Fate of Carbonates in the Earth’s Mantle (10-136 GPa)

Earth carbon cycle shapes the evolution of our planet and our habitats. As a key region of carbon cycle, subduction zone acts as a sole channel transporting supracrustal carbonate rocks down to the mantle, balancing carbon budget between the Earth’s surface and the interior, and regulating CO₂ concentration of the atmosphere. How carbonates evolve at depth is thus, a most fundamental issue in understanding carbon flux and carbon sequestration mechanism in the Earth. This study reviews prominent progresses made in the field of crystal chemistry of carbonates along subduction geotherms. It clearly finds that, in addition to common carbonates in the Earth’s crust, several new polymorphs of carbonates have been discovered to be stable under high-pressure and high-temperature conditions. This opens possibilities for oxidized carbon species in the deep Earth. However, metamorphic decarbonatation and reduction reactions restrict subducting carbonates to the top-mid region of the lower mantle. Specifically, subsolidus decarbonatation in the form of carbonates reacting with silicates has been proposed as an efficient process releasing CO₂ from slabs to the mantle. Besides, carbonate reduction in the metal-saturated mantle likely results in generation of super-deep diamonds and a considerable degree of carbon isotope fractionation. Review of these novel findings leads us to consider three issues in the further studies, including 1) searching for new chemical forms of carbon in the mantle, 2) determining the reduction efficiency of carbonates to diamonds and the accompanying carbon isotope fractionation and 3) concerning carbon cycle in subduction of continental crust.

Introduction

Carbon is a basic element that forms the Earth and the biosphere. Earth carbon cycle ties closely with the global climate, the origin and evolution of life and the carbon-based energy resources. The present review focuses on the carbon cycle in subduction zones. The crustal carbon is carried down to the Earth’s interior via subduction processes, modifying the redox state of the mantle, and the deep carbon exchanges with the crustal carbon via volcanism or metamorphic decarburization over geological timescales (Dasgupta and Hirschmann, 2010; Plank and Langmuir, 1993). This process is responsible for the recycling of a great amount of carbon and consequently, exerts profound influence upon the terrestrial ecosystem functions. In recent decades, the carbon balance between the Earth’s surface and the interior is wrecked. It has become our urgent desire to examine carbon sequestration capacity of different reservoirs and strive for a balance of carbon income and outcome. Extensive studies have been carried out on carbon flux from the Earth’s surface to the interior but obtained totally different results (Gorman et al., 2006; Johnston et al., 2011). Estimation on the removal amount of the subducted carbon from slabs varies from 35 to 100% (Kelemen and Manning, 2015). The large inconsistency is due mainly to our limited knowledge on behaviors of carbonates in subduction zones.

Carbonates (e.g. calcite CaCO₃, magnesite MgCO₃ and dolomite (Ca, Mg)CO₃) are predominantly carbon species in the Earth’s crust. Their structural stability in subduction zones is a key to measuring carbon flux. On the one hand, with the increase of the subduction depth, the pressure-temperature and the oxygen fugacity conditions become more and more favorable for reduced carbon species. In this respect, carbonates would be overtaken by diamond or graphite at mantle depths in excess of ∼250 km where the redox state begins to be buffered by iron-wüstite equilibrium due to disproportionation of Fe²⁺ to Fe³⁺ and metallic Fe⁰ (Frost et al., 2004). At greater depths, Fe²⁺ disproportionation to Fe³⁺ and Fe⁰ progressively increases, which likely results to diamond domains in the deep Earth (Rohrbach and Schmidt, 2011). On the other hand, however, the possibilities of crustal carbonates in the upper-mid mantle have been suggested by thermodynamic calculations (Sun et al., 2018) and experiments (Ghosh et al., 2009; Litvin et al., 2014). Besides, carbonate associations have been frequently discovered as inclusions in the transition-zone and the lower-mantle diamonds (e.g. Brenker et al., 2007; Kaminsky, 2012; Kaminsky et al., 2013; 2016). This indicates that the crystal chemistry of carbonates and the carbon sequestration mechanism in the deep Earth are far more complicated than previously thought.

