to graphite: Graphitization. *Elements* 10(6): 421–426. https://doi.org/10.2113/gselements.10.6.421, after Vandenbroucke M and Largeau C (2007) Kerogen origin, evolution and structure. *Organic Geochemistry* 38(5): 719–833. https://doi.org/10.1016/j.orggeochem.2007.01.001.</figcaption></figure>

sedimentary basins, or during prograde metamorphism in the shallow forearc of subduction zones margins, in accretionary prisms, and in contact metamorphic aureoles hosted in organic-rich sedimentary rocks. The contribution of thermogenic degassing of hydrocarbons induced by igneous intrusions may release carbonic fluids in amounts high enough to affect global climate conditions. For example, it has been proposed that the end-Triassic mass extinction and initial Eocene global warming may have been caused by voluminous intrusions of mantle-derived melts in organic-rich sedimentary basins at large igneous provinces (Capriolo et al., 2021; Svensen et al., 2004).

Beyond the conditions favorable to extensive thermogenic degassing, the contribution of reduced solid carbon materials to geological carbon degassing may be minimal in closed systems (Connolly, 1995; Pattison, 2006), even though high-temperature thermogenic degassing may still be present (Boutier et al., 2024b). At those conditions, graphitic carbon solubility becomes the dominant mechanism of carbon mobilization from organic carbon solids to fluids, especially in open systems flushed by metamorphic aqueous fluids (Fig. 5e) (Tumiati et al., 2020, 2022; Vitale Brovarone et al., 2020b; Zhang et al., 2018b). The carbon-rich gasses produced by aqueous dissolution of graphitic carbon should be considered abiotic in origin (Fig. 7), even though – in the absence of carbonate minerals, see below – their carbon isotope composition may reflect a biotic origin.

The solubility of graphitic carbon in aqueous metamorphic fluids depends on the redox state, on the pressure and temperature conditions, and on their physicochemical properties (e.g., crystallinity). The amount of carbon that can be dissolved from organic materials to aqueous fluids is higher for strongly oxidized or strongly reduced conditions, and is minimized for intermediate redox conditions at the so-called “water maximum” of carbon-saturated C-O-H fluids (Fig. 7b) (Connolly and Cesare, 1993; Holloway, 1984). Pressure decreases the solubility of graphitic carbon, whereas temperature has an opposite effect (Fig. 7c). Crystalline organic matter is less soluble than amorphous organic matter (Fig. 7b) (Toffolo et al., 2023 and references therein; Tumiati et al., 2020, and references therein; Vitale Brovarone et al., 2020b). Experimental work has shown that the solubility of organic matter, from amorphous to graphitic carbon forms, is dependent on both the structure of the considered material and its surficial properties (Toffolo et al., 2023). Increasing temperature during prograde metamorphism promotes the progressive crystallization of organic matter, increasing its solubility. In subduction zones, increasing pressure during prograde metamorphism may, depending upon thermal regimes, counterbalance this effect (Fig. 7c). Compared to a pure C-O-H system, the presence of silicate minerals can increase the solubility of graphitic carbon in aqueous fluids by up to 30%, as a consequence of a decrease in water activity probably related to organic species containing Si-O-C and Si-O-Mg bonds (Tumiati et al., 2017).

Experimental work has demonstrated that, for graphite-aragonite-water equilibria at fH₂ = FMQ (equivalent to fO₂ expressed as ΔFMQ = +0.61 log units) at subarc conditions, the largest fraction of dissolved CO₂ derives from graphite solubility, whereas carbonate dissolution at the investigated conditions produces bicarbonate ions mostly (Tumiati et al., 2022). In turn, the isotope
composition of the experimentally produced, graphite-derived CO₂ reflected carbonate-CO₂ equilibria, with implications on the interpretation of volcanic gas compositions and their slab source materials.

As discussed for carbonate minerals, interactions between organic matter and H₂-rich fluids may promote the mobilization of carbon from solid phases to light hydrocarbons in fluids. Experimental results obtained at pressure-temperature conditions consistent with subduction zone and upper mantle settings indicate that graphite hydrogenation to form CH₄ may be a very efficient process (Peña-Alvarez et al., 2021). Experiments of diamond hydrogenation showed similarly fast, yet slightly lower reaction rates compared to graphite (Peña-Alvarez et al., 2021). Natural examples of graphite hydrogenations are also documented (Vitale Brovarone et al., 2017).

