Forms and fluxes of carbon: Surface to deep

1 FORMS AND FLUXES OF CARBON: 1 SURFACE TO DEEP 2 Alberto Vitale Brovarone1,2,3*, Kevin Wong1, Donato Giovannelli4,5,6,7,8, Benoit de Pins4, Fabrice Gaillard9, 3 Malcolm Massuyeau10, Fabrizio Nestola11, Martha Pamato11, Isabelle Daniel12 4 1Dipartimento di Scienze Biologiche, Geologiche e Ambientali (BiGeA), Alma Mater Studiorum Università di Bologna, Bologna, 5 Italy 6 2Institut de Minéralogie, de Physique des Matériaux et de Cosmochimie (IMPMC), Sorbonne Université, Muséum National 7 d’Histoire Naturelle, UMR CNRS 7590, IRD UR206, 75005 Paris, France 8 3Institute of Geosciences and Earth Resources, National Research Council of Italy, Pisa, Italy 9 4Department of Biology, University of Naples Federico II, 80138 Naples, Italy 10 5National Research Council—Institute of Marine Biological Resources and Biotechnologies (CNR-IRBIM), 60125 Ancona, Italy 11 6Department of Marine and Coastal Science, Rutgers University, New Brunswick, NJ 08901, USA 12 7Marine Chemistry, Geochemistry Department—Woods Hole Oceanographic Institution, Falmouth, MA 02543, USA 13 8Earth-Life Science Institute, Tokyo Institute of Technology, Tokyo 152-8552, Japan 14 9CNRS, BRGM, ISTO, UMR 7327, Université d’Orléans, Institut des Sciences de la Terre (ISTO), 45071, Orléans, France 15 10Institut für Mineralogie, Westfälische Wilhelms-Universität Münster 16 11Department of Geosciences, University of Padova, Padova, Italy 17 12Universite Claude Bernard Lyon1, LGL-TPE, UMR 5276, CNRS, Ens de Lyon, Universite Jean Monnet Saint-Etienne, 18 Villeurbanne, 69622, France 19 *Corresponding author: Alberto VITALE BROVARONE ; [email protected] 20 21 22 23 24 2 1. Introduction 2 25 2. Origins of deep carbon 3 26 3. Carbon at the interface between the geosphere and biosphere 8 27 3.1. Carbon’s central role in life 8 28 3.2. The deep subsurface biosphere 11 29 3.3. Impact of life and organic carbon on the deep carbon cycle 12 30 3.4. Biogenic carbon: from the surface to the subsurface of the planet 13 31 3.5. Effect of life on carbon fluxes to and from the deep Earth 16 32 4. Deep carbon forms 18 33 4.1. Carbon in the Earth’s core 19 34 4.2. Carbon at the interface between the core and the lowermost mantle 22 35 4.3. Carbon throughout the Earth’s mantle 24 36 4.3.1. Reduced lower mantle and transition zone 25 37 4.3.2. Depleted oxidized upper mantle 28 38 4.3.3. Sub-continental lithospheric mantle 29 39 4.3.4. Sub-oceanic lithospheric mantle 32 40 4.4. Deep carbon in the crust 33 41 4.4.1. Continental crust 34 42 4.4.2. Oceanic crust and upper lithospheric mantle 37 43 4.5. Deep abiotic organics 41 44 5. Carbon movements within and between deep reservoirs 46 45 5.1. Movements of carbon from the crust to the deep Earth 46 46 5.1.1. Subduction of crustal carbon 46 47 5.1.2. Subduction erosion 50 48 5.2. Movements of carbon from depth to the crust 50 49 5.2.1. Tectonic movements of solid carbon from the deep Earth to the surface 50 50 5.2.2. Metamorphic and magmatic processes of carbon mobilization 51 51 5.3. Carbon reprecipitation at depth 59 52 5.4. Movements of deep carbon in melts 59 53 3 5.4.1. Carbon in intraplate magmatism 59 54 5.4.2. Carbon in arc basalts 64 55 5.5. Diamonds as clues on deep and super-deep carbon movements 65 56 5.5.1. Diamonds over (geological) time 66 57 5.5.2. Fluxes of deep carbon in diamonds 67 58 5.5.3. Diamonds and their solid carbon inclusions 68 59 5.5.4. Diamonds as trackers of global carbon recycling and geobiological evolution 68 60 6. Secular variation of surficial versus deep carbon reservoirs or steady state? 73 61 7. Concluding remarks 76 62 References 77 63 64 65 Abstract 66 Carbon is an essential element for the coevolution of Earth and life, and its largest fraction is stored in the 67 deep Earth. The availability of carbon at the Earth’s surface or shallow subsurface over the past 3.8 billion 68 years has been modulated by deep carbon processes, and have played a fundamental role in the emergence 69 and diversification of life. In turn, major geobiological changes at the Earth’s surface have profoundly 70 affected the chemical fingerprints and morphological forms of carbon entering the deep Earth through 71 subduction. This chapter presents an overview of the geology of deep carbon, from its origins and forms 72 throughout the Earth’s history, to its movements and fluxes between shallow and deep reservoirs. 73 74 Keywords Abiotic Organics; Biotic and Abiotic Carbon; Carbon budgets; Carbon Cycle; Carbon Fluxes; 75 Carbon in Melts; Deep Carbon; Coevolution of Earth and Life; Origin of Carbon; Metamorphic Carbon 76 Degassing; 77 78 Key points/Objectives box 79 - We present the origins and forms of deep carbon throughout the Earth’s history 80 - We summarize the reservoirs of deep carbon from the core to the crust 81 - We review the available estimates of carbon fluxes from the surface to the deep Earth and vice-82 versa 83 4 84 85 Keywords 86 87 1. Introduction 88 Carbon (C) is an essential element for life and the functioning of our planet. Its presence and 89 quantity in the biosphere and atmosphere has largely been controlled by deep Earth processes throughout 90 the Earth’s history (Berner, 1994). Deep carbon is recycled through tectonic (or mechanical), metamorphic, 91 magmatic, and fluid-rock interaction processes, and — within the parameter space for life — through 92 biologically mediated processes. However, the largest fraction of Earth’s carbon is stored in deep Earth 93 reservoirs such as the mantle and the core (Dasgupta and Hirschmann, 2010). Because of that, the 94 distribution, movements, and fluxes of deep carbon within and between geological reservoirs are still 95 largely unconstrained and model-dependent (Dasgupta, 2013; Wood et al., 2013). The long-term carbon 96 cycle on our planet is primarily controlled by plate tectonics, with convergent plate margins and subduction 97 driving the input of surface carbon into the deep Earth and volcanism and diffuse degassing acting as the 98 main pathway for the return of carbon to the Earth’s surface (Jarrard, 2003; Kelemen and Manning, 2015; 99 Lee et al., 2019; Plank and Manning, 2019; Werner et al., 2019). 100 Defining the boundary between surface and deep carbon is not straightforward, as the two 101 reservoirs are intimately connected. Deep carbon represents the largest fraction of carbon on Earth and 102 controls its long-term cycle (~1023 moles; Lee et al., 2019). On the other hand, surface carbon represents a 103 much smaller mass (1021–1022 moles; Hayes and Waldbauer, 2006; Javoy et al., 1982), but is central for 104 and profoundly affected by biological activity, with important implications on the types, sizes and chemical 105 fingerprints of both solid carbon being transferred from the surface to the deep Earth at subduction zones, 106 and deep carbon degassing into the atmosphere. The emergence and diversification of life over the Earth’s 107 history have substantially affected the diversity, distribution, and signatures of carbon-bearing geological 108 materials (Hazen et al., 2019), thereby affecting deep carbon processes and chemical fingerprints. This 109 includes subsurface life that acts on deep carbon to alter its composition and pathways of cycling, besides 110 being per se a reservoir of deep carbon . 111 5 Given the diverse interconnections among processes acting at various depths within the crust and 112 mantle, and the complex interplay between geological and biological processes, the potential solution to 113 identify a boundary between surface and deep carbon can be found in the respective sizes and response 114 times of carbon in the various reservoirs. Lee et al. (2019) defined the limit between endogenic (deep; 115 Earth’s lithosphere and mantle) and exogenic (surface; oceans, atmosphere, biosphere, reactive marine 116 sediments) carbon at 107 Gt mass of carbon (where 1 Gt = 1015 g) and 10,000–100,000 years of residence 117 time, with endogenic carbon being richer in carbon and slower in response time relative to exogenic carbon. 118 In the lithosphere and mantle reservoir, the residence time of carbon is estimated to be in the range of 1 119 million to 1 billion years, in comparison to the <100,000 year timescales typical of exogenic carbon cycles 120 (Lee et al., 2019). 121 Starting from the origin of deep carbon on Earth, this chapter presents a summary of the most 122 common forms of deep carbon, their movements — from solids to fluid phases — and fluxes. Although 123 carbon in the biosphere is considered exogenic (response time <10,000 years; Lee et al., 2019), this chapter 124 also presents some key aspects of carbon in the biosphere. This is because of the central role of life in 125 controlling the amounts, forms, and chemical signature of carbon not only at the Earth’s surface, but also 126 in the subsurface where endogenic and exogenic processes together act to modulate the fluxes of carbon 127 between the surface and the deep Earth. 128 2. Origins of deep carbon 129 Isotopic constraints on Earth’s bulk carbon are consistent with a chondritic origin (Broadley et al., 130 2022). The carbon content of non-carbonaceous enstatite chondrites and carbonaceous chondrites could 131 explain the concentration of carbon and the 13C/12C ratio of the bulk silicate Earth (BSE; Hirschmann, 2018; 132 Marty et al., 2016). A similar chondritic provenance of most life-forming elements, i.e., hydrogen and 133 nitrogen, is also consistent with the elemental and isotopic abundances for these elements (Broadley et al., 134 2022). It therefore seems likely that carbon and other life-forming elements were delivered to Earth during 135 planetary accretion. The geological history of carbon thus started with the earliest planetary events that 136 have shaped planet Earth (Figure 1). 137 Carbon was first delivered by chondrite-like materials on small planetary bodies that formed within 138 the first 1–3 millions of years of the solar system, via a mechanism termed ‘pebble accretion’ (Johansen 139 and Lambrechts, 2017; Li et al., 2016). During this early stage, radioactive heating was intense enough to 140 induce melting of these bodies, enabling magma ocean and core differentiation. Upon this early 141 differentiation, carbon was greatly partitioned into the core (Li et al., 2016), but due to the combination of 142 6 low gravity and high temperature, outgassing and loss of gaseous carbon was certainly prevailing. This 143 implies that the carbon content of these small planetary bodies was low. It was only when the size of these 144 planet-forming bodies reached diameters of ~1000 km that gravity was strong enough to limit gas escape. 145 Carbon incorporation in the planet interior then became considerable, with the particularity that under the 146 prevailing reducing conditions, a significant, amount of carbon must have been dissolved in the forming 147 core (Li et al., 2016; see Section 4.1) and outgassed into the atmosphere (Gaillard et al., 2022a). The 148 proportion of carbon being dissolved in the magma ocean (the Hadean BSE) was in comparison small 149 (Gaillard et al., 2022a,b). 150 Once pebble accretion ended, planetary growth occurred via collisions, a highly energetic process 151 capable of partly or totally ejecting the atmosphere into space (Chen and Jacobson, 2022). Since the magma 152 ocean atmosphere was primarily composed of carbon species, collisions mostly caused carbon loss. 153 Secondary atmospheres were subsequently formed by the degassing of the planetary interior after these 154 impacts. The primordial carbon delivered by chondrites has therefore been mostly lost to space or 155 partitioned into the core, which, in both cases, constitutes an irreversible mechanism, subtracting some 156 carbon from the BSE (Gaillard et al., 2021). Accordingly, non-carbonaceous (>3000 ppm C) and 157 carbonaceous (several wt.% C) chondrites actually contain more carbon than the BSE (Gaillard et al., 2021), 158 which has a carbon concentration ranging from 140 ppm (Hirschmann, 2018) to >500 ppm (Marty et al., 159 2013; 2016; see also Section 4.3). It has also been hypothesized that large impacts, such as that forming the 160 Moon, could have contributed a substantial amount of carbon and other volatiles to the BSE, as opposed to 161 facilitating overall carbon loss (Grewal et al., 2019). 162 On large bodies such as the Earth, a thick magma ocean could develop with high pressures 163 prevailing at its base. When pressure exceeds 15–20 GPa, ferric-iron-rich silicate melts (with Fe3+/Fetot 164 ~0.1–0.2) are stable in equilibrium with molten Fe-rich alloys (Armstrong et al., 2019). At lower pressure, 165 under otherwise similar conditions, such melts would contain only ~5% ferric iron. Vigorous convection 166 from the base to the top of the magma ocean implies that such ferric-iron-rich silicate melts would buffer a 167 high oxygen fugacity when exposed to surficial conditions. This implicitly assumes that the Fe-rich molten 168 alloys remain at the base of the magma ocean due to negative buoyancy forces. Consequently, while early 169 and thin magma oceans outgassed mostly CO, late and deep magma oceans outgassed mostly CO2. The 170 early redistribution of oxygen from CO2 at shallow conditions to Fe2O3 (i.e., Fe3+) in the deep magma ocean 171 may have caused the first diamond precipitation in the interior of Earth as illustrated in Figure 1 (Armstrong 172 et al., 2019). 173 7 The solidification of the magma ocean is a complicated stage, blending thermomechanical and 174 thermochemical processes (Ballmer et al., 2017; Hier-Majumder and Hirschmann, 2017; Korenaga, 2021). 175 It is yet unclear how magma ocean carbon could behave during this solidification. If a deep basal magma 176 ocean survived this process (Labrosse et al., 2007), we could expect substantial quantities of dissolved 177 carbon species to be sequestered therein. 178 Most models suggest that the atmosphere forming at the end of the magma ocean was CO2-179 dominated, with pressures ranging from 90–200 bar depending on the size of the degassing magma ocean 180 (Gaillard et al., 2022a; Sossi et al., 2020). Assuming water was efficiently degassed from the solidifying 181 mantle during the early Hadean, it has been proposed that its condensation in the form of early oceans could 182 have efficiently pumped in atmospheric carbon into the deep Earth via carbonate precipitation in a context 183 of early plate tectonics and vigorous rock alteration (Sleep et al., 2011; Korenaga, 2021). This implies an 184 early sequestration of deep carbon in the mantle via subduction of the surficial carbon. 185 A comparison between the mass of carbon outgassed by the magma ocean in the early Hadean and 186 the present-day mass of surficial carbon (atmosphere + oceans + crust) indicates surprisingly similar 187 numbers of 110±50×1021 g C (Gaillard et al., 2022a). This mass of surficial carbon is also indistinguishable 188 from the mass of carbon constituting the dense Venusian atmosphere. A similar conclusion has been 189 reached for nitrogen (Gaillard et al., 2022a), which can be explained by the very low fluxes of nitrogen 190 involved in both present-day planetary outgassing and burying via subduction with respect to the size of 191 the atmospheric reservoir (Gaillard et al., 2022a; Hirschmann, 2018). This may indicate that a surficial 192 reservoir of carbon has been formed early in the history of planet Earth and has been only moderately 193 affected by the subsequent 4.5 billion years of geological history. The balance of subduction versus 194 outgassing might have maintained a steady state regime, although there are still many uncertainties in the 195 actual fluxes of carbon involved in such geodynamic cycles (see Section 5; Dasgupta and Hirschmann, 196 2010; Iacono-Marziano et al., 2009; Kelemen and Manning, 2015) and it is difficult to discuss such 197 processes further. 