In this study, we systematically review the high-pressure behaviors of carbonates along from cold to the hottest subduction geotherms, aiming at three scientific issues: 1) Establishing a comprehensive phase relation of main carbonates in the Earth’s mantle. Our first concern is whether carbonates will melt and breakdown at high-pressure and high-temperature conditions, or preserve to greater depths even approaching the core-mantle boundary (CMB). 2) Next, close attention are paid upon the reduction mechanism of carbonates in reactions with the metallic mantle during subducting processes, as that would be an impenetrable barrier to carbonate subduction. 3) Meanwhile, decarbonization of carbonated peridotites and eclogites (including mid-ocean ridge basalt, MORB) by the surrounding silicates should also be examined, as these processes exhibit the possibility of CO₂ release as an equilibrium phase in the deep mantle, which sets fundamental control on mantle metasomatism.

Phase Diagram of Carbonates in the Mantle

CaCO₃, MgCO₃, (Ca,Mg)CO₃ are three main inorganic carbon substances in the Earth’s crust. They are also dominant carbon species recovered in sublithospheric diamonds. Carbonates show versatile behaviors such as phase transition, dissolution, melting, decomposition and decarburization at mantle conditions. For example, CaCO₃ exhibits exceeding solubility in aqueous fluids and may extract a considerable portion of carbon from slabs. But Ca²⁺ in carbonates would be replaced by Mg²⁺ upon sinking, and the solubility of dolomite reduces significantly if carbonates become Mg-rich (Farsang et al., 2021). Naturally Fe²⁺ substituting Mg²⁺ forms a complete MgCO₃-FeCO₃ solid solution, and a minor addition of Fe²⁺ significantly affect the phase boundary of carbonates especially at lower-mantle pressures (Liu et al., 2015; Tao et al., 2014). The following section presents a summary of phase relation of CaCO₃, MgCO₃ and (Ca,Mg)CO₃ and their ferrous series at depths.

Calcite CaCO₃

The stable polymorph of CaCO₃ at ambient conditions is calcite (R-3c, Z = 6) where [CO₃]²⁻ species form parallel layers and Ca²⁺ are six-fold coordinated by O²⁻. Upon pressures above ∼2 GPa, calcite reconstructively converts to aragonite (Pmcn, Z = 4) with [CaO₆]-octahedra turning into [CaO₉]-polyhedra. Aragonite has long been considered as the dominant CaCO₃-structure in the Earth’s interior from the upper mantle to the top of the lower mantle. Yet its density is lower than the principle constituents of the mantle (Litasov et al., 2007), which casts doubts on the presence of CaCO₃ in the deep Earth. Accumulations of experimental and theoretical data lately, however, create a whole new landscape of CaCO₃ inside the Earth. One significant finding is CaCO₃-VII (P2₁/c-l, Z = 4), which was put forward by Gavryushkin et al. (2017) as favorable in the pressure range of 30–50 GPa below1200 K (Figure 1). Bayarjargal et al. (2018) and Li et al. (2018) subsequently expanded its temperature scope to the melting point and suggest CaCO₃-VII as a predominant candidate at mantle depths of 690–1010 km. Increasing temperature would lower the pressure of aragonite → CaCO₃-VII transformation (Zhang ZG. et al., 2018). CaCO₃-VII is characterized with non-planar [CO₃]²⁻ units that are rotated by ∼45° relative to those in calcite/aragonite and are tilted ∼90° to one another. It is considered as an intermediate in the aragonite → post-aragonite transformation wherein the [CO₃]²⁻ units undergo a ∼90° rotation (Pickard and Needs, 2015; Smith et al., 2018). Conversion of aragonite to CaCO₃-VII is accompanied by an uprush of 2% for density, and an increase of 9, 2 and 8% for shear, longitudinal and average sound velocities. The difference in wave velocities increase by ∼6% when 10% CaCO₃ is present with respect to a pure pyrolitic composition (Bayarjargal et al., 2018). Therefore, the potential preferred orientation of CaCO₃-VII may contribute to seismic discontinuities at ∼700 km in the vicinity of slabs (Kaneshima, 2013; Yang and He, 2015; Bayarjargal et al., 2018).