## Carbon reprecipitation at depth

Not all the carbon being mobilized from metamorphic and metasomatic processes is degassed to Earth's surface. Precipitation of carbonate minerals, graphite, organic compounds, or diamonds from geological fluids or melts is a common process taking place at virtually any depth geological setting. Carbon saturation, redox, pressure/temperature variations and fluid-rock chemistry are some of the key parameters controlling the precipitation of carbon minerals from fluids or melts. At least some diamonds are thought to represent precipitates from deep/ultradeep carbonic fluids (Smith et al., 2016; Stachel et al., 2017; Thomassot et al., 2007). Voluminous or trace amounts of graphite form from percolating fluids and their interactions with rocks and/or other fluids and in rocks and geological contexts spanning mafic and ultramafic rocks at convergent and divergent margins (Luque et al., 2014), and virtually any rock types at various metamorphic, magmatic, and hydrothermal settings within the crust (Fig. 5f,g) (Boutier et al., 2024b; Galvez et al., 2013; Luque et al., 2014; Rumble et al., 1986; Tao et al., 2018). This also extends to other forms of condensed carbon materials, precipitation of which may serve as a carbon source for microbial pathways in the deep subsurface (Debret et al., 2022; Ménez et al., 2018). Carbonate precipitation in upper mantle rocks has been documented in multiple settings. Carbonate veins and carbonation processes, i.e., pervasive replacement of silicate rocks by carbonate minerals, recycling deep carbon in fluids are ubiquitous in crustal and mantle rocks, as demonstrated by natural and experimental data (Fig. 5h,i) (Consuma et al., 2020; Bouilhol et al., 2022; Hu et al., 2021; Kelemen et al., 2011; Menzel et al., 2018; Okamoto et al., 2021; Peng et al., 2021; Piccoli et al., 2016; Sieber et al., 2018). The so-called cold nose of the forearc mantle wedge of many subduction zones represents a characteristic zone of intense reprecipitation of dissolved carbon through carbonation processes (e.g., De Obeso et al., 2022; Kelemen et al., 2022; Menzel et al., 2018, 2024). In the forearc region of Costa Rica, carbon and noble gas data from hot springs indicate that a large fraction — as high as 91% — of carbon in fluids released from devolatilization reactions in the subducting slab reprecipitate as carbonate minerals in the crust before reaching the surface (Barry et al., 2019). Within the biosphere, microbial activity may also contribute to the modulation of carbon degassing through the precipitation of carbonate minerals or organic biomass (see Section “Carbon at the interface between the geosphere and biosphere”).

## Movements of deep carbon in melts

### Carbon in intraplate magmatism

Partial melting in the asthenospheric (convecting) mantle is an inescapable consequence of the presence of CO₂ and H₂O (Fig. 8). Nominally anhydrous minerals can store small amounts of H₂O, which results in a decrease of their melting temperature from 10 °C to ~100 °C; the more H₂O, the lower the melting temperature (Bell and Rossman, 1992; Bolfan-Casanova, 2005; Demouchy and Bolfan-Casanova, 2016; Green, 1973; Hirschmann et al., 2009). In contrast, carbon and CO₂ are insoluble in mantle minerals at asthenospheric conditions, but can be stored as carbonate minerals in relative oxidizing conditions. Accordingly, various constraints from experimental petrology have long shown that, independently of its concentration in the system, carbon stored as carbonates decreases the solidus of mantle rocks by several hundred degrees (Dasgupta, 2018; Green, 1973; Wallace and Green, 1988; Falloon and Green, 1989; Hammouda and Keshav, 2015).

Geochemical surveys have identified intermingled depleted and enriched mantle regions with H₂O and CO₂ contents ranging from a few ppm to a few wt% (tentative average: 50–250 ppm, Hirschmann, 2010; Marty, 2012; Le Voyer et al., 2017; Shimizu et al., 2019). This is thus expected to produce minute amounts of partial melts in the asthenosphere, of the order of 0.1 vol% to several vol %. Possible variations in pressure and temperature in the asthenospheric mantle are not expected to prevent the presence of incipient melts. Oxidation state could however preclude incipient melting if carbon is present as diamonds rather than oxidized carbon (i.e., carbonate minerals). Oxygen fugacity conditions in the convective upper mantle are however broadly oxidizing enough to enable incipient melting from 50 km down to 200–300 km depth (Brey et al., 1983; Gaillard et al., 2015; Kushiro, 1975; Moussallam et al., 2019; Rohrbach and Schmidt, 2011; Stagno et al., 2013; Wyllie et al., 1983; White and Wyllie, 1992; Zhang et al., 2024). At greater depths, diamonds are expected to prevail, preventing the occurrence of melting (Rohrbach and Schmidt, 2011; Stagno et al., 2013).