198 Towards the end of Earth formation, a significant meteoritic bombardment — the so-called ‘late 199 veneer’ — delivered significant portions of highly siderophile elements to the Earth in chondritic 200 proportions, thereby elevating their BSE concentrations despite core differentiation (e.g., Walker, 2009; 201 Wang and Becker, 2013). A volatile-rich late veneer, as suggested by Wang and Becker (2013), could have 202 supplied a substantial proportion (20%–100%) of the carbon now comprising the BSE. However, 203 subsequent studies have cast doubt on the possibility of a solely late veneer contribution to BSE volatiles, 204 including carbon. A simple mass calculation has shown that the mass of material required to account for 205 8 the BSE carbon inventory (~2±1 wt.%) is equal to or greater than the mass delivered by the late veneer 206 (0.1–2.0 wt.%; Mikhail and Füri, 2019). Ruthenium, a highly siderophile element replenished in the BSE 207 by the late veneer, has an isotopic composition in Earth’s mantle distinct from carbonaceous chondrite-208 derived material; this suggests that the late veneer was unlikely to be composed of the volatile-rich 209 carbonaceous chondrites from which Earth’s carbon is derived (Fischer-Gödde and Kleine, 2017). Instead, 210 it is reasonable to suggest that the bulk of Earth’s carbon was obtained during the principal stages of 211 planetary accretion, as discussed above, with a minor fraction of carbon subsequently added by a late veneer 212 (e.g., Hirschmann, 2016). This is supported by isotope fractionation modeling of sulfur and chalcophile 213 elements, which suggests that planetary differentiation is sufficient to account for volatile element 214 concentrations in the BSE with less than 40% delivered by the late veneer (Wang et al., 2023). 215 Once the magma ocean was totally solidified, Earth volcanism was mostly mafic (i.e., forming 216 basaltic lavas), accompanied by superhot lavas such as Mg-rich komatiites (Figure 1, top). No constraint 217 allows us to discuss the carbon outgassed by these magmas. The oldest available constraints are probably 218 encapsulated within diamonds; diamonds older than 3.0 Ga host peridotitic inclusions, whereas younger 219 diamonds prevalently host eclogitic inclusions (Shirley and Richardson, 2011; see Section 5.5 for a full 220 discussion on diamond inclusions). This is possibly a signature of the initiation of plate tectonics at the end 221 of the Archean eon. 222 Cratons are thick (>200 km) and cold lithospheric blocks which have experienced little, if any, 223 geological activity since the Archean. They host kimberlites, which constitute the dominant form of 224 extrusive rocks containing diamonds (see Section 5.5.1). Mantle xenoliths found in kimberlites and in many 225 other intraplate lavas have revealed that mantle metasomatism, i.e., melt-fluid-rock interactions taking place 226 at the base of lithosphere (O’Reilly and Griffin, 2013), is an important process, modifying the physical and 227 compositional properties of cratonic mantle lithosphere and possibly storing an important amount of carbon 228 (see Section 4.3.3; Aulbach et al., 2017; Hirschmann, 2018). To some extent, lithospheric diamonds can be 229 seen as resulting from such metasomatism, and Ca-Mg carbonates have also been widely observed. A 230 hidden carbon reservoir might thus have been continuously formed over time (Hammouda et al., 2021) by 231 melt-fluid-rock interactions taking place in cratonic, non-cratonic, and even oceanic lithosphere (Aulbach 232 et al., 2017; Delpech et al., 2023; Foley and Fischer, 2017; Hammouda et al., 2021; Hirano and Machida, 233 2022; Hirschmann, 2018; O’Reilly and Griffin, 2013). 234 The final picture of the geological history of carbon corresponds to the present-day outgassing of 235 volcanoes — mid-ocean ridges, petit spots, hot spots, volcanic arcs, etc. (Figure 1) — which are sourced 236 by various reservoirs built by 4.5 billion years of geodynamic processes with a key role played by 237 9 subduction processes (Aiuppa et al., 2017; Dasgupta and Hirschmann, 2010; Hammouda et al., 2021; 238 Hirano and Machida, 2022; Plank and Langmuir, 1998). Carbon isotopic measurements of volcanic gasses 239 in arc settings suggest the major involvement of sedimentary carbonates in degassing, with the possibility 240 of crustal assimilation processes short-circuiting the deep carbon cycle (Section 5.4; Iacono-Marziano et 241 al., 2009; Mason et al., 2017).242 16 deep ocean; Middelburg, 2011). Nitrification, the oxidation of ammonia to nitrite and nitrate, is supposed 424 to be the dominant chemolithotrophic metabolism in the deep ocean, while diverse ecosystems (e.g., 425 hydrothermal vent, cold seeps, mud volcanoes, whale falls, oxygen minimum zones) host diverse 426 communities of chemolithotrophs, further contributing to the total rate of non-phototrophic biogenic carbon 427 fixation in the ocean (0.77 Gt C yr-1; Middelburg, 2011). 428 All the biogenic carbon that is not recycled back into inorganic carbon (by respiration or fermentation) 429 can eventually be buried under layers of sediment. There, the fate of biogenic carbon is an important 430 component of the process of diagenesis, the process by which sedimentary deposits are transformed into a 431 sedimentary rock over the years (Figure 2). This is largely influenced by local biotic processes (La Rowe 432 et al., 2020). While, in the upper layers of shallow sediments, aerobic respiration remineralizes biogenic 433 carbon, their deeper layers, and, a fortiori, all deep-ocean sediments, are mostly subjected to anaerobic 434 respiration and fermentation processes. The most prevalent one is sulfate reduction (Canfield and Des 435 Marais, 1991). Sulfate-reducing microorganisms use the abundant oceanic sulfate as an electron acceptor 436 to oxidize biogenic carbon (Jørgensen et al., 2019). Interestingly, sulfate is not the most favorable oxidizer 437 thermodynamically (as compared to nitrate or ferric oxide), however its greater abundance in seawater 438 makes it a dominant electron acceptor in the anaerobic oxidation of biogenic carbon in marine sediments. 439 Deeper below, where sulfate is scarce, another metabolism, methanogenesis, is prevalent (e.g., Bojanova et 440 al., 2023). Methanogens reduce organic matter or CO2 into methane, releasing biogenic carbon that can 441 seep up. Between the sulfate and methane zones lies the sulfate-methane transition zone (SMTZ), where 442 both species cohabit. Within the SMTZ, a third metabolism, the anaerobic oxidation of methane, is 443 dominant. Anaerobic oxidation of methane uses the sulfate to oxidize methane produced below by 444 methanogens, thereby releasing CO2 and hydrogen sulfide. Stable isotope analysis can track these stratified 445 metabolisms, as each reaction leaves a characteristic isotopic signature, yielding a range of signatures in 446 diagenetic samples (Meister, 2019). 447 448 A part of biogenic carbon in sediment remains unalterable. Chemically, intermolecular cross-links 449 (Kunhi Mouvenchery et al., 2012) or abiotic modifications, such as sulfurization or peptide deamination 450 (Abdulla et al., 2018; Burdige, 2005), decrease the bioavailability of the organic matter in the sediments. 451 Physically, the phenomenon of sorption of organic matter to mineral surfaces also protects biogenic carbon 452 remineralization by encapsulating it, denaturing extracellular enzymes, and limiting molecular diffusion. 453 In the long run, the quantity of biogenic carbon buried therefore depends on the balance between a 454 continuous input from sedimenting biological matter and its oxidation by local microbial life, in parallel 455 17 with the ongoing process of sorption and the abiotic alteration of the matter, protecting this carbon from 456 further oxidation. 457 Finally, another component of the deep carbon cycle is the ‘carbonate pump’ which contributes to 458 carbon sequestration in the subsurface. Calcifying organisms — especially coccolithophores — use 459 carbonate ions in seawater to build calcite platelets called coccoliths. Upon their death, these organisms 460 contribute to the accumulation of biogenic carbonate sediments on the ocean floor. Over geological 461 timescales, these sediments can undergo lithification, forming carbonate rocks such as limestone (Folk, 462 1980). 463 Of the 10 Gt of biogenic organic carbon transported from the surface to the deep ocean annually, only 464 0.02–0.2 Gt C yr-1 is eventually incorporated in ocean sediments (Cartapanis et al., 2018, Dunne et al., 465 2007, Hayes et al., 2021, Jahnke, 1996). Most of this biogenic carbon is therefore accumulating at various 466 depths until it is eventually respired, effectively sequestering it away from surface ocean and atmospheric 467 exchange over timescales ranging from months to millennia, depending on the depth. 468 469 3.5. Effect of life on carbon fluxes to and from the deep Earth 470 Over geological timescales, organic-matter-rich sediments can be subducted into the Earth's mantle 471 along convergent plate boundaries (Figure 2; Section 5.1.1). Microbial communities can survive in 472 sediments that are subducted to depths of several kilometers, where they continue to influence carbon 473 transformations (Barry et al., 2019; Debret et al., 2022; Fryer et al., 2020; Plümper et al., 2017). In 474 particular, these communities can rely on the interaction with fluids associated with serpentinizing forearc 475 mantle: reduced volatile species from these fluids (e.g., H2, CH4) encounter more oxidized redox couples 476 from the subducted crust, thereby supporting deep chemolithotrophic life (Mottl et al., 2003; Ohara et al., 477 2012; Vitale Brovarone et al., 2020a). Chemolithotrophy leads to the formation of partially reduced organic 478 carbon compounds (e.g., lipids, carbohydrates, proteins, hydrocarbons), which may either be subducted 479 further into the deep mantle or, heated under high-pressure conditions, can decompose and release 480 remineralized carbon (CO2, CH4). 481 The subduction of biogenic carbon has significant implications for volcanic activity and the global 482 carbon cycle. Organic carbon and carbonates carried into subduction zones can decrease the melting point 483 of the mantle, leading to enhanced magma generation (Plank and Manning, 2019; see Section 5.4). This 484 18 magma, when it reaches the Earth's surface through volcanic eruptions, releases substantial amounts of CO2 485 back into the atmosphere together with variable amounts of reduced carbon compounds (e.g., Giggenbach, 486 1996; Symonds et al., 1994). Degassing and rising fluids containing reduced carbon species (and other 487 reduced compounds, e.g., H2, NH3, H2S) have the potential to fuel subsurface life on their way to the surface, 488 as evidenced by the abundance of chemolithotrophic microbial communities thriving in various volcanic 489 and gas-seeping ecosystems. Information regarding the quantitative impact of subsurface life on the carbon 490 recycled to the surface are limited, but suggests that potentially large amounts of the carbon compounds 491 (and other volatiles) recycled to the surface might be altered and sequestered by the subsurface biosphere 492 in the overriding crust. For example, recent work suggests that subsurface microbial communities present 493 in the forearc region of convergent margins can sequester into biomass up to 90% of the carbon recycled to 494 the surface, reducing the carbon going to the mantle by up to 40% (Barry et al., 2019, Fullerton et al., 2021). 495 Dissolved organic carbon in the sampled forearc springs showed δ13C signatures indicative of 496 chemolithotrophic activity utilizing CO2 originating from the subducting plate (Fullerton et al., 2021). 497 Additionally, near surface microbial communities associated with areas where reduced fluids and gas are 498 recycled to the surface utilize these compounds for their metabolism, further influencing the diversity and 499 fate of carbon molecules and impacting the carbon cycle. In this context, the interactions between 500 subsurface life and the solid Earth play a significant role in the deep carbon cycle, impacting and altering 501 the flux between the Earth's interior and surface. Constraining the contribution of life in altering the quality 502 and quantity of carbon recycled to the surface will be of key importance to understand the interactions and 503 feedbacks between biology and the solid Earth504 19 4. Deep carbon forms 505 The deep forms of carbon are diverse and include multiple redox states and coordination numbers 506 depending on their origin and location within the deep Earth (Figure 3). The most common solid carbon 507 forms are carbonate minerals, organic matter, graphite, diamond, and carbides. Their movement is largely 508 controlled by redox changes and solubility. This section covers the deep reservoirs of the Earth from the 509 core to the crust, passing through the mantle. 510 511 4.1. Carbon in the Earth’s core 512 Earth’s core is located 2890 km below the surface and is primarily made of iron (85%) and nickel 513 (5%). Small amounts of lighter siderophile elements, including hydrogen, oxygen, sulfur, silicon, and 514 carbon are mandatory to match the density deficit and velocity excess of the Preliminary Reference Earth 515 Model (PREM; Dziewonski and Anderson, 1981). As the Earth’s core constitutes 32% of the mass of our 516 planet, even minute amounts of carbon in the core (between 0.1–0.5 wt.% C depending on estimates) make 517 it the largest terrestrial carbon reservoir, containing approximately 90% of the bulk Earth carbon budget 518 (see Fischer et al., 2020, and references therein). 519 As the Earth’s core is not directly accessible, estimates of its carbon content rely on our 520 interpretation of various observations. The seismic velocity and elastic parameters of the Earth’s core 521 indicate that it is ∼10% less dense than iron-nickel alloy under the core pressure-temperature range (136–522 364 GPa and 4000 K at the core-mantle boundary, CMB) primarily due to the presence of light elements. 523 Comparison of elemental concentrations in BSE relative to primitive meteorites and the volatility trend of 524 elements indicates that the core is enriched in sulfur, hydrogen, and carbon as far as light volatile elements 525 are concerned. Isotopic fractionation during core formation leads to one of the highest estimates of core 526 carbon concentrations (0.5–2.0 wt.%; Wood et al., 2013). 