FIGURE 1

FIGURE 1. Phase diagram of CaCO₃, MgCO₃, (Ca,Mg)CO₃ and the ferrous series under high pressures and temperatures. Equilibrium boundary between Gr. and Dm. is based on Bundy et al. (1996). Pressure-Temperature-profiles of hot subduction slab (red arrow) stagnant in the transition zone and the coldest slab (navy-blue arrow) penetrating into the lower mantle (an extrapolated curve) are referred to Syracuse et al. (2010). The gray zone indicates the Mantle Geotherm (Katsura et al., 2010). Curve [1] represents the phase boundary between Arag. and CaCO₃-VII and Curve [2] between CaCO₃-VII and Post-Arag. (Gavryushkin et al., 2017; Li et al., 2018). Transformation from Post-Arag. to P2₁/c-h is indicated by Curve [3] (Gavryushkin et al., 2017; Zhang ZG. et al., 2018). Phase relation between CaCO₃-VI (P-1) and Arag. (13.5–25 GPa; ∼1100 K) is denoted by Curve [4] (Pickard and Needs, 2015; Bayarjargal et al., 2018). Curve [5] is an extrapolated curve showing the transformation from Arag. to CaCO₃-V (disordered) (Litasov et al., 2017; Zhao et al., 2019) or an unknown structure. Melting curves of CaCO₃ and the boundary of decomposition of CaCO₃ melt are referred to Litvin et al. (2014) (Curve [6]), Zhang and Liu, (2015) (Curve [7]) and Li et al. (2017) (Curve [8]). Structural evolution from MgCO₃ to C2/m (Curve [9]) is based on Oganov et al. (2008) and Boulard et al. (2011). MgCO₃-P-1 has a limited stable regime between 85 and 101 GPa below ∼1100 K (Pickard and Needs, 2015). Spin crossover of Fe²⁺ in ferromagnesite (Curve [10]) is from Liu et al. (2015). Phase boundaries from Fe²⁺-rich Mg-carbonates-R-3c to Pmm2 (Curve [11]) is plotted by Liu et al. (2015) and to C2/m (Curve [12]) by Boulard et al. (2011; 2012; 2015). MgCO₃ melt and the decomposition to diamond are measured from Litvin et al. (2014) (Curve [13]) and Solopova et al. (2015) (Curve [14]). Phase transition from Fe-Dol. to Fe-Dol.-II is denoted by Curve [15] and further to Fe-Dol.-III by Curve [16] (Merlini et al., 2012). Abbreviations: Gr., graphite; Dm., diamond; Arag., aragonite; Post-Arag., post-aragonite; Fe-Dol., Fe²⁺-bearing dolomite; Fe-Dol.-II, Fe²⁺-bearing dolomite-II; Fe-Dol.-III, Fe²⁺-bearing dolomite-III.

In the meanwhile, Gavryushkin et al. (2017) took the initiative in observing the occurrence of aragonite-II (P2₁/c), another model structurally relevant to aragonite at 35 GPa and its absence at 50 GPa. The phase domain of aragonite-II was reproduced by Bayarjargal et al. (2018) under non-hydrostatic pressures of 30–38 GPa. Yet never aragonite-II was a pure phase in both studies. By contrast, at pressures of 23–38 GPa, Li et al. (2018) observed only CaCO₃-VII generation at temperatures up to 2300 K. Zhang ZG. et al. (2018) provided thermodynamic arguments for the marginal metastability of aragonite-II at 30–40 GPa, and with temperature rising, aragonite-II became increasingly unfavorable. It is open to question for the presence and stability of aragonite-II, which is subjected to kinetic barriers of equilibrium phase transformations in experiments and uncertainties in calculations.