Pressure and temperature variations in the asthenosphere are, however, expected to significantly affect the chemical nature and volume fraction of melts produced by CO₂-H₂O-bearing peridotite melting. High-pressure and low-temperature melting tends to produce low-SiO₂ carbon-rich melts (i.e., carbonatites), whereas low-pressure and high-temperature melting tends to produce silicate melts (i.e., basalts). In between, all intermediate compositions are possible (Brey et al., 2008; Dasgupta et al., 2007, 2013b; Gudfinnsson and Presnall, 2005; Massuyeau et al., 2015, 2021). Fig. 9 illustrates this variability in melt composition versus depth for a 'normal' mantle (140 ppm CO₂; 240 ppm H₂O; mantle potential temperature = 1350 °C; see Massuyeau et al., 2021). The
<figure><img src="image_10.png" /><figcaption>**Fig. 8** Composition (a) and volume (b) of magmatic liquids equilibrated in the ambient convective upper mantle in terms of SiO₂, CO₂ and H₂O, as calculated with MAGLAB (Massuyeau et al., 2021). The inset represents the adiabatic profile used for the calculations, considering a normal mantle with 140 ppm CO₂ and 240 ppm H₂O.</figcaption></figure>

<figure><img src="image_11.png" /><figcaption>**Fig. 9** Comparison of solidus curves for a bulk peridotite composition in volatile-free (bold black line), H₂O-bearing (thin blue line), and CO₂-bearing (dashed red line) systems, presented in a temperature and pressure/depth diagram. Dry solidus after Hirschmann (2000); H₂O-bearing solidus after Hirschmann et al. (2009) for 50 and 750 ppm H₂O; CO₂-bearing solidus from Dasgupta (2013, 2018). The solid mantle adiabatic pressure-temperature path corresponds to a potential temperature of 1350 °C. fO₂ variations versus depth as well as the nature of the various carbon-bearing phases (fluid-carbonate-diamond/graphite) are indicated on the left part of the figure.</figcaption></figure>

melt compositions transition relatively abruptly from carbonatites to basalts in response to small variations in pressure and temperature (Dasgupta et al., 2007, 2013b; Stagno and Frost, 2010). This reflects the strongly non-ideal mixing properties between molten carbonate minerals and molten silicate minerals, which often results in liquid-liquid immiscibility (Brooker and Kjarsgaard, 2011; Freestone and Hamilton, 1980; Massuyeau et al., 2015; Nabyl et al., 2020; Novella et al., 2014). Immiscibility between silicate and carbonate melts is a popular mechanism to explain the formation of many carbonatites found at the surface (Nabyl et al., 2020; Yaxley et al., 2022). It has often been described in experiments conducted at moderate pressure and in variably fractionated, low-temperature alkaline melts (e.g., Nabyl et al., 2020; Weidendorfer et al., 2017). Yet, at asthenospheric pressure-temperature conditions, such immiscibility has only been observed during experimental simulations of partial melting of eclogite or pelitic rocks (Hammouda, 2003; Kiseeva et al., 2012; Litasov and Ohtani, 2010; Thomsen and Schmidt, 2008; Yaxley and Brey, 2004), whereas peridotite melting does not seem to produce immiscible carbonate-silicate melts. Indeed, increasing pressure and water content
tends to decrease the non-ideal carbonate-silicate mixing, enabling the formation of intermediate melt compositions, such as the famous kimberlites, which are in equilibrium with mantle peridotites at depth of 120–250 km (i.e., pressure of 4–8 GPa; Massuyeau et al., 2021; Stamm and Schmidt, 2017).

The chemical composition of these intraplate magmas are diverse, and range from low to high SiO₂ contents, which seemingly reflects the interruption of the adiabatic melting regime at variable depths as illustrated in Fig. 9 (Gudfinnsson and Presnall, 2005; Massuyeau et al., 2021; Tappe et al., 2007). This interruption at variable depth is controlled by the thickness of the (non-convective) lithosphere, which ranges from >200 km beneath cratons, where kimberlites are found, to ~60 km in some oceanic domains, where basanites or alkali basalts are extracted from the convective asthenosphere. Therefore, the combination of the CO₂-driven mantle melting regime and the variable thickness of the lithospheric lids provides a simple explanation of the variability of melt compositions found in intraplate domains.