527 The relative enrichment of the core in carbon relative to the BSE was acquired when the core 528 segregated from the mantle silicate reservoir during the collisional accretion stage of the Earth, which 529 extended 10–100 million years after the initial solar nebula processes (see Section 2; Barboni et al., 2017; 530 Jacobson et al., 2014). The metal delivered to the growing Earth that ultimately formed the core existed as 531 droplets that equilibrated at least partially with the magma ocean, while themselves sinking due to the higher 532 density of the metal phase relative to silicate melt. The environment, depth, and redox conditions under 533 20 which carbon partitioned between silicate and metal phases largely dictated the carbon content of the 534 Earth’s core. At low to moderate pressure and high temperature, carbon is siderophile. Our modern societies 535 have largely taken advantage of this behavior, expanding our understanding of carbon partitioning behavior 536 in metals through the production and intensive use of steel (i.e., iron carbide), and developing adjustable 537 properties depending on the addition of other siderophile elements (e.g., chromium and often nickel to 538 obtain stainless steel). The high cosmochemical abundance of carbon, coupled with its low BSE abundance 539 (0.01%), has led to the suggestion that carbon dissolved extensively into the core, and that its concentration 540 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 541 inner core, which may contain high-pressure carbides such as Fe3C or Fe7C3 (Wood, 1993; Dasgupta et al., 542 2013a). Therefore, Fe-C alloy binaries have been extensively investigated under extreme pressure-543 temperature conditions. Their density and elastic properties were measured, as a viable composition model 544 of the core must account for the density deficit and seismic characteristics of the core. For the inner core, 545 both experiments and theory suggest that carbon and hydrogen are key impurity elements to explain the 546 Poisson ratio (Vp/Vs) observed in the inner core as compared to pure iron. Among candidate Fe-C alloys 547 and carbides, Fe7C3 is a good candidate from its ability to match the anomalously low Vs and high Poisson 548 ratio of the inner core, in addition to reproducing the density deficit and having the highest electrical 549 resistivity among all Fe-C alloys (C. Zhang et al., 2018). A caveat is its density, which may be too low to 550 match the inner core density, and would more generally imply a high carbon concentration in the core. This 551 contradicts experiments performed at the pressures and temperatures expected in a deep magma ocean 552 during formation of Earth’s core which demonstrate that carbon becomes much less siderophile at depth 553 (Fischer et al., 2020). Fischer et al. estimated that the core likely contains a maximum of 0.09±0.04 to 554 0.20±0.10 wt.% C, which is in agreement with most of the theoretical atomistic and experimental mineral 555 physics studies that were performed at or close to core pressures. 556 As the composition of the core is central to our understanding of the Earth’s accretion and the 557 terrestrial magnetic field, hundreds of studies focusing on mineral physics, cosmochemical constraints, or 558 geochemical experiments have attempted to decipher the composition of the Earth’s inner or outer core and 559 provide a match with the seismic properties of the core. Depending on the chosen theoretical or 560 experimental parameters and strategies, the pressure-temperature range covered by these studies, and the 561 cocktail of light elements considered, results vary. Recent results still propose that the carbon concentration 562 in the outer core spans one order of magnitude between 0.1 and 2.0 wt.% (Bajgain et al, 2021). Existing 563 data are yet insufficient to resolve the competing models of core composition because of limited data 564 coverage in the relevant challenging pressure-temperature-composition space and uncertainties in 565 experimental measurements and theoretical simulations. In their recent review, Hirose et al. (2021) propose 566 21 a likely range of composition for the outer and inner core, with approximately 0.2 wt.% and 0–1.3 wt.% C 567 respectively, depending on the concentration of other light elements. 568 Future progress and better constraints on the core composition and carbon budget could arise from 569 the investigation of the interactions between melt and solid at the inner core boundary (ICB), as techniques 570 and methods progressively allow us to reach ICB pressure (330 GPa); ICB temperatures, which may be as 571 high as 7000 K, have not yet been attained. This would inform us on the evolution of Earth's magnetic field 572 through time. The thermal and compositional evolution of Earth’s core is related to the growth of the inner 573 core. Since light elements control the liquidus temperature of Fe alloy and therefore the crystallization of 574 the core during secular cooling, they could substantially increase the thermal conductivity of the outer core 575 with profound implications for core thermal evolution and geodynamics. Inner core viscosity and dynamics 576 are also certainly dependent on the light element composition through diffusion rate, grain growth rate and 577 shear modulus, with virtually nothing known at this stage. 578 Moving upwards through the Earth to the outermost part of the core, seismology has revealed a 579 ~400 km thick low-velocity and low-density layer named the E’ layer (Figure 3), which is characterized by 580 a steep Vp gradient (Brodholt and Badro, 2017; Helffrich and Kaneshima, 2010; Kaneshima, 2018; 581 Kaneshima and Helffrich, 2013; Kaneshima and Matsuzawa, 2015). The E’ layer is interpreted as a remnant 582 of the chemical exchange between the early formed core and the whole-Earth magma ocean. The stagnation 583 of such a layer over time necessitates buoyancy relative to the underlying convecting outer core, which 584 could be explained by an enrichment in oxygen and depletion in silicon, but probably not at the 100-km 585 scale. Presence of a significant amount of carbon in the E’ layer is unlikely, as adding carbon slightly 586 decreases the density of liquid iron and increases its Vp (Morard et al., 2017; Nakajima et al., 2015). 587 Despite the challenges to reach the pressure and temperature conditions characteristic of the CMB 588 and deeper, the knowledge of the physical properties of Fe alloys and melts including carbon and other light 589 elements has incredibly improved over the last decades. Even for the lowest estimated carbon contents 590 (~0.2 wt.%), the largest carbon reservoir of planet Earth is its core, which hosts at least 90% of the whole 591 Earth carbon budget and is where carbon has been sequestered since the final stages of core formation. 592 593 4.2. Carbon at the interface between the core and the lowermost mantle 594 The lowermost mantle above the CMB, referred to as the D” layer, displays multiple seismic 595 features (Figure 3). It is also a strong thermal boundary layer, with temperatures increasing from an average 596 22 mantle adiabat of about 2500 K at 2600 km depth (e.g., Stixrude et al., 2009) to a CMB temperature of 597 about 4000 K. The D” layer is both vertically and laterally highly heterogeneous, and it is layered on a 598 kilometer scale in the lower 50 km (Sidorin et al., 1999; Thybo et al., 2003). Because of diffraction of 599 seismic waves by the core, the resolution in this layer is lower than in the overlying mantle and hinders 600 access to fine details. Velocity profiles show large lateral heterogeneities in this region. Estimates of the 601 thickness of the D” layer suggest that it ranges from 100 km to about 400 km. Within this layer, unusually 602 low wavespeeds have been measured beneath the African continent and the Pacific Ocean, and these regions 603 are known as large low-shear-velocity provinces (LLSVPs) and exhibit a density excess of about 1.25% 604 compared to the surrounding mantle. Most active volcanic hotspots on Earth’s surface are located above 605 and at the boundaries of LLSVPs. 606 Adjacent to the LLSVPs, seismic observations show ultra-low velocity zones (ULVZs), often in 607 contact with the CMB. ULVZs are not ubiquitous. They also have a density excess and a low seismic 608 velocity feature as indicated by their name. Their Vp and Vs reduction is about 10% and 30%, respectively 609 (e.g., Lay, 2005). Although the presence of ULVZs and LLSVPs is not questioned, their origin and 610 composition remain uncertain and there are multiple conceptual models that are consistent with seismic 611 tomography results and the geochemistry of the hotspot lavas of oceanic islands, which are commonly 612 assumed to result from the partial melting of large rising mantle plumes deeply rooted in the LLSVPs and 613 ULVZs (see reviews by Garnero et al., 2016; Weis et al., 2023). 614 The ULVZs may represent partially leaky ‘windows’ between the D” layer and the outer core. Such 615 localized core-mantle exchanges are supported by geochemical investigations of plume-related volcanic 616 rocks. Recent isotopic measurements have revealed a negative correlation between the 182W/184W and 617 3He/4He ratios in major plume-related oceanic island basalts (OIB) suites (see review by Weis et al., 2023). 618 Plumes giving rise to OIBs with low 182W/184W ratios might have sampled core metal, most likely via some 619 ULVZs located in the root zones of plumes (Rizo et al., 2019). 620 The location of LLSVPs and ULVZs between regions of sinking tectonic plates is intimately 621 associated with the long-term dynamics of the convecting mantle, although it is unclear whether they are 622 inherited from the early evolution of the mantle and related with remnants of the deep magma ocean, or 623 they developed over time (e.g., McNamara, 2019; Flament et al., 2022). As long-lived structures, stable or 624 slowly mobile at the CMB over hundreds of millions of years, they will be negatively buoyant relative to 625 the convecting mantle, and their composition is likely correlated to the observations of their increased 626 intrinsic density and seismic velocity reduction. 627 23 Among the multiple models of LLSVPs and ULVZs, we focus here on the potential role of carbon-628 bearing phases that may explain the density excess and low shear wave pattern of LLSVPs. Enhanced iron 629 content in deep mantle mineral assemblages is the easiest way to account for the density excess and low 630 velocity of LLSVPs and ULVZs, and low partial melting could also explain the low-velocity features. Liu 631 et al. (2016) proposed that metallic Fe-C produced in subducted slabs beyond a depth of 250 km may 632 partially melt at CMB conditions, since the eutectic melting curve of the Fe-C system crosses the geotherm 633 near Earth’s CMB. At the eutectic, the Fe-C melt could contain 2 wt.% C at the P-T conditions of the CMB 634 (Fei and Brosh, 2014; Lord et al., 2009). If concentrated into isolated patches, such a metallic melt could 635 induce the seismically observed density and velocity features of ULVZs as it is twice as dense as the 636 surrounding mantle. These hypotheses connect peculiar features near Earth’s CMB to subduction of the 637 oceanic lithosphere through the deep carbon cycle. Depending on the melt fraction and wetting angle, Liu 638 et al. (2016) distinguished three types of dynamics of the Fe-C melt in ULVZs that could have different 639 seismic signatures, and may explain the laterally varying electrical conductivity in the D” layer: a) ULVZs 640 containing solid phases that have become iron-rich through reaction with Fe-C melt; b) ULVZs containing 641 9–16 vol.% non-wetting Fe-C melt coexisting with a small degree of silicate melt in a solid silicate matrix; 642 c) ULVZs containing 5–11 vol.% Fe-C melt that wets the solid silicate matrix and will ultimately merge 643 with the core, explaining some of the W isotopic signature sampled by OIBs (Rizo et al., 2019). 644 The seismic features of the hot and dense LLSVPs, including seismic anisotropy, are generally 645 consistent with the presence of bridgmanite, CaSiO3-perovskite (or Davemaoïte; Thomson et al., 2019). 646 They have recently appeared to be also consistent with the high-pressure polymorphs of CaC2O5 (Wang et 647 al., 2024), with a structure that is based on CO4 groups (tetracarbonates) with sp3-hybridized carbon. In 648 particular the symmetry change from orthorhombic Fdd2 to monoclinic C2 leads to a density increase of 649 5.8% and a Vs decrease of 7.4%, and to elastic parameters of CaC2O5-C2 consistent with seismic 650 observations within LLSVPs. The orthorhombic to monoclinic transition in CaC2O5 occurs at 80 GPa, 651 broadly consistent with the top of LLSVP piles. More generally, the elastic properties of the polymorphs 652 of CaC2O5 that may form by the reaction between CaCO3 and CO2 in deep subducting slabs from transition 653 zone conditions and deeper are in good agreement with PREM. This would connect the deep carbon cycle 654 from the oceanic crust to the lowermost mantle and back to OIBs. Given similarities in elastic properties 655 between CaSiO3-perovskite and CaC2O5-C2, it may be challenging to evaluate the actual amount of 656 oxidized carbon present in LLSVPs. Their fate may be better estimated through their isotopic signature in 657 OIB suites. 658 Should there be remnants of the basal magma ocean, sp3 hybridized tetracarbonates are less siderophile, 659 and thus large amounts of it may be stranded instead in the deep mantle. These tetracarbonates may 660 24 therefore provide a possible source for carbon-rich emissions registered at the surface in OIB suites 661 (Cerantola et al., 2023). 662 663 4.3. Carbon throughout the Earth’s mantle 664 Carbon is a trace element in the Earth’s convecting mantle. Estimates of carbon content in the bulk 665 silicate Earth generally reconstructed from noble gas content are estimated to vary between 500 and 1000 666 ppm (Marty et al. 2013, 2016). Ratios of C with elements that partition in a similar manner in the mantle, 667 such as Ba or Nb, yield a BSE estimate of 110±40 ppm for the mantle, and 140±40 ppm for the BSE 668 (Hirschmann, 2018). When estimated from elements that behave geochemically like carbon but are 669 refractory in a cosmochemical sense, concentrations of ~192 ppm and ~350 ppm C are derived from U and 670 Ba respectively (Lee et al., 2019). 671 The concentration of carbon in the lower and upper mantle is substantially different. Measured carbon 672 fluxes near hot spots are much higher than those for the depleted mid-ocean ridge basalt (MORB)-source 673 mantle indicative of a higher concentration of carbon in the lower mantle. The most comprehensive and 674 detailed analysis of the carbon content in MORBs showed very-fine-scale variations in their mantle source, 675 with carbon fluxes from the mantle to the surface varying by over three orders of magnitude at the segment 676 scale along mid-ocean ridges (Hauri et al., 2019). Integration over the entire length of ridges yields a global 677 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 678 C in the depleted mantle source of MORB (Cottrell et al., 2019, Hauri et al., 2019; Le Voyer et al., 2018). 679 Measurement of carbon solubility in the major minerals of the upper mantle, transition zone and lower 680 mantle shows that carbon is virtually insoluble in the silicate minerals, at the level of a ppm or less (Keppler 681 et al., 2003; Shcheka et al., 2006). This implicitly requires that carbon is stored in accessory, nominally 682 carbon-rich phases such as elemental carbon in the form of graphite or diamond, carbides, hydrocarbons in 683 regions rich in hydrogen or carbonates, with oxidation numbers ranging between -4 and +4 depending on 684 pressure, temperature and oxygen fugacity (fO2). As such, the diversity of carbon-rich phases is diagnostic 685 of the vertical and lateral heterogeneities in the convecting mantle. Over geological time, the forms of 686 carbon-rich phases are controlled by the degree of oxygenation of the mantle and are diagnostic of its 687 evolution through time. 688 689 25 4.3.1. Reduced lower mantle and transition zone 690 Rock samples from depths of 300 km and below are scarce; the deeper, the scarcer. Observations 691 rely almost exclusively on inclusions within the few super-deep diamonds that have been collected through 692 time in cratons (see Section 5). Observations of natural samples are complemented by experimental and 693 theoretical studies to assess the redox state and carbon-hosting phases in the deep mantle. 694 The deepest natural messengers from the mantle are the ‘CLIPPIR’ diamonds (Cullinan-like, Large, 695 Inclusion Poor, Pure, Irregular, and Resorbed) that formed at lower mantle conditions and were brought 696 back to the surface (see Shirey et al. 2013, 2019, and references therein). Their most common inclusion is 697 a composite metallic Fe-Ni-C-S mixture. The metallic inclusions are comprised of cohenite (Fe,Ni)3C, 698 interstitial Fe-Ni alloy, segregations of Fe-sulfide, and minor occasional Fe-Cr-oxide, Fe-oxide, and Fe-699 phosphate. A thin fluid layer of CH4 and lesser H2 is trapped at the interface between the inclusion and 700 surrounding host diamond. The mixture is commonly interpreted to have been trapped as a molten metallic 701 liquid. Their depth is bracketed to depth between 360–750 km by inclusions containing high-pressure 702 silicate minerals such as retrogressed CaSiO3-perovskite (Davemaoïte) or majoritic garnet. 