At 42 GPa and 1400 K, a further phase transition occurs from CaCO₃-VII to post-aragonite (Pmmn, Z = 2) with the Clapeyron slope of the phase boundary being -233 K/GPa (Bayarjargal et al., 2018; Li et al., 2018). Post-aragonite exhibits much higher wave velocity anisotropies than major silicates and other carbonates (Huang et al., 2017). This reconstructive transformation leads to measurable decreases in shear (−7.0%) and longitudinal (−4.7%) wave velocities around 1010 km in consideration of the pyrolitic mantle composition with 10% CaCO₃ (Li and Yuen, 2014; Niu, 2014; Bayarjargal et al., 2018). Post-aragonite was once believed as the best host structure for CaCO₃ in the bulk of the lower mantle and it transformed to a pyroxene-like structure C222₁ at the CMB (Oganov et al., 2008). This notion, however, should be deliberated. Another pyroxene-phase P2₁/c-h as recommended above ∼65 GPa, which was evolved from collapse of P2₁/c-l without energy barriers (Pickard and Needs, 2015; Huang et al., 2017; Zhang ZG. et al., 2018). Both P2₁/c-h and C222₁ are featured with [CO₄]^4- species, but the former is made up of pyroxene chains stacked out-of-phase while the latter of parallel chains. In spite of pyroxene-phases, Lobanov et al. (2017) insisted that post-aragonite perseveres until ∼105 GPa, afterwards it undergoes the lowest-enthalpy path to become P2₁/c-h along with conversion of sp²-to sp³-hybridized carbon. But Smith et al. (2018) documented that CaCO₃-VII itself is capable of sp²-sp³ hybridization exchange and rather than any known CaCO₃-polymorphs, an entirely new phase came in above ∼50 GPa. There exist remarkable discrepancies as to the stability of CaCO₃-polymorphs under mid-mantle pressures as well as the phase boundaries, due likely to the kinetic hindrance of the transformations or limitations of different approaches, but sp³-hybridized carbon does merit special notice because of its thermodynamical stability in the megabar regime.

Melting of CaCO₃ was observed at ∼1880 K at 3 GPa (Irving and Wyllie, 1975) and monitored using Raman Spectroscopy above ∼1900 K at 10 GPa (Litvin et al., 2014), with a smoother slope by conductivity measurement (Li et al., 2017). The melting curve derived from First-principles calculations shifted slightly to lower temperatures, where a knee point occurred at 13 (1) GPa and 1970 (40) K that was attributed to conversion between CaCO₃-V and aragonite. Aragonite was detected in the recovered products in the pressure-temperature region where CaCO₃-V is supposed to be stable (Zhang and Liu, 2015; Li et al., 2017) due likely to the unquenchable character of CaCO₃-V. CaCO₃ had a broad area of congruent melting up to ∼3500 K, above which, it started to decompose. Decomposition of CaCO₃ resulted in generation of graphite that transformed to diamond if annealed to lower temperatures of 700–2700 K (Bayarjargal et al., 2010; Spivak et al., 2011). All these studies pointed out that melting and decomposition of CaCO₃ were realized at temperatures above most subduction geotherms. That is, almost the whole mantle falls into the solid regime of CaCO₃ polymorphs.

CaCO₃ with complex structural chemistry can be understood from the following two perspectives. 1) In calcite-members (CaCO₃, CdCO₃, MnCO₃, FeCO₃, ZnCO₃, MgCO₃, etc), CaCO₃ has the largest ionic radius of Ca²⁺ that can bear various pressures by adjusting coordination numbers (Oganov et al., 2008). 2) In aragonite-family (CaCO₃, SrCO₃, PbCO₃, BaCO₃, etc), CaCO₃ contains the closest matching in distance between Ca²⁺ and [CO₃]^2- species that results to the most variation in C-O bond lengths and deviation from [CO₃]²⁻-coplane (Antao and Hassan, 2009). The great variety of CaCO₃ polymorphs opens possibilities of stable CaCO₃ component in the deep mantle. According to the results above, we plot Figure 1 for to obtain an overview of phase evolution of CaCO₃ to be aragonite (≤∼25 GPa) → CaCO₃-VII (25–42 GPa) → post-aragonite (42–80 GPa) → P2₁/c-h or C222₁ (80–137 GPa), which is dominantly affected by pressure.