These melts are produced by adiabatic incipient melting of the asthenosphere and constitute ~0.1 to several vol% of the mantle, depending on the bulk concentrations of CO₂-H₂O and the potential temperature. On top of these thermochemical variables, melt accumulation and mantle compaction processes can affect melt behavior in the mantle (Grégoire et al., 2006; Havlin et al., 2013; Keller et al., 2017; Massuyeau et al., 2021; Soltanmohammadi et al., 2018). These accumulation/migration processes are driven by the peculiar CO₂ mantle melting regimes, which produce a range of melt densities (Agee, 2008; Dobson et al., 1996; Jing and Karato, 2011; Massuyeau et al., 2023; Lange and Carmichael, 1987; Ritter et al., 2020; Sakamaki et al., 2006; Suzuki et al., 1995), viscosities (Behrens and Schulze, 2003; Dingwell, 2007; Dingwell et al., 2022; Dobson et al., 1996; Giordano et al., 2008; Kushiro, 1976; Ritter et al., 2021; Sakamaki et al., 2013; Whittington et al., 2009), and mantle permeabilities (Grégoire et al., 2006; Keller et al., 2017), inducing variable melt buoyancies along the melting column and contributing to volatile redistribution (Bekaert et al., 2021; Dasgupta, 2013; Gaillard et al., 2019; Schettino and Poli, 2020). In many cases however, calculated melt mobilities are lower than 1 cm yr⁻¹, which implies that they are less mobile than the convective mantle itself. It is only in regions where melt fractions are higher than the 'normal' mantle (due to higher CO₂-H₂O or higher potential temperatures) that melt mobility could exceed mantle convection rates (Gaillard et al., 2019) and thus lead to migration/accumulation/extraction processes. In most cases, however, one expects that such incipient melts do not reach the surface and remain trapped at the lithosphere-asthenosphere boundary, causing metasomatism of the lithosphere (see Section "Sub-continental lithospheric mantle").

We do not have in situ analyses of these CO₂-driven melting processes, but electrical and seismic geophysical observations have been used to interpret in near real-time the formation and interconnection of these incipient melts in the asthenosphere (Chantel et al., 2016; Gardés et al., 2020; Hammouda and Laporte, 2000; Kawakatsu et al., 2009; Kawakatsu and Utada, 2017; Laumonier et al., 2017; Minarik and Watson, 1995; Massuyeau et al., 2021; Rychert et al., 2020; Schmerr, 2012; Selway and O'Donnell, 2019; Selway et al., 2019; Sifré et al., 2014; Yoshino et al., 2010, 2018). The interpretation of remote sensing geophysical data is not unequivocal as several geological processes could explain one geophysical proxy. The recently discovered 'petit-spots,' a new class of volcanoes, might represent some leakage of these asthenospheric melts, permitted by the bending and stretching of the oceanic lithosphere prior to enter into subduction (Hirano et al., 2006; Hirano and Machida, 2022). These melts are very CO₂-rich and may indicate incipient melts trapped at the lithosphere-asthenosphere boundary. The seminal question is then how broadly distributed those melts are. This requires the analysis of constraints from experimental petrology, mantle geochemistry, geophysics and an understanding of the occurrence of petit-spots. Hammouda et al. (2021) suggested a global layer of such incipient melts at the lithosphere-asthenosphere boundary, but mantle heterogeneities with carbon-depleted and carbon-enriched regions might speak for local melt enrichments.

## Carbon in arc basalts

At subduction zones, arc volcanism transfers some CO₂ from the slab to the Earth's surface. Hydrous silicate melts, which are formed in the mantle wedge, can transport carbon as dissolved carbonate minerals in basaltic melts or in supercritical fluids. Due to large uncertainties and large observed variations in CO₂ fluxes at arcs (Aiuppa et al., 2019; Mason et al., 2017), few estimates of the average CO₂ content of primary arc magmas are available. Outgassing rates of CO₂ at volcanic arcs are indeed highly variable, with the greatest CO₂ emitters involving an overwhelming contribution from thermally destabilized sedimentary carbonate rocks (Aiuppa et al., 2019; Iacono-Marziano et al., 2009; Mason et al., 2017). In a review paper, Plank and Manning (2019) suggested that up to 1 wt% CO₂ could be dissolved and transported in primary arc basalts. This is 10 times greater than the average CO₂ contents of primary mid-ocean-ridge basalts. This estimate is indirect as it is based on C/S ratio measured in volcanic gasses, with sulfur (i.e., mostly SO₂ outgassing) being affected by complex processes during transcrustal magma transport (e.g., Mungall et al., 2015). In spite of these unknown, it appears that arc magmas have higher CO₂ concentrations than mid ocean ridge basalts; in terms of total CO₂ fluxes however, emissions from mid-ocean ridges and arcs are, within uncertainties, equivalent (Kelemen and Manning, 2015; Fig. 4). Below we analyze the dynamic of carbon transfer throughout the plumbing system of arc volcanoes.