703 The Fe-Ni-C-S inclusions in CLIPPIR diamonds are unique physical samples of an Fe-rich metallic 704 liquid from the deep mantle, which confirm the fundamental process of Fe2+ disproportionation at depth, 705 also called ‘redox freezing’. Indeed, the Earth’s mantle oxygen fugacity is buffered by ferric-ferrous (Fe3+-706 Fe2+) equilibria and decreases with increasing depth. This has been measured in natural garnet peridotites 707 (Woodland and Koch, 2003), and confirmed by experiments that tested garnet oxythermobarometry 708 equilibria under relevant high-pressure conditions (Stagno et al., 2013, and references therein). Theoretical 709 analysis of the mantle fO2 evolution with increasing depth (Ballhaus, 1995; Ballhaus and Frost, 1994) 710 predicts that an isochemical mantle composition becomes reduced by ∼0.6 log units in fO2 per GPa as a 711 result of molar volume changes of redox equilibria and changing mantle mineralogy. At lower mantle 712 conditions, experimental studies have shown that aluminous bridgmanite and majoritic garnets are sinks 713 for ferric iron, which in turns forces ferrous iron disproportionation into ferric silicates and Fe-metal (e.g., 714 McCammon, 1997; Frost and McCammon, 2008). This is thought to generate up to about 1 wt.% metal in 715 the lower mantle and establishes a carbon-reducing environment in much of the deep mantle, where the 716 forms of carbon would be diamond, metal and carbides (Figure 3). 717 Experimental work has also assessed the relevant carbon phases in the lower mantle. At the base 718 of the upper mantle, the Fe-Ni-C solidus at 10 GPa is low (1150–1250 °C) compared to mantle geotherms, 719 which may indicate that Fe-Ni-C alloys are stable as molten phases, potentially limited in size to the grain 720 scale. Cohenite [Fe,Ni]3C observed in the CLIPPIR diamonds likely forms together with metal and diamond 721 32 incompatible trace element ratios) is sufficient to account for the considerable fluxes of diffuse CO2 910 degassing at continental rifts (Iddon and Edmonds, 2020; Wong et al., 2023). 911 912 4.3.4. Sub-oceanic lithospheric mantle 913 Carbon may also be sequestered within sub-oceanic lithospheric mantle. At mid-ocean ridges, 914 asthenospheric carbon mobilized by mantle melting is focused towards the ridge axis, where a significant 915 quantity degasses into the water column. In addition, a portion of the mobilized carbon is directed towards 916 the flanks of the spreading ridge, where it is sequestered through the metasomatism of the juvenile sub-917 oceanic lithospheric mantle (Keller et al., 2017). As oceanic plates age, their underlying mantle lithosphere 918 will thicken, and by doing so it will accumulate additional small-fraction carbon-rich asthenospheric melts 919 in the developing lithosphere. Furthermore, xenoliths collected from the base of the Pacific plate suggest 920 that later extensional stresses in the lifetime of the oceanic plate such as plate flexure may induce further 921 metasomatism (Pilet et al., 2016). Carbon present in the oceanic lithosphere in this manner is believed to 922 contribute to observed seismic low-velocity zones and velocity contrasts (Keller et al., 2017). The controls 923 on CO2 sequestration within oceanic lithosphere include mantle temperature, mantle CO2 content, and melt 924 productivity/fertility (Keller et al., 2017), which are known to vary substantially across the present-day 925 ridge system (e.g., Le Voyer et al., 2019). In combination with the difficulty in accessing the lower oceanic 926 lithosphere, there is likely to be substantial uncertainty in quantifying the carbon budget of the lithosphere 927 underlying oceanic plates. Following geodynamic models and tomographic estimates of lithospheric mass, 928 recent first order estimates suggest that 0.5–2.2×1021 g C may be stored within present-day unaltered sub-929 oceanic lithospheric mantle (Gibson and McKenzie, 2023; Müller et al., 2022). 930 931 4.4. Deep carbon in the crust 932 In the continental crust, solid carbon will take multiple different forms depending on the pressure, 933 geothermal gradient, and oxidation state of the local environment (e.g., Hazen et al., 2013a; Oganov et al., 934 2013). Voluminous solid continental carbon phases can be broadly subdivided into two principal forms. 935 The largest fraction of crustal carbon is represented by inorganic carbon in carbonate minerals, in which 936 oxidized carbon is present within the bicarbonate anion (CO32-) bonded in mineral structures to a wide 937 selection of cations, which account for about 80% of total crustal carbon (Berner, 2004; Hayes and 938 33 Waldbauer, 2006). The second largest form of solid crustal carbon, accounting for about 20% of total 939 carbon, is organic matter, in which carbon is present in its reduced form (Hazen et al., 2013a). It includes 940 biogenic organic matter, petrogenic or detrital graphitic carbon in sedimentary rocks, and abiotic forms of 941 condensed reduced carbon (Beyssac and Rumble, 2014; Galy et al., 2008; Luque et al., 1998). 942 A large fraction of crustal solid carbon, both organic and inorganic, in the crust formed through biological 943 processes and their successive solid-state transformations. Carbonate minerals can be synthesized by 944 various organisms that grow a carbonate shell, or result from biology mediated chemical reactions. Over 945 the past 150 million years, Deposition of calcareous plankton in deep-sea sediments, mainly coccoliths and 946 and foraminifera, has increased over the past 150 million years and at present accounts alone for more than 947 half of global carbonate sedimentation (Berner 2004; Milliman, 1974). Organic carbon is primarily 948 produced by biology and accumulates in rocks during sedimentary processes and is progressively 949 transformed to carbonaceous materials and graphitic carbon by diagenetical and metamorphic processes. 950 Carbonate minerals and organic matter can also form abiotically in the crust, through multiple processes 951 ranging from silicate weathering at the Earth’s surface, to deep crustal fluid-rock interactions. The 952 following sections describe the forms and distribution of carbon in the main crustal reservoirs. 953 954 4.4.1. Continental crust 955 The continental crust is an important reservoir of geological carbon. Recent estimates derived from 956 statistical analysis of lithological distributions suggest that ~20% of present-day continental crust is 957 composed of carbonate rocks (Walton and Shorttle, 2024), which operate as the largest store of crustal 958 carbon at the present-day (Hirschmann, 2018). However, the distributions and forms of carbon on and 959 within present-day and geologically historic continental crust is highly heterogeneous, complicating our 960 capacity to estimate the carbon budget of the crust. Here, we summarize the current understanding of carbon 961 reservoirs in the continental crust. 962 Crustal carbonate and organic matter are distributed within both the sedimentary and crystalline 963 materials that comprise the bulk continental crust (Hirschmann, 2018). Of these two carbon forms, 964 carbonate is perhaps more commonly analyzed, whilst data on reduced carbon is sparse in comparison 965 (Derry, 2014; Hartmann et al., 2012). Carbon on continental crust dates back to 4.0 Ga, coinciding with 966 some of the earliest records of life within sediments (Schidlowski, 1988). Primordial Earth had a primarily 967 silicate crust, which has been gradually carbonated over time as a consequence of continuous deep-Earth 968 34 degassing (Alcott et al., 2024). From present-day records of lithological preservation, it is estimated that 969 ~5% of total preserved Precambrian crustal volume is comprised of carbonates, which are principally 970 biogenic stromatolitic and abiotic precipitated carbonates deposited in oceanic environments and preserved 971 on continental crust (Walton and Shorttle, 2024). After the Cambrian and Ordovician radiations of 972 biomineralising marine life (Gilbert et al., 2022), there is a substantial increase in the proportion of 973 carbonate preserved on continental crust (~20–25%); similarly, a decrease in the fraction of accumulated 974 carbonates is recorded within crustal sections in the aftermath of the Permian-Triassic extinction (Walton 975 and Shorttle, 2024). This ongoing carbonate accumulation within Earth’s continental crust has had 976 substantial consequences for the co-evolution of life and the geosphere and atmosphere (e.g., Alcott et al., 977 2024; Duncan and Dasgupta, 2017, Isson et al., 2019). The abundance of carbonate present with the bulk 978 Earth crust is therefore both affected by and in turn affects surface biogeochemical cycles (see Section 3). 979 In the continental lithosphere (crust and uppermost non-convecting mantle), the principal form 980 hosting carbon is carbonate minerals (Hazen et al., 2013a; Oganov et al., 2013). Of these minerals, the most 981 abundant carbonate minerals in the Earth’s crust are calcite (β-CaCO3) and dolomite (CaMg(CO3)2), which 982 are present within massive sedimentary and metamorphic carbonate formations that account for 90% of 983 crustal carbonates (Reeder, 1983). Calcite is additionally one of the most common cement minerals found 984 within siliciclastic sedimentary rocks (Manning et al., 2013). Aragonite (λ-CaCO3), the orthorhombic 985 polymorph of the trigonal calcite, is a metastable form of carbonate which principally originates from 986 biomineralization, and is therefore found in substantial proportions on continental shelves where it is 987 sourced from the dissolved bicarbonate anion in seawater by marine biomineralizers (e.g., Gilbert et al., 988 2022). Other key but minor carbonate minerals include magnesite (MgCO3), which forms through the 989 carbonation of Mg-rich igneous and metamorphic rocks or through authigenic precipitation (e.g., Scheller 990 et al., 2021), and siderite (FeCO3), a principal component of Precambrian banded iron formations (e.g., 991 Klein, 2005). At the time of writing, 614 carbonate minerals are registered by the online mineral database 992 mindat.org (i.e., filtering by ‘CO32-’; retrieved June 2024), illustrating the broad and diverse nature of 993 carbonate minerals on Earth at the present day (Hazen et al., 2013b). Carbonates are also generated through 994 chemical weathering of silicate rocks, which consumes atmospheric CO2 (Gaillardet et al., 1999; Hartmann 995 et al., 2012; Stewart et al., 2019; Urey, 1952) and has hence been correlated to long-term climate modulation 996 (e.g., Gernon et al., 2021; Johansson et al., 2018). 997 Oxidized carbon in crystalline continental basement is present in igneous and metamorphic rocks 998 as carbonate minerals or as fluid inclusions (e.g., Lowenstern, 2001; Skelton, 2011). In magmatic systems, 999 mobile carbon, present as a vapor phase or dissolved within silicate magmas, typically takes the form of 1000 35 CO2 (typically 1–50 mol%), with its more reduced gaseous forms CO and CH4 much lower in concentration 1001 (typically ≪1 mol%; Giggenbach, 1996; Symonds et al., 1994). While carbon concentrations are enhanced 1002 in peralkaline/alkaline compositions (Ni and Keppler, 2013), CO2 is highly insoluble in silicate magmas 1003 (e.g., Dixon et al., 1995), and will degas deep within the Earth’s crust to form fluid phases (Section 5.3). 1004 The majority of this CO2 is expected to be degassed into the atmosphere or oceans, although magmatic 1005 carbon-bearing fluids, among other carbon sources, may contribute to the formation of hydrothermal solids 1006 (e.g., Bénézeth et al., 2013). An additional minor reservoir of crustal carbon is in rare low-melt-fraction 1007 volcanic rocks, such as carbonatites, nephelinites, and kimberlites (e.g., Jones et al., 2013). 1008 In near-surface environments, reduced forms of carbon (i.e., C0 and hydrocarbons) are attributed to 1009 biological processes, and are hence commonly referred to as simply ‘organic carbon’ (Derry, 2014). This 1010 carbon is present almost exclusively within sediments, such as deep-sea/lacustrine shales or coals, and are 1011 derived from organic material resistant to microbial processing (e.g., microbial membranes or plant lignin; 1012 Manning et al., 2013). Burial of organic material in marine environments is the second largest sink of 1013 atmospheric CO2 after silicate weathering and carbonate precipitation, a process that is enhanced by fluvial 1014 transport of organic carbon from continents to oceans (Galy et al., 2008). At higher crustal pressures, 1015 reduced and dehydrated carbon in sediments takes the form of graphite (Hazen et al., 2013a), which is a 1016 key contributor to the electrical conductivity of continental crust (Glover, 1996). Graphite may also form 1017 within some igneous and metamorphic rocks as an accessory phase precipitated from reduced C-O-H fluids 1018 (e.g., Luque et al., 1998). Under specific circumstances, organic carbon in the crust will undergo thermal 1019 maturation to generate biogenically derived hydrocarbons (e.g., Sephton and Hazen, 2013). Further thermal 1020 maturation will result in the decarboxylation and dehydrogenation of organic compounds, resulting in the 1021 gradual oxidation and aromatisation of organic hydrocarbons to conclude in graphite (see Section 4.5 for a 1022 full discussion on deep abiotic organics). 1023 Carbon in the continental crust will be remobilized during metamorphism in various forms such as 1024 CO2, CH4, or more complex ionic species (e.g., Manning et al., 2013; Sverjensky et al., 2020; Section 5). 1025 Destabilization of solid carbon during metamorphism may result from a number of tectonic or volcanic 1026 processes that result in disequilibrium pressure-temperature conditions, including volcanism (e.g., Black 1027 and Gibson, 2019; Ganino and Arndt, 2009; Johnston et al., 2011; Mason et al., 2017), localized microbial 1028 and thermogenic breakdown to form CH4 (Section 5.2.2.2; e.g., Etiope et al., 2019; Etiope and Sherwood 1029 Lollar, 2013), and regional metamorphism (Section 5.2.2; Ague, 2000; Becker et al., 2008; Groppo et al., 1030 2013, 2017; Kerrick and Caldeira, 1998; Stewart and Ague, 2018;). As a result, the lengths of tectonic 1031 settings, in particular those of arcs and subduction zones, are believed to be correlated with periods of 1032 36 icehouse-greenhouse activity on Earth (Gernon et al., 2022; McKenzie et al., 2016; Pall et al., 2018). 1033 Retrograde metamorphism may in turn consume mobile carbon in fluids to precipitate minerals such as 1034 carbonates and/or graphite (Hu et al., 2021; Piccoli et al., 2016; 2018; 2021; Scambelluri et al., 2016). 1035 Finally, solid carbon in carbonates will be mobilized by processes such as oceanic and atmospheric 1036 acidification (e.g., Doney et al., 2009; Eyssautier-Chuine et al., 2016) and reverse weathering (Isson and 1037 Planavsky, 2018). 