Magnesite MgCO₃ and Ferromagnesite

MgCO₃ maintained its rhombohedral structure (R-3c; Z = 6) up to at least ∼80 GPa and ∼2300 K, thereafter it turned into MgCO₃-II (C2/m; Z = 6) until the CMB (Isshiki et al., 2004; Maede et al., 2017) (Figure 1). MgCO₃-II is featured with [C₃O₉]⁶⁻-rings that consists of three [CO₄]⁴⁻-tetrahedra with asymmetric C-O bonds, and Mg²⁺ are in eight- or ten-fold coordination (Oganov et al., 2008; Boulard et al., 2011). Another silicate-like polymorph C222₁ was once proposed as the most thermodynamically favorable for MgCO₃ above ∼102 GPa (Panero and Kabbes, 2008), but it was dismissed later. By calculations, Pickard and Needs (2015) suggested a new phase of P-1 to be the best candidate at pressures of 85–101 GPa in a limited temperature regime that is slightly lower than the coldest subduction geotherm. That is the reason why in most attempts we observed MgCO₃ directly transformed to MgCO₃-II (Boulard et al., 2011; Maeda et al., 2017; Martirosyan et al., 2018). The stability of C2/m was also confirmed in Fe²⁺-rich MgCO₃ (Ferromagnesite-II) above 80 GPa between 1850 and 2500 K. Ferromagnesite-II hasd a higher density and was temperature quenchable. Especially it could accommodate a considerable amount of Fe³⁺ (magnetite Fe₄O₃ and/or a new phase of Fe₄(CO₄)₃) that resulted from self-redox reaction between Fe²⁺ and [CO₃]²⁻ (Boulard et al., 2011, 2012, 2015; Chen et al., 2018). Special attention has been paid on spin transition of Fe²⁺ that drives a prior-transformation of ferromagnesite from R-3c to Pmm2 at ∼50 GPa and 1400 K (Liu et al., 2015), along with an abrupt volume collapse. The phase boundary of ferromagnesite from high spin to low spin state was tightly constrained around 60 GPa for temperatures above ∼1000 K (Liu et al., 2015). There might be a positive correlation between the Fe²⁺-content and the spin transition pressure (Spivak et al., 2014). Or there is just a negligible compositional effect due to the long Fe²⁺-Fe²⁺ distances in the structure separated by [CO₃]²⁻ species (Lin et al., 2012; Hsu and Huang, 2016; Weis et al., 2017). Some doubts remain regarding the presence of intermediate-spin state (Hsu and Huang, 2016; Weis et al., 2017).

MgCO₃ melt congruently at temperatures above the Mantle Geotherm (Litvin et al., 2014; Solopova et al., 2015). An increasing separation in melting points between CaCO₃ and MgCO₃ with pressure was expected to partially account for the majority of carboantitic melt found on the Earth’s surface being highly calcic (Hammouda et al., 2015; Li et al., 2017). MgCO₃ decomposed to oxides and diamond above ∼3000 K (Spivak et al., 2013; Litvin et al., 2014; Solopova et al., 2015), and the energy cost for CaCO₃ is much higher (Santos et al., 2019). Melting curve of ferromagnesite deviated seriously from that of pure MgCO₃. Similar to FeCO₃, melting loop of ferromagnesite was asymmetrical and had a thermal minimum for Fe²⁺-rich compositions (Tao et al., 2014; Kang et al., 2015). At mantle conditions, ferromagnesite was energetically advantageous over decomposition (Boulard et al., 2011; 2012), which generated magnetite-magnesioferrite solid solution, graphite and CO₂ (Kang et al., 2016). So melting and decomposition of CaCO₃-MgCO₃ series bounded the phase fields of these carbonates, and that cast doubts on releasing free CO₂ in the deep Earth.