The solubility of carbon in molten basalts is low and it increases broadly linearly with pressure so that it reaches ~1 wt% at pressure equivalent to the base of the continental crust (~1 GPa, corresponding to 30 km depth; Iacono-Marziano et al., 2012). As most other volatile components (i.e., water, sulfur and halogens) are much more soluble in silicate melts than CO₂, CO₂ is the first volatile component released by magma during their transcrustal journey. Considering furthermore that most rising melt remains locked in the crust and evolves as intrusive bodies, this implies that a large portion of magmatic CO₂ in arcs is outgassed at great depth and must exist as supercritical fluids in the crust. Variation in melt compositions can affect CO₂ solubility and thus the depth of fluid saturations (Iacovino et al., 2021; Wieser et al., 2022). Alkaline primary melts dissolve more CO₂ and thus reach fluid-saturation at shallower depth. If primary arc basalts contain 1 wt% CO₂, depending on melt compositions they will start
forming a supercritical CO₂-rich fluid phase at depths ranging from 30 to 10 km as they ascend through the crust. These CO₂-rich deep fluids rise through the crust in a way that can be decoupled from the magmatic flow (Edmonds et al., 2022). Thus, CO₂ can be released from stagnant melts in the crust (non-eruptive), which is not the case of other magmatic volatile components bar noble gasses. It is thus difficult to estimate and model the global emissions of CO₂ by arc volcanoes since these emissions start from the deep crust, yield surficial fluxes that are very diffused (Chiodini et al., 2005; Werner et al., 2019), and can originate from a combination of intrusive/extrusive events (Edmonds et al., 2022). Using remote satellite data compiled in the period 2005–2015, Fischer et al. (2019) have shown that non-eruptive CO₂ degassing by volcanoes largely exceeds that emitted during volcanic eruptions (30 times more) at planetary scale. This survey also suggests that a significant portion of subducted carbon remains stored in the arc lithosphere, possibly in the form of metasomatic carbonates or at shallower levels, in aquifers. Finally, as suggested by Iacono-Marziano et al. (2009), intrusion of magma in crustal reservoirs hosted by sedimentary carbonate rocks can lead to melt-carbonate interactions massively producing CO₂; Mason et al. (2017) and Aiuppa et al. (2019) provide regional and global evidence for the importance of this mechanism, which adds some complexity in our analyses of the volcanic CO₂ emissions at a planetary scale.

## Diamonds as clues on deep and super-deep carbon movements

Diamond is the highest-pressure form of solid carbon and is stable over a depth range of hundreds of kilometers from the upper mantle down to the transition zone and into the lower mantle (Day et al., 2023, and references therein). Diamonds are classified in two very different categories: lithospheric, which formed beneath the cratonic areas between 120 and 130 and 200–220 km depth and sub-lithospheric (also called super-deep diamonds), which formed at depths greater than 280–300 km and down to at least 700–800 km depth within the lower mantle or along the subducting slabs (Nestola et al., 2023; Timmerman et al., 2023). From the above (see Section “Carbon throughout the Earth’s mantle”), it is evident that diamonds can form within very different geological environments from the upper to the lower mantle and can therefore have peridotitic and eclogitic origin (and much rarely websteritic; Stachel et al., 2022a). These precious stones are therefore unique minerals that can provide crucial information regarding the movements of carbon within the deep Earth in space and time.