1038 Much like other elements in the continental crust, the concentration of carbon is assessed by determining 1039 the volumetric quantities of each lithology in the crust, and then further assigning each lithology with a 1040 carbon concentration characteristic of representative rocks exposed at the Earth’s surface (Derry, 2014; 1041 Hartmann et al., 2012; Rudnick and Gao, 2003, 2014). A significant source of uncertainty in estimating 1042 crustal carbon budgets is therefore the uncertainty of the mass of the crust itself, which must be considered 1043 in addition to carbon concentrations within characteristic lithologies (Derry, 2014). Some studies estimating 1044 the carbon content of the continental crust do not consider carbon on continental shelves (e.g., Wedepohl, 1045 1995), which accounts for as much as 31% of present-day continental area (Cogley, 1984). Further 1046 uncertainty is introduced by sampling bias: the uppermost continental crust is perhaps the most interrogated 1047 owing to its easier accessibility (Derry, 2014; Hartmann et al., 2012). However, lithologies at the Earth’s 1048 surface are prone to chemical weathering, which increases their carbon content over time as a result of 1049 secondary carbonate formation (Hartmann et al., 2012). Estimates of carbon contents in the mid-lower crust 1050 are rarer in comparison (Derry, 2014), and are dependent on predictions of the lithological composition of 1051 the lower crust, which may be supplemented by seismic observations (e.g., Wedepohl, 1995; Yaroshevsky, 1052 2006). Finally, both oxidized and reduced forms of carbon must be considered to provide a complete picture 1053 of continental carbon budgets. In the bulk continental crust, the proportion of reduced carbon relative to the 1054 total carbon budget is considered to range from 0.10–0.28 (Derry, 2014; Gao et al., 1998; Hartmann et al., 1055 2012; Wedepohl, 1995; Yaroshevsky, 2006); however this is highly dependent on sample selection, and 1056 varies substantially between lithology types. For example, sediments typically have a greater proportion of 1057 reduced carbon than crystalline basement (e.g., Yaroshevsky, 2006). 1058 A recent summary of the present-day carbon budget of the continental crust is provided by 1059 Hirschmann (2018). Literature values assessed by Hirschmann (2018) converge on values of 76±16×1021 g 1060 C hosted in sedimentary rocks and 15±8×1021 g C within crystalline rocks for a total of 91×1021 g C, and is 1061 based on previous estimates of crustal carbon budgets (e.g., Gao et al., 1998; Hayes and Waldbauer, 2006; 1062 Hunt, 1972; Wedepohl, 1995; Yaroshevsky, 2006). This budget is distributed across the solid exosphere, 1063 37 including carbon in oceanic domains (see following section), at a concentration of 27±5 ppm (Hirschmann, 1064 2018). 1065 1066 4.4.2. Oceanic crust and upper lithospheric mantle 1067 At the present day, the total mass of carbon in oceanic crust and lithosphere is on the order of ~1.0–1068 1.5×1022 g C (Hirschmann, 2018; Müller et al., 2022). The majority of this carbon (~60%) is stored within 1069 pelagic sediments, with the remainder comprising carbonate minerals in the oceanic basement and sub-1070 oceanic lithospheric mantle (Hirschmann, 2018; Müller et al., 2022). Because seawater-rock interactions in 1071 the oceans do not only affect crustal sequences but also the uppermost part of the oceanic lithospheric 1072 mantle, this section presents an overview of the upper oceanic lithosphere affected by magmatic and 1073 hydrothermal processes at mid-ocean ridges. 1074 1075 4.4.2.1. Altered lithospheric mantle 1076 Seawater percolates through permeable uppermost oceanic crust and lithospheric mantle where it 1077 may be exposed, e.g., at mid-ocean ridges or through faulting (e.g., Fisher, 2005). Carbonate minerals can 1078 form during serpentinisation through the coupled hydration and carbonation of olivine to form serpentine 1079 minerals and carbonates (e.g., Power et al., 2013; Kelemen et al 2011; Menzel et al., 2018). Ultramafic 1080 rocks can therefore have significant potential for rapid carbonate formation upon reaction with CO2-bearing 1081 fluids (e.g., Kelemen et al., 2011; Kelemen and Matter, 2008; Matter and Kelemen, 2009). Significant 1082 natural carbon sequestration is expected where mantle peridotite is exposed to seawater at oceanic settings 1083 and subject to carbonation (Alt et al., 2013). Serpentinisation and coupled carbonation of the oceanic 1084 lithosphere may therefore occur via a number of different tectonic processes: a) widespread exposure and 1085 faulting of peridotite at (ultra-)slow spreading centers (Alt et al., 2013; Cannat et al., 2010; Dick et al., 1086 2003; Kelemen et al., 2011; Merdith et al., 2020; White et al., 2001) and locally at the ridge axis at faster 1087 spreading centers and oceanic transform zones (Alt et al., 2013; Klein et al., 2024); b) exposure of 1088 continental lithospheric mantle within magma-poor continent-ocean transitions (Albers et al., 2021; 1089 Grevemeyer et al., 2022; Liu et al., 2023; Sawyer et al., 1994; Schwarzenbach et al., 2013); c) faulting of 1090 oceanic plates during bending at the outer rise as they enter subduction zones (van Avendonk et al., 2011; 1091 38 Christensen and Ruff, 1988; Faccenda, 2014; Grevemeyer et al., 2018; Lefeldt et al., 2012; Ranero et al., 1092 2003). 1093 There is a higher concentration of carbon near exposed serpentinite seafloor as a consequence of 1094 seawater accessibility. While uppermost oceanic serpentinite is primarily inorganic carbon, deeper 1095 serpentinite (>50–100 m) principally hosts organic carbon, which is sourced from methanogenic 1096 microorganisms within serpentinites that thrive on the hydrogen gas byproduct of serpentinisation 1097 (Schwarzenbach et al., 2013) and abiotic forms of organic matter (e.g., Sforna et al., 2018; Ménez, 2020). 1098 Most data on carbon concentrations in altered peridotite are obtained from seafloor drill samples, with the 1099 remainder determined from ophiolites (Kelemen and Manning, 2015). Carbon in these altered peridotites 1100 have an average bulk concentration of 681±45 ppm C (Kelemen and Manning, 2015), although 1101 concentrations from some settings may be as high as 17.1 wt.% CO2 (Albers et al., 2021). Carbonated 1102 serpentinites therefore host a minor yet not insignificant proportion of the overall oceanic lithosphere 1103 carbon budget, estimated as ~0.2–0.3×1021 g C (Merdith et al., 2020; Müller et al., 2022). 1104 1105 4.4.2.2. Altered oceanic crust 1106 In the high-permeability upper crust, carbon, sourced from magmatic fluids and seawater, 1107 precipitates as the carbonate minerals calcite and aragonite within rocks and veins at low temperatures 1108 (<100 °C; Alt and Teagle, 1999). Most basaltic crust alteration will occur within the first few tens of 1109 millions of years after crust formation (Alt and Teagle, 1999; Coogan and Dosso, 2015; Coogan and Gillis, 1110 2018; Coogan et al., 2016). The combination of pore infilling through carbonate precipitation and gradual 1111 sedimentation will limit subsequent circulation of water through oceanic crust over geological time, thereby 1112 decreasing carbon sequestration as the crust ages (Grevemeyer et al., 1999); overlying carbonate sediments 1113 may themselves be incorporated into the basaltic crust as recrystallized carbonate minerals to facilitate this 1114 process (Gillis and Coogan, 2011). High degrees of basalt alteration will also limit the availability of 1115 exposed fresh material required to precipitate carbonate, thereby limiting further carbonate precipitation in 1116 aged crust (Albers et al., 2023). Another control on crustal carbonate formation is bottom-water 1117 temperature, which promotes faster fluid-rock reactions resulting in carbonate precipitation (Coogan and 1118 Gillis, 2013; Gillis and Coogan, 2011). 1119 Estimates of bulk oceanic crustal CO2 concentration suggest a range of 0.18–0.22 wt.% CO2 (Alt 1120 and Teagle, 1999; Kelemen and Manning, 2015). Most CO2 is concentrated in the uppermost 300 m of 1121 39 oceanic crust at an estimated mean concentration of 1.85 wt.% (Müller and Dutkiewicz, 2018), which 1122 represents two-thirds of the total CO2 in typically ~7 km-thick oceanic crust (Gillis and Coogan, 2011; 1123 Staudigel et al., 1989). Estimates of present-day oceanic crust carbon storage using these concentration 1124 estimates range from 2–4×1021 g C (Hirschmann, 2018; Kelemen and Manning, 2015; Müller and 1125 Dutkiewicz, 2018). At least a fraction of this carbon is present in the form of organic compounds 1126 (Shilobreeva et al., 2011). 1127 1128 4.4.2.3. Sediments 1129 The key parameters to consider when assessing the carbon budget of pelagic oceanic sediments are 1130 the mass of present-day sediment, and the amount of carbon present within said sediment. Both parameters 1131 are highly heterogeneous in present-day oceans (Conrad, 2013; Dutkiewicz et al., 2016; Hayes et al., 2021), 1132 and are assessed respectively through seafloor seismic reflection surveys and ocean drilling programs 1133 (Plank, 2014; Plank and Langmuir, 1998; Rea and Ruff, 1996). The principal controls on the thickness of 1134 sediment at any given point on the seafloor are a) the distance of the point from the nearest passive 1135 continental margin, a measure of the sedimentary input to that point, and b) the age of the oceanic crust, a 1136 measure of time over which sedimentation has occurred (Dutkiewicz et al., 2017; Olson et al., 2016). Large 1137 rivers will accelerate local oceanic sedimentation rates; for example, the Ganges river fan possesses a 1138 column of sediment of 7 km thickness at its maximum (Curray, 2014). 1139 Carbon-bearing sediments can be subdivided into carbonate (predominantly CaCO3) and organic 1140 carbon, which represent the oxidized and reduced forms of carbon respectively (e.g., Clift, 2017). 1141 Accumulation of significant thicknesses of oceanic sedimentary carbonate is a geological novelty on Earth 1142 (e.g., Gilbert et al., 2022). While metazoan carbonate biomineralization is known from the Ediacaran 1143 onwards (e.g., Grant, 1990; Wood, 2011), the onset of extensive neritic biomineralization during the Marine 1144 Mesozoic Resolution was the principal herald of widespread and significant marine carbonate production 1145 representative of present-day carbonate sedimentation in the oceans (Erba, 2006; Falkowski et al., 2004; 1146 Ridgwell, 2005). 1147 Sedimentary carbon accumulation on oceanic crust is limited by the carbonate/calcite compensation 1148 depth (CCD), which marks the depth below which the rate of inorganic carbonate dissolution exceeds the 1149 rate of deposition (e.g., Broecker, 2008; Derry, 2022; Dutkiewicz et al., 2019; Ridgwell and Zeebe, 2005; 1150 van Andel, 1975). Within present-day oceans, the CCD corresponds to 4.5–5.0 km below sea level 1151 40 (Dutkiewicz et al., 2016; Zeebe, 2012). Terrestrial sediment fluxes (Derry, 2022; Walling and Fang, 2003), 1152 biological productivity (Kumar et al., 1995; Sarmiento et al., 2004), and sea-surface temperature and 1153 salinity (Dutkiewicz et al., 2016), all act in concert to affect carbonate deposition or dissolution, thereby 1154 controlling the position of the CCD. The parameterisation of these factors permits the extrapolation of 1155 carbonate sediment thicknesses to all present-day ocean basins (Dutkiewicz et al., 2019). Furthermore, the 1156 thermal subsidence of ocean basins with age will mean that old oceanic crust will carry negligible 1157 sedimentary carbonate, in contrast to shallow seafloor in regions of high biological productivity 1158 (Dutkiewicz et al., 2016, 2019; Plank, 2014). On the other hand, organic carbon is rapidly consumed by 1159 biological processes such as respiration (e.g., Higgins et al., 2009). Preservation of organic carbon in 1160 sediments therefore requires rapid deposition and burial, for example, beneath regions of high biological 1161 productivity or at deep sea fans where sedimentation rates are higher (Galy et al., 2007; Plank and Manning, 1162 2019). Temperature during sediment burial additionally has a strong effect on organic carbon sediment 1163 preservation (Maliverno and Martinez, 2015). Finally, the weathering of silicates and organic carbon in 1164 pelagic sediments will fix substantial CO2 into carbonates over geological time (Wallmann et al., 2008). At 1165 the present day, the mass of carbon within sediments is estimated to be on the order of 10×1021 g C 1166 (Dutkiewicz et al., 2019; Hirschmann, 2018; Yaroshevsky, 2006); pelagic sediments are therefore the 1167 largest reservoir of carbon on oceanic plates at the present day. 1168 1169 4.5. Deep abiotic organics 1170 Carbon dioxide is generally considered as the dominant carbon species in the modern Earth (Figure 1171 1). However, as conditions become more reducing at depth, CH4 and H2, together with smaller 1172 hydrocarbons, become stable at the expense of CO2 or other more oxidized carbon species within the 1173 stability field of graphite or diamond (Sverjensky et al., 2020). In the mantle, the transition from CO2-1174 dominated to CH4-dominated regimes occurs at depths of 125–140 km (~4 GPa) in sub-continental 1175 domains, and at shallower depths of ~90 km (~2.5 GPa) in the sub-oceanic mantle. The chemical complexity 1176 of the fluid organics that could be present is exemplified by the wealth of organics found in the fluid 1177 inclusions of diamonds from the Urals placers from ~200 km depth, though their exact origin remains 1178 unknown (see Table 4 in Sobolev et al., 2019). They contain aliphatic hydrocarbons up to C17, cyclic 1179 hydrocarbons up to C15, oxygenated hydrocarbons, and heterocyclic compounds, together with 1180 nitrogenated and sulfonated compounds (plus CO2 and H2O). This is supported by experimental and 1181 theoretical work. Above 2 GPa, heating pure CH4 at high temperature triggers its polymerization to higher 1182 alkanes (C3), with ethane the most abundant species (Kolesnikov et al., 2009). Theoretical simulations have 1183 41 shown that formic acid is thermodynamically more stable than the products of the water-gas shift reaction 1184 (CO2 and H2) above 3 GPa and at ~750–1150 °C (Stolte et al., 2021). This result is in good agreement with 1185 the spectroscopic signature of carboxylic functional groups coating diamond inclusions in the deeply 1186 subducted metamorphic rocks of the Lago di Cignana in the Western Alps at P ≥ 3.2 GPa and T ~600 °C 1187 (Frezzotti, 2019), and with thermodynamic models suggesting that organic acids play an important role in 1188 diamond nucleation and growth in deep Earth (Sverjensky et al., 2014; Sverjensky and Huang, 2015). 1189 Similarly, but at shallower depths, the graphite or other condensed carbon forms present in many crustal 1190 and mantle rocks such as those described at mid-ocean ridges of within subduction contexts — and now 1191 emplaced and incorporated in the continental crust — may not be of biotic origin (i.e., transformation of 1192 subducted biogenic organic matter), but instead has precipitated from deep carbonic fluids (e.g., Boutier et 1193 al., 2024b; Debret et al., 2022; Evans et al., 2002; Luque et al., 1998; Ménez et al., 2018) or from the 1194 interaction between reduced fluids and carbonates (Galvez et al., 2013; Malvoisin et al., 2012; Peng et al., 1195 2021; Tao et al., 2018; Vitale Brovarone et al., 2017). 1196 Expanding on the carboxylic acids, experimental work has shown that pressure expands miscibility 1197 gaps in C-O-H fluids, and immiscible fluids are generated at pressures in excess of 1.5 GPa (600–700 °C), 1198 with formic acid yielding both aqueous and dry methane and ethane fluids ±CO2 and ±H2 depending on the 1199 oxygen fugacity (Li, 2017). At higher pressures (up to 4.6 GPa) and lower temperatures (300–350 °C), 1200 aqueous Naor Ca-acetate yields aqueous and mineral carbonate (as expected from the speciation of C(IV) 1201 at high pressure; Facq et al., 2014), soluble light alkanes (C1–C3), and immiscible hydrocarbons; the higher 1202 the pressure, the greater the amount of hydrocarbons and methane generated (see F. Huang et al., 2017; J. 1203 Huang et al., 2023 for details). At 350 °C and above, immiscible hydrocarbons tend to become aromatic, 1204 and may transform into polycyclic aromatic hydrocarbons (PAH) and ultimately into graphite, depending 1205 on temperature and composition. Natural fluid inclusions from high-pressure metamorphic rocks support 1206 the immiscibility of CH4, and potentially the immiscibility of other hydrocarbons in aqueous fluids at high 1207 pressure conditions (Giuntoli et al., 2024; Sverjensky et al., 2020; Vitale Brovarone et al., 2017). 