The elastic behaviors of MgCO₃ associate closely to its structural chemistry, which besides, helps to locate the presence of subducted carbonates in the mantle. At mantle pressures, the shear wave velocities in MgCO₃ were smaller than those in silicates. Marcondes et al. (2016) thus associated the low-velocity zones near the bottom of the lower mantle with the presence of MgCO₃. In the light of the density reduction, Litasov et al. (2008) announced that MgCO₃ had a limited effect on seismic profiles of principle assemblages in the mantle, but Yang et al. (2014)proposed to detect its presence by regionally seismic anisotropies because of its unusual high compressional (V_P) and shear (V_S) anisotropies. Existence of minor Fe²⁺ as well as the spin transition opened a new window on the elastic properties of ferromagnesite (Fu et al., 2017; Stekiel et al., 2017). For high spin state, the sound velocities decreased with Fe²⁺-content, whereas the anisotropy of the sound velocities increased. Upon pressure, anisotropies of V_P and V_S increased until spin transition occurred and decreased afterwards. So a drastic softening of sound velocities across spin transition may occur in slabs enriched with ferromagnesite in the mid-lower mantle. Increases of both V_P and V_S velocities were observed after spin transition. Enrichment of Fe²⁺ led to larger differences in sound velocities and velocity anisotropies for low spin state as well. In this regard, Fe²⁺-rich MgCO₃ regions may show higher seismic contrast compared with principle minerals in the deep mantle.

To sum up (Figure 1), along subduction geotherms MgCO₃ maintains initial R-3c structure up to ∼80 GPa and hereafter converts to C2/m, with sp²-sp³ transition further upholding the phase in the lowermost mantle. Ferromagnesite undergoes spin transition in the mid-mantle and changes to Pmm2-structure subsequently. The drop in the unit-cell and the accordingly uprush of the density expand the stability field of ferromagnesite to higher pressures, yet disproportionation of Fe²⁺ could highly disintegrate the structure at lower-mantle conditions. Temperature exerts a greater impact on phase relations of ferromagnesite relative to MgCO₃. Neither melting nor decomposition occurs in MgCO₃ or ferromagnesite over the entire mantle.

Dolomite (Ca,Mg)CO₃ and Ankerite

Dolomite-(Ca,Mg)CO₃ is a common carbonate mineral in marine sediments, carbonated eclogite and ophiolite (Tao et al., 2020; Shatskiy et al., 2021). At the topmost of the upper mantle, dolomite broke down into a mixture of magnesite-MgCO₃ and aragonite-CaCO₃, the field of the coexistence opening towards higher temperature with pressure (Tao et al., 2014). A monoclinic phase (dolomite-II) set in driven by rotation of [CO₃]²⁻ groups at 14–20 GPa under ambient temperature (Santillán et al., 2003; Merlini et al., 2019; Stekiel et al., 2019). This new phase was expected to enhance the thermodynamic stability of (Ca,Mg)CO₃ up to 40 GPa over MgCO₃+CaCO₃. Above 41 GPa, a triclinic structure (dolomite-III, Z = 8) formed (Merlini et al., 2019).