Diamonds may experience long vertical journeys through the Earth's interior, from the deep mantle to the surface, as documented by ultra-deep diamonds recovered at the Earth's surface (Nestola et al., 2018). They can also move horizontally for thousands of kilometers, likely transported on the 'shoulders' of tectonic plates. For example, diamonds emplaced within the North Australian Craton and now found within the Meratus Mountains in Southeast Borneo, Indonesia, testify to horizontal movements of lithospheric carbon in diamonds for thousands of kilometers (Kueter et al., 2016). Lastly, eclogitic diamonds related to subduction zones demonstrate the transfer of carbon from the crust to hundreds of kilometers downwards into the mantle within subducting slabs (Lorenzon et al., 2023).

The possibility of tracking the movements of diamonds within and among different geological environments inside the Earth with diamond ages would represent an unprecedented tool to retrieve key missing tiles of solid carbon geodynamics between 120 and 130 and 700–780 km depth range (Gu et al., 2022; Nestola et al., 2018; Thomson et al., 2016).

## Diamonds over (geological) time

Most (if not all) carbon reaching depths greater than 150 km, temperatures above 1320 °C, and oxygen fugacity below -1.8 log units relative to the FMQ buffer within the Earth's mantle precipitate as diamonds (Stagno and Frost, 2010). Diamond precipitation is thought to happen through oxidation or reduction of carbon-bearing species like CO₂ and/or CH₄ carried by fluids and/or melts that interact with the surrounding rocks (Day et al., 2023). This well-accepted formation process is likely occurring at present within the mantle, even if no diamonds are presently being erupted to the Earth's surface. There are no doubts that every day, new diamonds are being produced within the Earth's mantle, but they are unable to reach the surface because deep, kimberlitic eruptions have not occurred on our planet since at least ~30 Ma (Batumike et al., 2008). Kimberlites are rare, mafic volcanic rocks erupted directly from deep regions of the mantle; these rocks are able to transport diamonds to the Earth's surface and this is why they are intensely explored on cratonic areas worldwide. The reason why kimberlites are no longer erupted is not certain, with recent work suggesting that kimberlitic eruptions are promoted by specific slab flux and subduction angles of tectonic plates (Mather et al., 2023). These authors reconstructed the slab dip angles of subduction segments for the past 170 million years to understand pulses in kimberlitic eruptions, concluding that high subduction rates promote the return flow of fertile mantle, which in turn increases partial melting and magma generation within the deep mantle (Mather et al., 2023). This geodynamical model could help predict when kimberlite eruptions restart on the Earth and, consequently, when new diamonds will be possibly sampled and transported to the Earth's surface. The youngest kimberlite erupted only about 12,000 years ago (Igwisi Hills volcanoes, Tanzania; Brown et al., 2012), but unfortunately did not transport any diamonds. It cannot be excluded that this young kimberlitic volcanic activity can indicate a reactivation of these global geological processes. Although the age of kimberlitic events is known, constraining the age of diamonds in them remains a major scientific challenge.

Diamond ages are constrained through numerous analytical approaches, and span between 3.5 Ga and 220–140 Ma (for a review, see Smit et al., 2022; Fig. 10a); this demonstrates that diamonds have formed over most of the life of our planet. In particular, peridotitic lithospheric diamonds are the oldest known diamonds, while lithospheric eclogitic diamonds are generally not older than 3.0 Ga (Fig. 10; Shirey and Richardson, 2011; Smit et al., 2022). Unfortunately, just a few super-deep diamonds have
<figure><img src="image_12.png" /><figcaption>**Fig. 10** Diagrams concerning diamond age and isotopic composition. (a) Illustration showing the historical occurrence of diamond growth events on Earth (Fig. 1 in Howell et al., 2020). Diamond symbols represent eclogitic diamond growth events; square symbols represent peridotitic diamond growth events, with additional color differentiation for harzburgitic diamonds (magenta) and herzolitic diamonds (blue); circle symbols represent fibrous growth events, which constrain the age of the corresponding kimberlite eruption. (b) Box plot showing δ¹³C data recorded from diamonds (Fig. 7 in Howell et al., 2020). Redrawn after Howell D, Stachel T, Stern RA, Pearson DG, Nestola F, Hardman MF, Harris JW, Jaques AL, Shirey SB, Cartigny P, Smit KV, Aulbach S, Brenker FE, Jacob DE, Thomassot E, Walter MJ, and Navon O (2020) Deep carbon through time: Earth's diamond record and its implications for carbon cycling and fluid speciation in the mantle. *Geochimica et Cosmochimica Acta* 275: 99–122. https://doi.org/10.1016/j.gca.2020.02.011.</figcaption></figure>

been dated as of today, and based on the existing data, the oldest specimens are not older than 650–700 Ma (Timmerman et al., 2023), therefore appearing to be much younger than the oldest lithospheric diamonds.