1208 As of today, the complexity of the fluid organics observed in deep fluid inclusions has been partly 1209 reproduced in experiments (McCollom, 2013), emphasizing that highly mobile and likely highly reactive 1210 hydrocarbons are stable at sub-lithospheric conditions, provided that the temperature is low enough to 1211 kinetically prevent the formation of graphite or diamond, as their compositions are mainly controlled by 1212 the local water activity and oxygen fugacity. 1213 Whilst lithospheric subsurface environments should be the exclusive realm of CO2 as discussed in 1214 previous sections, there are numerous occurrences of CH4 reported in vents, seeps, springs and aquifers at 1215 48 crustal contaminants (e.g., Straub et al., 2020), it is possible to identify signatures of eroded crust in arc 1396 magmas (e.g., Stern, 2020; Straub et al., 2015). Estimates of present-day subduction erosion typically range 1397 from 30–150 km3 Myr-1 per km length of subduction zone (e.g., Clift, 2017; Straub et al., 2020; Vannucchi 1398 et al., 2016); it is therefore likely that erosion of the overriding plate can contribute towards the carbon 1399 budget of subduction. However, it is difficult to determine the contributions of this contribution beyond 1400 qualitative assessment. It is known that the rate of material transfer during erosion depends on several 1401 factors, including the topography and sediment cover of the subducting plate, high convergence rates, and 1402 subduction channel fluid dynamics (Clift and Vannucchi, 2004; Straub et al., 2020; Vannucchi et al., 2012). 1403 Furthermore, eroded convergent margins typically have overlying plates exposing igneous arc basement, 1404 which bears less carbon than other lithologies (Clift, 2017; Yaroshevsky, 2006); the lithology of the 1405 overriding plate will be a further contributor to the overall carbon budget of subduction erosion. Further 1406 characterisation of the subduction erosion process, and therefore also of volume and mass fluxes of eroded 1407 overlying crust at subduction zones, will be required in order to completely capture the quantities of carbon 1408 contributed to subduction from this process. 1409 1410 5.2. Movements of carbon from depth to the crust 1411 5.2.1. Tectonic movements of solid carbon from the deep Earth to the surface 1412 At convergent margins, tectonic or mechanical movements of solid carbon from mantle depths to 1413 the crust or within the crust are primarily controlled by off-scraping or diapirism of downgoing slab material 1414 or eroded mantle wedge material (see reviews by Klein and Behn, 2021; Plank and Manning, 2019). 1415 Carbon-bearing subducted materials, especially subducted sediments but also large volumes of subducted 1416 oceanic igneous and mantle rocks, can be removed from the downgoing slab, reducing the amount of 1417 subducted carbon reaching the convective mantle. A dominant fraction of this material is removed from the 1418 downgoing slab at trenches to form accretionary prisms (Tewksbury-Christle et al., 2021), At greater depths 1419 within the forearc to arc region, the fate of this material is uncertain; subducted sediments may be removed 1420 from the slab to form diapirs that partially melt or do not melt as they rise into the mantle wedge, and be 1421 emplaced at the base of the crust (Behn et al., 2011; Ducea et al., 2021; Kelemen and Behn, 2016), or be 1422 exhumed as tectonic slices or mélanges along the plate boundary to form orogenic complexes (Agard et al., 1423 2009; Guillot et al., 2009). The fluxes of solid carbon associated with these tectonic or mechanical 1424 movements of deep rocks towards the surface are largely unknown. It has been suggested that off-scraping 1425 of sediments at trenches may remove a fraction of subducted sediments as high as 35% of global subducted 1426 49 sediments (Straub et al., 2020). Nevertheless, the ultimate amount of solid carbon being mobilized through 1427 tectonic or mechanical processes varies from one subduction zone to another, or even subduction zone sub-1428 segments relative to adjacent sub-segments, based on the specific tectonic setting, the type and amounts of 1429 subducted sediments, and the thermal regimes controlling the contextual mobilization of carbon through 1430 metamorphic or magmatic degassing (Plank and Manning, 2019). 1431 1432 5.2.2. Metamorphic and magmatic processes of carbon mobilization 1433 The mobilization of carbon from solids to fluids requires changing conditions and is strongly 1434 dependent upon the closed or open behavior of the considered systems. Changes in pressure and/or 1435 temperature conditions, changing redox, and/or on the presence and nature of fluids are the main parameters 1436 controlling the mobilization of carbon from the main solid carbon phases, namely carbonate minerals, 1437 organic matter, diamond (Figure 5), and fluids and melts. Such conditions can be achieved through multiple 1438 processes and at various geological contexts. These processes are summarized in the following sections. 1439 1440 5.2.2.1. Mobilization of carbon from carbonate minerals 1441 The three main processes mobilizing carbon from carbonate minerals are decarbonation reactions, 1442 carbonate dissolution, and melting. In the forearc to subarc regions of subduction zones as well as in the 1443 continental crust, decarbonation and dissolution are the most important processes, whereas carbonate 1444 melting is expected to take place at greater depths in the mantle, or during anatexis in the lower crust. 1445 Decarbonation is the process of carbonate breakdown through carbonate-silicate reactions in 1446 response to raising temperature. Also referred to as ‘Urey reactions’ (Urey, 1952), decarbonation can be 1447 summarized by the generic reaction: 1448 MCO3 + SiO2 → MSiO3 + CO2 1449 where M represents a divalent cation, most typically Ca2+ or Mg2+. 1450 Examples of decarbonation reactions are common in metamorphic rocks in metasedimentary, 1451 metamafic, and ultramafic rocks, and involve a broad range of carbonate (calcite, dolomite, ankerite) and 1452 silicate (quartz, serpentine, and many others) reactants (Fig. 6a). Besides the characteristic calc-silicate 1453 products, decarbonation reactions are also identified by isotopic shifts in the residual carbonate fraction. 1454 50 Residual carbonate is expected to track δ13C depletion owing to the positive ΔCarb-CO2; e.g., Baumgartner 1455 and Valley, 2001; Bottinga, 1969). Decarbonation reactions do not require the presence of an aqueous fluid 1456 to produce carbon-bearing fluids. Nevertheless, the presence of aqueous fluids, for example during 1457 infiltration of external fluids or in response to local dehydration reactions, has been shown to enhance 1458 decarbonation reactions. Decarbonation is believed to contribute substantially to metamorphic carbon 1459 degassing at convergent margins (Gorman et al., 2006; Kelemen and Manning, 2015; Kerrick and Connolly 1460 2001), with collisional metamorphism producing higher CO2 fluxes relative to subduction metamorphism, 1461 and cumulative fluxes as high as about 6–84 Mt C yr-1 (Stewart et al. 2019; Figure 4). Experimental studies 1462 and thermodynamic modeling results predict decarbonation reactions and the formation of CO2-enriched 1463 fluids to be particularly efficient at subarc depths (Gorman et al., 2006; Gonzalez et al., 2016). 1464 In addition to decarbonation reactions, carbonate dissolution also contributes to the mobilization of 1465 deep carbon from solid to fluid phases. Carbonate dissolution has only recently been added to the inventory 1466 of processes contributing to deep carbon budgets and fluxes (Ague and Nicolescu, 2014; Facq et al., 2014; 1467 Frezzotti et al., 2011; Kelemen and Manning, 2015). The lack of previous assessments was related to the 1468 only recent establishment of values for the dielectric constant of water at high-pressure conditions 1469 (Svejensky et al., 2014). Carbonate dissolution may be effective at mobilizing carbon into aqueous fluids 1470 in the absence of silicate minerals, i.e., when decarbonation reactions cannot take place. Carbonate 1471 solubility is dependent upon pressure, temperature, and redox conditions (Fig. 6b,c). Experimental studies 1472 on aragonite solubility demonstrated that, under oxidized conditions, the speciation of carbonic fluids 1473 produced by carbonate dissolution is strongly dependent upon the pressure and temperature conditions, 1474 with bicarbonate ions dominating at pressure <3 GPa to at least 400 °C and carbonate ions at higher pressure 1475 and temperature conditions (Facq et al., 2014). CO2 was found to be present in appreciable amounts below 1476 2–3 GPa (Li, 2017). Reduced conditions, instead, have been shown to boost carbonate solubility in aqueous 1477 fluids, with methane controlling the speciation of carbonic aqueous fluids resulting from carbonate 1478 dissolution (Lazar et al., 2014) (Fig. 6c). Natural examples of carbonate dissolution are documented in 1479 metamorphic terranes such as the Cyclades (Ague and Nicolescu, 2014) and Corsica (Piccoli et al., 2016; 1480 2021). 1481 An increasing body of literature has documented examples of natural carbonate reduction at various 1482 metamorphic settings (Boutier et al., 2024b; Galvez et al., 2013; Tao et al., 2018; Peng et al., 2021Vitale 1483 Brovarone et al., 2017; Wang et al., 2022) (Fig. 5b). Although at least some of these cases can be explained 1484 by carbonate dissolution in aqueous fluids under reducing conditions, the immiscibility of H2 in 1485 metamorphic fluids under a vast pressure-temperature range may promote carbonate-H2 reactions in the 1486 51 absence of an initial aqueous fluid. In this instance, the term carbonate hydrogenation should be preferred. 1487 A simple example of carbonate reduction is 1488 CaCO3 + 4H2 → Ca(OH)2 + CH4 + H2O 1489 where the breakdown of carbonate is promoted by hydrogenation. Carbonate reduction and hydrogenation 1490 are documented in subduction zone metamorphic rocks as a consequence of fluid-rock interactions such as 1491 serpentinization (Boutier et al., 2024b; Peng et al., 2021; Vitale Brovarone et al., 2020a), or during prograde 1492 or early retrograde metamorphism of weakly altered oceanic crustal rocks (Wang et al., 2022). The same 1493 processes are also expected to occur at mantle pressures and temperatures (Scott et al., 2004). Fluxes of 1494 carbon through these processes are largely unquantified. Fluxes of CH4 produced through metamorphic 1495 decarbonation reactions of ultramafic and mafic rocks are estimated at 1.5 Mt CH4 yr-1 (Vitale Brovarone 1496 et al., 2017) and up to about 11 Mt CH4 yr-1 (Zhang et al., 2023). 1497 1498 5.2.2.2. Mobilization of carbon from organic matter, graphite, and diamond 1499 During diagenesis and low-temperature (<300 °C) metamorphism, organic matter 1500 undergoes a series of chemical reactions that progressively release heteroatoms such as hydrogen, sulfur, 1501 nitrogen, and oxygen, in addition to carbon. Following laboratory experimental results and observations in 1502 natural settings, this process results in the carbonization and – at higher temperature – graphitization of 1503 organic matter (Vendenbroucke and Largeau, 2007) (Figure 7a). With progressive loss of heteroatoms 1504 during carbonization (or kerogen transformation) and graphitization, carbon is released in the fluid in the 1505 form of hydrocarbon oils and gasses (Tissot and Welte, 1984; Vendenbroucke and Largeau 2007), 1506 promoting the formation of a progressively purer and more structured organic matter, or graphitic carbon 1507 (Fig. 5c) (Beyssac and Rumble, 2014). Light hydrocarbon gasses produced during this process are referred 1508 to as thermogenic gasses and, as well as for oils, are referred to as biotic in origin (Fig. 5d) (i.e., formed 1509 through the transformation of biogenic matter; Etiope, 2015). Studies conducted in low-grade metamorphic 1510 terranes suggest that thermogenic degassing is efficient up to about 300 °C, whereas at higher temperatures 1511 the contribution of this process is considered negligible (Mullis et al., 1994; Tarantola et al., 2007). 1512 Thermogenic degassing is primarily controlled by temperature and does not require an external agent such 1513 as aqueous fluids. Such conditions are expected during diagenesis in sedimentary basins, or during prograde 1514 metamorphism in the shallow forearc of subduction zones margins, in accretionary prisms, and in contact 1515 metamorphic aureoles hosted in organic-rich sedimentary rocks. The contribution of thermogenic degassing 1516 of hydrocarbons induced by igneous intrusions may release carbonic fluids in amounts high enough to affect 1517 52 global climate conditions. For example, it has been proposed that the end-Triassic mass extinction and 1518 initial Eocene global warming may have been caused by voluminous intrusions of mantle-derived melts in 1519 organic-rich sedimentary basins at large igneous provinces (Capriolo et al., 2021; Svensen et al., 2004). 1520 Beyond the conditions favorable to extensive thermogenic degassing, the contribution of reduced 1521 solid carbon materials to geological carbon degassing may be minimal in closed systems (Connolly, 1995; 1522 Pattison, 2007), even though high-temperature thermogenic degassing may still be present (Boutier et al., 1523 2024b). At those conditions, graphitic carbon solubility becomes the dominant mechanism of carbon 1524 mobilization from organic carbon solids to fluids, especially in open systems flushed by metamorphic 1525 aqueous fluids (Fig. 5e) (Tumiati et al., 2020, 2022; Vitale Brovarone et al., 2020b; S. Zhang et al., 2018). 1526 The carbonic gasses produced by aqueous dissolution of graphitic carbon should be considered abiotic in 1527 origin (Figure 7), even though – in the absence of carbonate minerals, see below – their carbon isotope 1528 composition may reflect a biotic origin. 1529 The solubility of graphitic carbon in aqueous metamorphic fluids depends on the redox state, on 1530 the pressure and temperature conditions, and on their physicochemical properties (e.g., crystallinity). The 1531 amount of carbon that can be dissolved from organic materials to aqueous fluids is higher for strongly 1532 oxidized or strongly reduced conditions, and is minimized for intermediate redox conditions at the so-called 1533 “water maximum” of carbon-saturated C-O-H fluids (Figure 7b) (Connolly and Cesare, 1993; Holloway, 1534 1984). Pressure decreases the solubility of graphitic carbon, whereas temperature has an opposite effect 1535 (Figure 7c). Crystalline organic matter is less soluble than amorphous organic matter (Figure 7b) (Tumiati 1536 et al., 2020, and references therein; Vitale Brovarone et al., 2020b). Increasing temperature during prograde 1537 metamorphism promotes the progressive crystallization of organic matter, increasing its solubility. In 1538 subduction zones, increasing pressure during prograde metamorphism may, depending upon thermal 1539 regimes, counterbalance this effect (Figure 7c). 1540 Experimental work has demonstrated that, for graphite-aragonite-water equilibria at fH2 = FMQ 1541 (equivalent to fO2 expressed as ΔFMQ = +0.61 log units) at subarc conditions, the largest fraction of 1542 dissolved CO2 derives from graphite solubility, whereas carbonate dissolution at the investigated conditions 1543 produces bicarbonate ions mostly (Tumiati et al., 2022). In turn, the isotope composition of the 1544 experimentally produced, graphite-derived CO2 reflected carbonate-CO2 equilibria, with implications on 1545 the interpretation of volcanic gas compositions and their slab source materials. 