Naturally Fe²⁺ substitutes Mg²⁺ in dolomite with various contents. Incorporation of Fe²⁺ strongly decreased the order degree of the structure (Tao et al., 2014) and helprf maintain its stability to higher pressures and temperatures (Stekiel et al., 2019). A trigonal → orthorhombic transformation was firstly observed in (Ca,Mg)_1.9Fe_0.1(CO₃)₂ at ∼17 GPa and 300 K (Mao et al., 2011). This new orthorhombic phase persisted in upon laser-heating to ∼1600–1700 K at 27–31 GPa. At 36–83 GPa, it further transformed to a monoclinic structure. The pressure-induced spin transition of Fe²⁺ was manifested by a drop in the unit-cell volume at ∼47 GPa. This observation, however, was not reproduced in Fe-richer counterparts. A similar experimental study on a Fe²⁺-richer crystal CaMg_0.6Fe_0.4(CO₃)₂, suggested formation of two triclinic structures under pressures (Merlini et al., 2012). One was Fe-dolomite-II (P-1; Z = 4), a distortion of the dolomite structure associated by rotation of [CO₃]²⁻ groups, existing between 17 and 35 GPa. The coordination number of Ca²⁺ increased from six in Fe-dolomite to eight in Fe-dolomite-II. The other was Fe-dolomite-III that adopted the same P-1 symmetry but had eight formula units in the primitive cell. It exhibited structural relations to CaCO₃-III, as the [CO₃]²⁻ units are non-planar with carbon partly in four-fold coordination, some of which edge-shared with Ca²⁺ and Fe²⁺ polyhedra. Fe-dolomite-III had a broad stability domain to 43 GPa and the melting point (∼2600 K). Its density was demonstrably comparable to that of the mixture MgCO₃+CaCO₃ if taking into account the FeCO₃ component. The second-order transformation from Fe-dolomite to Fe-dolomite-II was similar to that of dolomite → dolomite-II, and it had been strongly confirmed by Merlini et al. (2019). But above 36 GPa, Merlini et al. (2019) suggested a first-order transition to a rhombohedral symmetry (R3) (Fe-dolomite-IIIb). In the pressure range of 70–120 GPa and at 2500 K, Fe-dolomite-IV with an orthorhombic symmetry (Pnma) formed. It was charactered with corner-sharing [CO₄]⁴⁻ units in three-fold rings, similar to MgCO₃-C2/m (Oganov et al., 2008). Different Fe²⁺ contents could be responsible for the disagreement. Previously, Solomatova and Asimow (2016) predicted a monoclinic phase C2/cas the stable structure for more Fe²⁺-enriched samples Ca(Mg_0.5,Fe_0.5)(CO₃)₂ and CaFe(CO₃)₂ up to 60 GPa. The C2/c phase cannot be achieved experimentally due to the large energy barrier and the correspondingly large volume collapse, resulting in conversion from dolomite to metastable structures.

Existing data has revealed the high-pressure behaviors of dolomite, as is summarized in Figure 1, but it is limited to draw a full understanding of the structural evolution of Fe-dolomite at pressures and temperatures. Issues deserve evaluation with respect to the effect of Fe²⁺-content and its spin transition on the thermal and elastic properties and the stability fields (Tao et al., 2014; Merlini et al., 2016), and possibility of energetically competitive structures. Besides, there remains a large area uncovered for pressures above ∼60 GPa. We now reach agreements on 1) a minor addition of Fe²⁺ stabilizes the rhombohedral structure of dolomite against decomposition and transformation under upper mantle conditions; 2) there stands a good chance for fine changes occurring at ∼17 and ∼35 GPa, and a decrease trend of crystal symmetry is expected with pressure.

Carbonates Reduction by Reduced Mantle

Given most subduction geotherms, CaCO₃, MgCO₃, (Ca,Mg)CO₃ and their ferrous series maintain solid polymorphs over a wide range of depths down to the CMB. But their presence in the mantle depends highly as well on oxygen fugacity, which is controlled by oxidation state of iron in minerals and melts (e.g. sulfides and silicates) (Bataleva et al., 2017). A large part of the mantle below 250 km is thought to be Fe⁰-sufficient and favorable for reductions of carbonates or carbonatitic melts, which is a potential barrier to carbonates further-subduction and a promising mechanism for sublithospheric diamond formation (Rohrbach and Schmidt, 2011; Stagno et al., 2011). Several carbon species uncovered from the lower-mantle diamonds, vary from oxidized phases (e.g. carbonates) to the reduced (e.g. methane, graphite and iron carbide) (Brenker et al., 2007; Kaminsky and Wirth, 2011; Kaminsky, 2012; Kaminsky et al., 2013, 2016; Smith et al., 2016). That strongly indicates that the redox reactions between subducted carbonates (oxidized) and metallic mantle (reduced) is responsible for sublithospheric diamond genesis (Figure 2).