## Fluxes of deep carbon in diamonds

Once the possibility for deep carbon in diamonds to move vertically and horizontally in the deep Earth is proven, an estimate of how much carbon can be transported by the movement of diamond can be determined. Firstly, it is important to consider that most diamonds (between 98% and 99%) are lithospheric (Stachel and Harris, 2008). This percentage can significantly vary among world localities, as recently reported by Curtolo et al. (2023) who, on a basis of 1069 specimens, show a percentage of lithospheric diamonds of ~95.4% for all reported Canadian diamonds. Only a few localities (Juina, Brazil; Walter et al., 2011; Kankan, Guinea, Stachel et al., 2000a) appear to contain an extremely high amount of super-deep diamonds. All things considered, it can be assumed that 98% is a representative percentage for lithospheric diamonds worldwide, leaving super-deep diamonds to account for the remaining 2%.

Garber et al. (2018), combining seismological, mineralogical and petrological constraints, buoyancy and electrical conductivity data from natural samples, and mass balance constraints for global carbon, proposed diamonds to account for 2 vol% of lithosphere. This most recent estimate can therefore be considered reasonable as it is in broad agreement with other authors reporting 3 vol% (Anand et al., 2004) or 1 vol% (e.g., Peltonen et al., 2002). With the 2 vol% of lithospheric diamonds, Garber et al. (2018) calculated the mass of carbon as diamond being 10¹⁸ kg (assuming the present-day conditions) for cratonic roots with a shape of an inverted cone with 1000 km base and 50 km height. The authors state that if such roots are similar at the global scale, then the total mass of carbon as diamond would result in about 10¹⁶ t (10²² g), in agreement with estimates of the SCLM carbon budget (Hirschmann, 2018; see Section “Sub-continental lithospheric mantle”).

With respect to the quantification proposed by Garber et al. (2018), we can consider that such mass of diamond would only refer to lithospheric diamonds. In addition to this, an additional 2% of super-deep diamonds should be considered, which would
correspond to further 2 × 10¹⁴ t, resulting in a final 1.02 × 10¹⁶ t worldwide. This amount corresponds to 2% of overall carbon on the planet, or alternatively ~1–10% of carbon in the present-day mantle (Garber et al., 2018; McDonough and Sun, 1995; Dasgupta and Hirschmann, 2010).

## Diamonds and their solid carbon inclusions

Laboratory experiments definitively show that a few carbonate minerals can be stable at pressure and temperature conditions at which diamonds form, such as magnesite, aragonite, a few polymorphs of calcite, dolomite, Na-carbonates (e.g., Cerantola et al., 2023; Litasov et al., 2020; Tao et al., 2014, and references therein). Diamonds and their inclusions can provide important information as for these solid carbon phases actually exist in the deep Earth.

In both lithospheric and sub-lithospheric diamonds, common inclusions of carbonate minerals such as calcite, magnesite and dolomite have been reported (Brenker et al., 2007; Stachel et al., 2022a). However, other, more rare carbonates have been identified in both lithospheric and sublithospheric diamonds including aragonite (Smith et al., 2022); siderite (Stachel et al., 2000b); nyereereite, Na₂Ca(CO₃)₂ (Howell et al., 2012); huntite, CaMg₃(CO₃)₄ (Agrosi et al., 2019); nahcolite, NaHCO₃ (Kaminsky et al., 2009); eitelite Na₂Mg(CO₃)₂ (Kaminsky et al., 2013).

Other carbon-bearing mineral inclusions have also been identified in diamonds in addition to carbonate minerals; these include carbide minerals such as cohenite, Fe₃C (Smith et al., 2016); moissanite, SiC (Kaminsky, 2012); haxonite, (Fe,Ni)₂C₆ (Kaminsky, 2012). Both carbonate minerals and carbide minerals, however, are not common inclusions in diamonds (they are actually extremely rare) and thus they possibly do not represent a significant fraction of the solid carbon related to diamonds.

## Diamonds as trackers of global carbon recycling and geobiological evolution

One of the most intriguing and geologically significant aspects in studying diamonds is the clear evidence that in some instances they show a biotic signature with respect to their carbon isotopic composition (i.e., δ¹³C), displaying values consistent with a biogenically-derived organic carbon source. A very recent and extensive review on the carbon isotopic compositions as a function of the different regions of diamond formation is provided by Stachel et al. (2022b). These authors report carbon isotopic compositions for 4307 worldwide diamonds, of which 3474 are of known paragenesis (mainly peridotitic versus eclogitic, with the websteritic fraction being extremely limited), and with δ¹³C ranging from -41‰ to +2.5‰ (VPDB).