1546 As discussed for carbonate minerals, interactions between organic matter and H2-rich fluids may 1547 promote the mobilization of carbon from solid phases to light hydrocarbons in fluids. Experimental results 1548 obtained at pressure-temperature conditions consistent with subduction zone and upper mantle settings 1549 53 indicate that graphite hydrogenation to form CH4 may be a very efficient process (Peña-Alvarez et al., 1550 2021). Experiments of diamond hydrogenation showed similarly fast, yet slightly lower reaction rates 1551 compared to graphite (Peña-Alvarez et al., 2021). Natural examples of graphite hydrogenations are also 1552 documented (Vitale Brovarone et al., 2017). 1553 1554 5.3. Carbon reprecipitation at depth 1555 Not all the carbon being mobilized from metamorphic and metasomatic processes is degassed at the 1556 Earth’s surface. Precipitation of carbonates, graphite, organic compounds, or diamonds from geological 1557 fluids or melts is a common process taking place at virtually any depth geological setting. Carbon saturation, 1558 redox, pressure/temperature variations and fluid-rock chemistry are some of the key parameters controlling 1559 the precipitation of carbon minerals from fluids or melts. At least some diamonds are thought to represent 1560 precipitates from deep/ultradeep carbonic fluids (Smith et al., 2016; Stachel et al., 2017; Thomassot et al., 1561 2007). Voluminous or trace amounts of graphite form from percolating fluids and their interactions with 1562 rocks and/or other fluids and in rocks and geological contexts spanning mafic and ultramafic rocks at 1563 convergent and divergent margins (Luque et al., 2014), and virtually any rock types at various metamorphic, 1564 magmatic, and hydrothermal settings within the crust (Fig. 5f,g) (Boutier et al., 2024b; Galvez et al., 2013; 1565 Luque et al., 2014; Rumble et al., 1986; Tao et al., 2018). This also extends to other forms of condensed 1566 carbon materials, precipitation of which may serve as a carbon source for microbial pathways in the deep 1567 subsurface (Debret et al., 2022; Ménez et al., 2018). Carbonate precipitation in upper mantle rocks has been 1568 documented in multiple settings. Carbonate veins and carbonation processes, i.e., pervasive replacement of 1569 silicate rocks by carbonate minerals, recycling deep carbon in fluids are ubiquitous in crustal and mantle 1570 rocks, as demonstrated by natural and experimental data (Fig. 5h,i) (Consuma et al., 2020; Bouilhol et al., 1571 2022; Hu et al., 2021; Kelemen et al., 2011; Menzel et al., 2018; Peng et al., 2020; Piccoli et al., 2016; 1572 Sieber et al., 2018). The so-called cold nose of the forearc mantle wedge of many subduction zones 1573 represents a characteristic zone of intense reprecipitation of dissolved carbon through carbonation processes 1574 (e.g., De Obeso et al., 2022; Kelemen et al., 2018, 2022; Menzel et al., 2018, 2024). In the forearc region 1575 of Costa Rica, carbon and noble gas data from hot springs indicate that a large fraction — as high as 91% 1576 — of carbon in fluids released from devolatilization reactions in the subducting slab reprecipitate as 1577 carbonate minerals in the crust before reaching the surface (Barry et al., 2019). Within the biosphere, 1578 microbial activity may also contribute to the modulation of carbon degassing through the precipitation of 1579 carbon minerals or organic biomass (see Section 3). 1580 54 1581 5.4. Movements of deep carbon in melts 1582 5.4.1. Carbon in intraplate magmatism 1583 Partial melting in the asthenospheric (convecting) mantle is an inescapable consequence of the 1584 presence of CO2 and H2O (Figure 8). Nominally anhydrous minerals can store small amounts of H2O, which 1585 results in a decrease of their melting temperature from 10 °C to ~100 °C; the more H2O, the lower the 1586 melting temperature (Bell and Rossman, 1992; Bolfan-Casanova, 2005; Demouchy and Bolfan-Casanova, 1587 2016; Green, 1973; Hirschmann et al., 2009). In contrast, carbon and CO2 are insoluble in mantle minerals 1588 at asthenospheric conditions, but can be stored as carbonates in relative oxidizing conditions. Accordingly, 1589 various constraints from experimental petrology have long shown that, independently of its concentration 1590 in the system, carbon stored as carbonates decreases the solidus of mantle rocks by several hundred degrees 1591 (Dasgupta, 2018; Green, 1973; Wallace and Green, 1988; Falloon and Green, 1989; Hammouda and 1592 Keshav, 2015). 1593 Geochemical surveys have identified intermingled depleted and enriched mantle regions with H2O 1594 and CO2 contents ranging from a few ppm to a few wt.% (tentative average: 50–250 ppm, Hirschmann, 1595 2010; Marty, 2012; Le Voyer et al., 2017; Shimizu et al., 2019). This is thus expected to produce minute 1596 amounts of partial melts in the asthenosphere, of the order of 0.1 vol.% to several vol.%. Possible variations 1597 in pressure and temperature in the asthenospheric mantle are not expected to prevent the presence of 1598 incipient melts. Oxidation state could however preclude incipient melting if carbon is present as diamonds 1599 rather than oxidized carbon (i.e., carbonates). Oxygen fugacity conditions in the convective upper mantle 1600 are however broadly oxidizing enough to enable incipient melting from 50 km down to 200–300 km depth 1601 (Brey et al., 1983; Gaillard et al., 2015; Kushiro, 1975; Moussallam et al., 2019; Rohrbach and Schmidt., 1602 2011; Stagno et al., 2013; Wyllie et al., 1983; White and Wyllie, 1992; Zhang et al., 2024). At greater 1603 depths, diamonds are expected to prevail, preventing the occurrence of melting (Rohrbach and Schmidt, 1604 2011; Stagno et al., 2013). 1605 Pressure and temperature variations in the asthenosphere are, however, expected to significantly 1606 affect the chemical nature and volume fraction of melts produced by CO2-H2O-bearing peridotite melting. 1607 High-pressure and low-temperature melting tends to produce low-SiO2 carbon-rich melts (i.e., 1608 carbonatites), whereas low-pressure and high-temperature melting tends to produce silicate melts (i.e., 1609 basalts). In between, all intermediate compositions are possible (Brey et al., 2008; Dasgupta et al., 2007, 1610 55 2013b; Gudfinnsson and Presnall, 2005; Massuyeau et al., 2015, 2021). Figure 9 illustrates this variability 1611 in melt composition versus depth for a ‘normal’ mantle (140 ppm CO2; 240 ppm H2O; mantle potential 1612 temperature = 1350 °C; see Massuyeau et al., 2021). The melt compositions transition relatively abruptly 1613 from carbonatites to basalts in response to small variations in pressure and temperature (Dasgupta et al. 1614 2007, 2013b; Stagno and Frost, 2010). This reflects the strongly non-ideal mixing properties between 1615 molten carbonates and molten silicates, which often results in liquid-liquid immiscibility (Brooker and 1616 Kjarsgaard, 2011; Freestone and Hamilton, 1980; Massuyeau et al., 2015; Nabyl et al., 2020; Novella et al., 1617 2014). Immiscibility between silicate and carbonate melts is a popular mechanism to explain the formation 1618 of many carbonatites found at the surface (Nabyl et al., 2020; Yaxley et al., 2022). It has often been 1619 described in experiments conducted at moderate pressure and in variably fractionated, low-temperature 1620 alkaline melts (e.g., Nabyl et al., 2020; Weidendorfer et al., 2017). Yet, at asthenospheric pressure-1621 temperature conditions, such immiscibility has only been observed during experimental simulations of 1622 partial melting of eclogite or pelitic rocks (Hammouda, 2003; Kiseeva et al., 2012; Litasov and Ohtani, 1623 2010; Thomsen and Schmidt, 2008; Yaxley and Brey, 2004), whereas peridotite melting does not seem to 1624 produce immiscible carbonate-silicate melts. Indeed, increasing pressure and water content tends to 1625 decrease the non-ideal carbonate-silicate mixing, enabling the formation of intermediate melt compositions, 1626 such as the famous kimberlites, which are in equilibrium with mantle peridotites at depth of 120–250 km 1627 (i.e., pressure of 4–8 GPa; Massuyeau et al., 2021; Stamm and Schmidt, 2017). 1628 The chemical composition of these intraplate magmas are diverse, and range from low to high SiO2 1629 contents, which seemingly reflects the interruption of the adiabatic melting regime at variable depths as 1630 illustrated in Figure 9 (Gudfinnsson and Presnall, 2005; Massuyeau et al., 2021; Tappe et al., 2007). This 1631 interruption at variable depth is controlled by the thickness of the (non-convective) lithosphere, which 1632 ranges from >200 km beneath cratons, where kimberlites are found, to ~60 km in some oceanic domains, 1633 where basanites or alkali basalts are extracted from the convective asthenosphere. Therefore, the 1634 combination of the CO2-driven mantle melting regime and the variable thickness of the lithospheric lids 1635 provides a simple explanation of the variability of melt compositions found in intraplate domains. 1636 These melts are produced by adiabatic incipient melting of the asthenosphere and constitute ~0.1 1637 to several vol.% of the mantle, depending on the bulk concentrations of CO2-H2O and the potential 1638 temperature. On top of these thermochemical variables, melt accumulation and mantle compaction 1639 processes can affect melt behavior in the mantle (Grégoire et al., 2006; Havlin et al., 2013; Keller et al., 1640 2017; Massuyeau et al., 2021; Soltanmohammadi et al., 2018). These accumulation/migration processes 1641 are driven by the peculiar CO2 mantle melting regimes, which produce a range of melt densities (Agee, 1642 2008; Dobson et al., 1996; Jing and Karato, 2011; Massuyeau et al., 2023; Lange and Carmichael, 1987; 1643 56 Ritter et al., 2020; Sakamaki et al., 2006; Suzuki et al., 1995), viscosities (Behrens and Schulze, 2003; 1644 Dingwell, 2007; Dingwell et al., 2022; Dobson et al., 1996; Giordano et al., 2008; Kushiro, 1976; Ritter et 1645 al., 2021; Sakamaki et al., 2013; Whittington et al., 2009), and mantle permeabilities (Grégoire et al. 2006; 1646 Keller et al., 2017), inducing variable melt buoyancies along the melting column and contributing to volatile 1647 redistribution (Bekaert et al., 2021; Dasgupta, 2013; Gaillard et al., 2019; Schettino and Poli, 2020). In 1648 many cases however, calculated melt mobilities are lower than 1 cm yr-1, which implies that they are less 1649 mobile than the convective mantle itself. It is only in regions where melt fractions are higher than the 1650 ‘normal’ mantle (due to higher CO2-H2O or higher potential temperatures) that melt mobility could exceed 1651 mantle convection rates (Gaillard et al., 2019) and thus lead to migration/accumulation/extraction 1652 processes. In most cases, however, one expects that such incipient melts do not reach the surface and remain 1653 trapped at the lithosphere-asthenosphere boundary, causing metasomatism of the lithosphere (see Section 1654 4.3.3). 1655 We do not have in situ analyses of these CO2-driven melting processes, but electrical and seismic 1656 geophysical observations have been used to interpret in near real-time the formation and interconnection of 1657 these incipient melts in the asthenosphere (Chantel et al., 2016; Gardés et al., 2020; Hammouda and 1658 Laporte, 2000; Kawakatsu et al., 2009; Kawakatsu and Utada, 2017; Laumonier et al., 2017; Minarik and 1659 Watson, 1995; Massuyeau et al., 2021; Rychert et al., 2020; Schmerr, 2012; Selway and O’Donnell, 2019; 1660 Selway et al., 2019; Sifré et al., 2014; Yoshino et al., 2010, 2018). The interpretation of remote sensing 1661 geophysical data is not unequivocal as several geological processes could explain one geophysical proxy. 1662 The recently discovered ‘petit-spots’, a new class of volcanoes, might represent some leakage of these 1663 asthenospheric melts, permitted by the bending and stretching of the oceanic lithosphere prior to enter into 1664 subduction (Hirano et al., 2006 ; Hirano and Machida, 2022). These melts are very CO2-rich and may 1665 indicate incipient melts trapped at the lithosphere-asthenosphere boundary. The seminal question is then 1666 how broadly distributed those melts are. This requires the analysis of constraints from experimental 1667 petrology, mantle geochemistry, geophysics and an understanding of the occurrence of petit-spots. 1668 Hammouda et al. (2021) suggested a global layer of such incipient melts at the lithosphere-asthenosphere 1669 boundary, but mantle heterogeneities with carbon-depleted and carbon-enriched regions might speak for 1670 local melt enrichments. 1671 1672 57 5.4.2. Carbon in arc basalts 1673 At subduction zones, arc volcanism transfers some CO2 from the slab to the Earth’s surface. 1674 Hydrous silicate melts, which are formed in the mantle wedge, can transport carbon as dissolved carbonates 1675 in basaltic melts or in supercritical fluids. Due to large uncertainties and large observed variations in CO2 1676 fluxes at arcs (Aiuppa et al., 2019; Mason et al., 2017), few estimates of the average CO2 content of primary 1677 arc magmas are available. CO2 outgassing rates at volcanic arcs are indeed highly variable, with the greatest 1678 CO2 emitters involving an overwhelming contribution from thermally destabilized sedimentary carbonates 1679 (Aiuppa et al., 2019; Iacono-Marziano et al., 2009; Mason et al., 2017). In a review paper, Plank and 1680 Manning (2019) suggested that up to 1 wt.% CO2 could be dissolved and transported in primary arc basalts. 1681 This is c. 10 times greater than the average CO2 contents of primary mid-ocean-ridge basalts. This estimate 1682 is indirect as it is based on C/S ratio measured in volcanic gasses, with sulfur (i.e., mostly SO2 outgassing) 1683 being affected by complex processes during transcrustal magma transport (e.g., Mungall et al., 2015). In 1684 spite of these unknown, it appears that arc magmas have higher CO2 concentrations than MOR basalts; in 1685 terms of total CO2 fluxes however, emissions from mid-ocean ridges and arcs are, within uncertainties, 1686 equivalent (Kelemen and Manning, 2015; Figure 4). Below we analyze the dynamic of carbon transfer 1687 throughout the plumbing system of arc volcanoes. 1688 The solubility of carbon in molten basalts is low and it increases broadly linearly with pressure so 1689 that it reaches ~1 wt.% at pressure equivalent to the base of the continental crust (~1 GPa, corresponding 1690 to 30 km depth; Iacono-Marziano et al., 2012). As most other volatile components (i.e., water, sulfur and 1691 halogens) are much more soluble in silicate melts than CO2, CO2 is the first volatile component released by 1692 magma upon their transcrustal journey. Considering furthermore that most rising melt remains locked in 1693 the crust and evolves as intrusive bodies, this implies that a large portion of magmatic CO2 in arcs is 1694 outgassed at great depth and must exist as supercritical fluids in the crust. Variation in melt compositions 1695 can affect CO2 solubility and thus the depth of fluid saturations (Iacovino et al., 2021; Wieser et al., 2021). 1696 Alkaline primary melts dissolve more CO2 and thus reach fluid-saturation at shallower depth. If primary 1697 arc basalts contain 1 wt.% CO2, depending on melt compositions they will start forming a supercritical 1698 CO2-rich fluid phase at depths ranging from 30 to 10 km as they ascend through the crust. These CO2-rich 1699 deep fluids rise through the crust in a way that can be decoupled from the magmatic flow (Edmonds et al., 1700 2022). Thus, CO2 can be released from stagnant melts in the crust (non-eruptive), which is not the case of 1701 other magmatic volatile components bar noble gasses. It is thus difficult to estimate and model the global 1702 emissions of CO2 by arc volcanoes since these emissions start from the deep crust, yield surficial fluxes 1703 that are very diffused (Chiodini et al., 2005; Werner et al., 2019), and can originate from a combination of 1704 64 deep diamonds close to 100 Ma (Bulanova et al., 2010) and in general it would appear that super-deep 1892 diamonds are younger than lithospheric diamonds (Smit et al., 2022). This still remains to be fully 1893 understood and a definitive answer will be provided once more super-deep diamond ages will be available. 1894 65 6. Secular variation of surficial versus deep carbon reservoirs or 1895 steady state? 1896 There has been significant discussion on whether the shallow Earth deep carbon cycle is balanced, i.e., 1897 if total subduction fluxes are accounted for by total volcanic degassing (Kelemen and Manning, 2015; Plank 1898 and Manning, 2019; Wong et al., 2019; Müller et al., 2022). An analysis of carbon flux balance for inputs 1899 to and output from the mantle suggest a steady state regime for the modern era, albeit with large 1900 uncertainties. Plank and Manning (2019) concluded that the highly variable speciation and quantity of 1901 subducted carbon, the diversity of pressure-temperature paths along slabs, and the possible crustal 1902 contamination by various sediments render the assessment of carbon inputs to versus outputs from the 1903 mantle highly dependent of the global geodynamic configurations. Connections between the Wilson cycles 1904 and global mantle carbon inputs/outputs are thus expected, with some links with the surficial and deep 1905 carbon reservoirs formed across the history of planet Earth. As briefly discussed in Section 2, Gaillard et 1906 al. (2022a) concluded that the mass of surficial carbon formed in the early Hadean, once the magma ocean 1907 solidified, was identical, within uncertainties, to the present-day mass of surficial carbon. This may rather 1908 indicate a steady state regime for the inputs to and output from the mantle back to 4.5 Ga. Nevertheless, 1909 secular variations in the distribution and amounts of different forms of carbon, such as organic and inorganic 1910 carbon, are well documented. To list only a few notable examples, the Great Oxidation Event has imparted 1911 substantial redox changes in both surface and deep environments, affecting the stability of different carbon 1912 forms; the Cambrian explosion led to a remarkable increase in the productivity of organic carbon (Husson 1913 and Peters, 2017); lastly, the Marine Mesozoic Revolution may have marked an unprecedented change in 1914 the deposition and accumulation of inorganic carbon in marine carbonates (e.g., Müller et al., 2022; 1915 Ridgwell, 2005). Although these secular variations may have not impacted the global fluxes of carbon 1916 inputs and outputs, and may have been promoted by deep-Earth processes (Eguchi et al., 2020), their 1917 chemical fingerprints can be traced through geological time in the rock record (e.g., Kump et al., 2011; 1918 Giuliani et al., 2022). 1919 Modern Earth carbon ingassing and outgassing of the mantle is summarized in Figure 11. Following 1920 Plank and Manning (2019), we select recent published carbon flux values for subducted slab lithological 1921 components. In addition, we review recent estimates of carbon outgassing fluxes at tectonic settings. To 1922 obtain a total carbon in/outgassing flux, we perform Monte Carlo simulations to obtain a median estimate 1923 and uncertainty interval. We randomly select an individual carbon flux estimate from each flux subcategory 1924 shown in Figure 11 (e.g., selecting one of the three subducting oceanic crust carbon estimates), and 1925 subsequently randomly sample a carbon flux value from the error range of that estimate (i.e., the range 1926 66 denoted by the error bars in Figure 11), assuming that the uncertainty is uniformly distributed within that 1927 range. An individual estimate for total in/outgassing is thereby obtained from the sum of each individual 1928 sampled flux value contributing to carbon in/outgassing. The in/outgassing fluxes presented in Figure 11 1929 are median values ±2 standard deviations (n = 100,000). 1930 On a first order we note that our values for in/outgassing are broadly in concordance with Plank 1931 and Manning (2019), in that there is rough parity within 2 standard deviations between ingassing and 1932 outgassing. However, the ingassing flux we present (107±26 Mt C yr-1) is slightly larger than the outgassing 1933 flux (70±36 Mt C yr-1). We attribute this to several reasons. Firstly, there is significant uncertainty in all 1934 fluxes presented in Figure 11, thereby complicating an appropriate comparison. We also note that several 1935 fluxes are introduced or are revised relative to Plank and Manning (2019), such as the role of peridotite, 1936 which is now divided into three categories based on the circumstances of peridotite carbonation. In 1937 particular we highlight the introduction of sub-oceanic lithospheric mantle as a reservoir for oceanic carbon, 1938 which has only recently been quantified (Keller et al., 2017; Müller et al., 2022). We also consider the large, 1939 uncertain rift fluxes of Wong et al. (2019), which are substituted for the intraplate diffuse fluxes of Plank 1940 and Manning, but do not account for the degassing of individual rift-related volcanoes in these settings. We 1941 also note that estimates of fluxes of solid carbon exhumed in metamorphic rocks is essentially unquantified 1942 (Plank and Manning, 2019). Finally, we note that there are several processes during subduction that are not 1943 considered here, such as sequestration of carbon into the forearc slab, mantle and crust as carbonate and/or 1944 reduced carbon (Barry et al., 2019; Debret et al., 2022; Fullerton et al., 2021; Hu et al., 2021; Piccoli et al., 1945 2019; Vitale Brovarone et al., 2017), which will affect the carbon that is outgassed during volcanism. It 1946 therefore remains an open question as to whether the balance between subduction and outgassing remains 1947 at parity at the present day (e.g., Kelemen and Manning, 2015), and how this balance may be sustained (if 1948 at all) through geological time (e.g., Müller et al., 2022; Wong et al., 2019). 1949 1950 67 7. Concluding remarks 1951 Efforts from a broad range of scientific disciplines have provided a substantial contribution to the 1952 understanding of carbon reservoirs and movements in the deep Earth. However, despite deep carbon being 1953 the largest carbon reservoir on Earth, it remains the least understood. This can be depicted by the large 1954 uncertainties in deep carbon reservoirs and fluxes (see Figure 4), which stems from a partial understanding 1955 of the range and distribution of deep forms of carbon and related processes. Considering the two most 1956 striking peculiarities of our planet, life and subduction, fundamental questions remain on how, to what 1957 extent, and since when life has influenced deep carbon. What is the response time of deep carbon to major 1958 surface biogeochemical changes? Can carbon isotope signatures classically interpreted as biotic (or 1959 contaminated by biotic carbon materials) be explained by purely abiotic processes? How much and how 1960 diverse are the abiotic organics potentially formed in the interior of planet Earth, thereby potentially 1961 rejuvenating Gold’s theory (Gold, 1992)? Would the answer challenge our models on when the first 1962 biogenic carbon reached the convective mantle? To what degree does the lack of a sound quantification of 1963 diffuse carbon degassing at the Earth’s surface affect our models of deep carbon inputs and outputs? The 1964 current diversification of geobiological approaches to deep carbon studies still suffers from a jungle of 1965 semantic obstacles regarding the types of carbon and its origins. Simple terms such as organic/inorganic 1966 carbon, abiotic, biotic, or biogenic carbon are used in multiple ways and with very diverse meanings among 1967 emerging communities. In turn, this growing diversification of geo-biological disciplines represents a 1968 fundamental milestone in the understanding of deep carbon, and how it has been shaping the coevolution 1969 of Earth and life since at least 3.8 Ga. It also offers an opportunity to question processes on other planetary 1970 bodies, as the James Webb Space Telescope continues to reveal new, detailed observations of exoplanets 1971 and their intrinsic processes. 1972 1973 Acknowledgements 1974 This chapter builds on the significant increase in knowledge gained under the auspices of the Deep Carbon 1975 Observatory decadal program. This work is part of the project that has received funding from the European 1976 Research Council (ERC) under the European Union’s Horizon 2020 research and innovation programme 1977 (Grant Agreement No. 864045, acronym DeepSeep). A MUR FARE (acronym DRYNK) grant to AVB and 1978 grant and MUR PRIN2022 (Grant No. 20224YR3AZ; acronym HYDECARB) to AVB and DG are also 1979 acknowledged. FG acknowledges the GASTON project (ANR-18-CE31-0021). 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The geological history of deep carbon summarized in a clockwise time line sketch (see Section 2 for details). 3504 It shows the evolution of deep carbon from the early accretion of planet Earth, where magma ocean processes prevailed 3505 and drove core-mantle-atmosphere exchanges, to the present-day solid-state convection regime wherein mantle-crust-3506 atmosphere exchanges are dominated by plate tectonics. 3507 121 3508 Figure 2. Summary of the major biological processes interacting with carbon that have direct effects and interactions 3509 with the deep carbon cycle. Major processes are colored according to the division between terrestrial (green), marine 3510 (blue) and deep subsurface (red) processes. Global estimates of biomass are reported (Bar-On et al., 2018). Gross 3511 (photosynthetic) primary productivity on land: 110–150×1015 g C yr-1 (Jian et al., 2022; Lu et al., 2024). Organic 3512 carbon diagenesis: 64–120×1012 g C yr-1 (Kump and Arthur, 1999). Heterotrophy at the ocean surface: 143 ×109 g 3513 C yr-1 (del Giorgio and Duarte, 2002). Gross (photosynthetic) primary productivity in the ocean: 103–150×1015 g C 3514 yr-1 (Huang et al., 2021). Water column chemosynthesis (in the photic zone, dark ocean and surface marine sediments): 3515 ~1×1015 g C yr-1 (Middelburg, 2011). Marine organic carbon burial and diagenesis: 20–200×1012 g C yr-1 (Cartapanis 3516 et al., 2018; Dunne et al., 2007; Hayes et al. 2021; Jahnke, 1996). Oceanic sediment methane production is 75–3517 320×1012 g C yr-1 (Valentine, 2002). Utilization of deep volatiles in the continental crust (and oceanic crust) is still 3518 unconstrained (subscript a), but likely to be significant; see Barry et al. (2019) and Fullerton et al. (2021) for estimates 3519 at the Costa Rican convergent margin (~108–1010 moles C yr-1). Both continental and oceanic subsurface 3520 chemosynthesis is currently unconstrained (subscript b). 3521 128 3603 3604 Figure 7. Organic carbon degassing processes. (a) Biotic pathways of solid organic carbon mobilization to form 3605 hydrocarbon oils and gasses. Van Krevelen diagram showing the transformation pathways for kerogens and 3606 mobilization of carbon as biotic hydrocarbon oils and gasses during carbonization and successive graphitization 3607 (modified from Buseck and Beyssac, 2014, after Vandenbroucke and Largeau, 2007). Kerogen types I, II, and III 3608 derive from lacustrine algae, marine micro-organisms, and terrestrial plants, respectively, from. The immature zone 3609 represents conditions prior to the release of large amounts of oil and gas hydrocarbons. The diagram also shows 3610 compositional data of graphitic carbon from selected metamorphic rocks from Grew (1974). (b-c) Abiotic pathways 3611 of solid organic carbon mobilization in aqueous fluids. Diagrams calculated with the Thermotopes-COH software 3612 (Boutier et al., 2024a). (b) C-O-H thermodynamic model at 500 °C and 1 GPa. The lower part of the diagram, below 3613 the carbon saturation curve, corresponds to carbon undersaturated conditions (no graphite/graphitic carbon/diamond 3614 stable). See Holloway (1984) and Connolly and Cesare (1993) for details. The subfigure also shows the predicted fluid 3615 compositions for carbon-saturated systems at water-maximum conditions (ΔFMQ = -0.64 at the model conditions), 3616 ΔFMQ = -2, and ΔFMQ = +0.5 for a carbon activity (aC) equal to one (graphite stable). Note the higher carbon 3617 concentrations in more reduced or more oxidized carbon-saturated fluids relative to the so-called water maximum. 3618 The carbon saturation curve for a less graphitic carbon form is also shown and calculated with aC equal to 10. Note 3619 the higher solubility for aC > 1, corresponding to less graphitic carbon forms less crystalline than graphite. (c) 3620 Pressure-temperature diagram showing total dissolved (XC), calculated by summing the molar fractions of CH4 and 3621 CO2 (XCH4 and XCO2, respectively), for a graphite/diamond-saturated fluid at water maximum. At these conditions, 3622 CH4 and CO4 have the same concentration. Other C-bearing fluid species are in negligible concentrations. It can be 3623 seen that the solubility of graphite/diamond increases with temperature and decreases with pressure. 3624 3625 129 3626 Figure 8. Composition (a) and volume (b) of magmatic liquids equilibrated in the ambient convective upper mantle 3627 in terms of SiO2, CO2 and H2O, as calculated with MAGLAB (Massuyeau et al., 2021). The inset represents the 3628 adiabatic profile used for the calculations, considering a normal mantle with 140 ppm CO2 and 240 ppm H2O. 3629 3630 130 3631 Figure 9. Comparison of solidus curves for a bulk peridotite composition in volatile-free (bold black line), H2O-3632 bearing (thin blue line), and CO2-bearing (dashed red line) systems, presented in a temperature and pressure/depth 3633 diagram. Dry solidus after Hirschmann (2000); H2O-bearing solidus after Hirschmann et al. (2009) for 50 and 750 3634 ppm H2O; CO2-bearing solidus from Dasgupta (2013, 2018). The solid mantle adiabatic pressure-temperature path 3635 corresponds to a potential temperature of 1350 °C. fO2 variations versus depth as well as the nature of the various 3636 carbon-bearing phases (fluid-carbonate-diamond/graphite) are indicated on the left part of the figure. 3637 3638 131 3639 Figure 10. Diagrams concerning diamond age and isotopic composition, redrawn after Howell et al. (2020). (a) 3640 Illustration showing the historical occurrence of diamond growth events on Earth (Figure 1 in Howell et al., 2020). 3641 Diamond symbols represent eclogitic diamond growth events; square symbols represent peridotitic diamond growth 3642 events, with additional color differentiation for harzburgitic diamonds (magenta) and lherzolitic diamonds (blue); 3643 circle symbols represent fibrous growth events, which constrain the age of the corresponding kimberlite eruption. (b) 3644 Box plot showing δ13C data recorded from diamonds (Figure 7 in Howell et al., 2020). 3645 3646 132 3647 Figure 11. Figure illustrating the carbon flux balance into (left) and out of (right) the mantle, following Plank and 3648 Manning (2019). All values shown are in Mt C yr-1, and represent the selected value as illustrated by the solid bars. 3649 The error bars show uncertainty in the estimate. The total input and output fluxes (±2 standard deviations) are 3650 determined by Monte Carlo error propagation assuming uniform uncertainties in individual values and by randomly 3651 selecting an estimate within each contribution (see Section 6 for further details). References, in order of first 3652 appearance from left to right: A+13: Alt et al., 2013; M+22: Müller et al., 2022; KM15: Kelemen and Manning, 2015; 3653 L+19: Li et al., 2019; C17: Clift, 2017; [B+17, Wo+19]: the two models of Brune et al. (2017) combined with present-3654 day average rift degassing flux of Wong et al. (2019); H+18: Hauri et al., 2018; T+18: Tucker et al., 2018; LV+19: 3655 Le Voyer et al., 2019; We+19: Werner et al., 2019; B+21: Bekaert et al., 2021; H+19: Hauri et al., 2019. 3656 3657