The carbon isotope composition of the mantle, obtained by studying carbonatites, kimberlitic carbonates, mantle xenoliths and volcanic CO₂, is δ¹³C = -5‰ (Deines, 2002). Most of the 1700 peridotitic diamonds in Stachel et al. (2022b) show δ¹³C varying between -2‰ and -8‰ with the modal abundance centered at -5‰. On the other hand, 1204 eclogitic diamonds analyzed by these authors show a much greater variability: while 50% of these diamonds also show an average value of -5‰ similar to the peridotitic types, the remaining 50% show a variable δ¹³C ranging from -10‰ to -41‰, the minimum value mentioned above.

Stachel et al. (2022b) also investigated 209 super-deep diamonds. From this group of diamonds, the so-called asthenospheric-transition zone diamonds (formed between 250 and 660 km depth; 55 samples) show a variability in δ¹³C between -28‰ and +0.7‰ with two modes at about -10‰ and -20‰; the 'lower-mantle' diamonds (formed at depth greater than 660 km, 154 samples) show a strong mode at δ¹³C = -5‰ and a second, weaker mode at about -25‰ (mainly due to the Brazilian diamonds from the Juina locality). The analysis by Stachel et al. (2022b) indicates that peridotitic lithospheric and lower mantle diamonds share the same mantle carbon source. Eclogitic lithospheric diamonds and asthenospheric-transition zone diamonds also share similar values of δ¹³C but in this case in addition to values of -5‰ they are also characterized by much lower values of δ¹³C. It appears clear that in order to search for diamonds showing a biogenic-like signature, we must look among eclogitic and asthenosphere-transition zone diamonds.

The presence of ¹²C-enriched carbon in diamonds has been an issue debated for many years. Seafloor sediments have been long considered as the most favored crustal carbon source given the similar δ¹³C values to surficial sedimentary carbonate minerals (δ¹³C ≈ 0‰) or organic matter (δ¹³C < -20‰). While sedimentary carbonate and graphitized organic carbon can potentially be retained in slabs and buried to significant depth (Duncan and Dasgupta, 2017), they cannot explain the combined carbon and nitrogen isotopic characteristics of eclogitic diamond. Diamonds formed from organic matter, in which carbon and nitrogen are coupled, should retain the positive δ¹⁵N values of organic matter in subducting sediments (Li et al., 2019); however, most (>60%) ¹³C-depleted eclogitic diamonds show δ¹⁵N < 0 (Cartigny et al., 2014). Furthermore, some eclogitic diamonds display extreme negative δ¹⁵N values, which has not been observed in subducting sediments. Therefore, the leading hypothesis is that the diamonds characterized by ¹²C-enriched carbon are related to the presence of subducted biotic organic carbon in altered oceanic crust (see Walter et al., 2011, and references therein). This view has been recently supported by Li et al. (2019), who not only show that the altered igneous portion of oceanic crust has a total carbon content similar to that of subducted sediments, but that the δ¹³C range of normal and biogenic carbonate and organic matter in altered oceanic crust (from -24‰ to +11‰) is similar to the range observed in diamonds (-41‰ to +3‰). In addition, altered oceanic crusts shows has nitrogen with a highly variable δ¹⁵N, assimilated as ammonium ions, into secondary clay minerals during alteration (Li et al., 2019). Clay minerals formed by low temperature (<100 °C) alteration generally show positive δ¹⁵N values and are isotopically similar to organic matter. On the other hand, clay minerals that form by alteration at higher temperatures (>250 °C) show δ¹⁵N as low as -16‰ and are related to the generation of ammonium ions by abiotic N₂ reduction. From these observations, Li et al. (2019) reported three possible carbon and nitrogen reservoirs with distinct isotopic signatures within altered oceanic crust that could mix with the mantle signature (δ¹³C = -5‰; δ¹⁵N = -4‰) to produce the isotopic signatures reflected by most eclogitic and websteritic diamonds. These three reservoirs are: (C1) 'normal carbonate' (negligible organic matter) + ¹⁵N-enriched (sedimentary-like) clay (δ¹³C = 0‰; δ¹⁵N = 16‰);