FIGURE 2

FIGURE 2. Experimental pressure-temperature conditions and run products of carbonates-Fe⁰ system. Blue diamond and green square represent formation of diamond (Dm.) and iron carbide (Fe-C), respectively, with uncertainties estimated for conditions marked. Diamond formation is always along with graphite generation. The dashed black curve is extrapolated from Zhu et al. (2018) for indicating the onset boundary of MgCO₃-Fe⁰ interaction.

Several studies have been conducted in petrologic context of subducted carbonate assemblages in contact with metallic mantle. One representative study (Palyanov et al., 2013) at 6.5–7.5 GPa and 1273–1873 K documented that (Mg_0.9Ca_0.1)CO₃-Fe⁰ interaction generated carbonatitic melts, magnesiowustite (Mg,Fe)O, cohenite Fe₃C and metastable graphite. Increasing temperature led to a higher degree of partial melting but a decrease in Ca²⁺-component in carbonate melt. Diamonds nucleation from carbonatitic melts (≥1673 K) was expected with heavier carbon isotope relative to the initial carbonate and high concentrations of nitrogen (1000–1500 ppm). In contrast, at iron front where carbide (Fe-C) melt formed, diamond crystallized with lower δ¹³C value and only 100–200 ppm of nitrogen mostly in the form of single substitutional aggregation. These chemical characters are analogous to that of natural sublithospheric diamonds. Besides, the accompanying products, iron carbide, graphite and (Mg,Fe)O are characteristic inclusions in natural stones. For simplicity, at the topmost of the upper mantle, MgCO₃ reduction by Fe⁰ is described as:

${\text{MgCO}}_3\left(\text{magnesite}\right)+{\text{Fe}}^{\text{0}}\left(\text{metal}\right)\to \left(\text{Mg,Fe}\right)\text{O}+{\text{Fe}}_3\text{C}+\text{C}\left(\text{diamond}\right)$(R3-1)

Thermodynamic calculations showed that over the entire mantle pressure range, C+(Mg,Fe)O is more stable than MgCO₃+Fe-C. Thus, presence of residual carbides in the run products is due to incomplete reaction. A similar experimental work (Martirosyan et al., 2015) later reproduced this observation but for only formation of graphite rather than diamond. That can be accounted for by kinetic impeding for graphite-to-diamond conversion. Increase of temperature and addition of catalysts are two effective methods for diamond formation (Tomlinson et al., 2011; Palyanov et al., 2013).

At higher pressure-temperature conditions, the onset boundary of MgCO₃-Fe⁰ interaction was defined linearly crossing 1100 K at ∼14 GPa and 1375 K at ∼40 GPa. The onset of the reaction and diamond nucleation seemed to have nothing to do with partial melting of Fe-C, whichrequired much higher temperatures (Morard et al., 2017). Crystallization of diamond slowed down drastically in 14–18 GPa, corresponding to the mantle depths of 410–525 km (Zhu et al., 2018). That implied a high-pressure kinetic for the reaction and accordingly, a huge reduction in diamond production at the transition zone. The mechanism of the high-pressure kinetic was unknown but it provided a plausible explanation for the abundance of sublithospheric diamond at 250–450 km in contrast to the rarity below. In a broader pressure-temperature range of 70–150 GPa and 1500–2600 K, MgCO₃-Fe⁰ interaction was exemplified by (Martirosyan et al., 2018):

${\text{MgCO}}_3\left(\text{magnesite}\right)+\text{Fe}\left(\text{metal}\right)\to \text{FeO}+\left(\text{Mg,Fe}\right)\text{O}+{\text{Fe}}_7{\text{C}}_3+\text{C}\left(\text{diamond}\right)$(R3-2)

Transformation from MgCO₃-R-3c to C2/m around 80 GPa exerted little effect on the reaction path. Diamonds nucleated and grew directly from carbonates at subsolidus conditions. Coexistence of FeO+(Mg,Fe)O in the run products suggested an immiscibility gap in MgO-FeO system (Dorfman et al., 2018; Martirosyan et al., 2018). Naturally for worldwide sublithospheric diamonds, (Mg,Fe)O solid solution is the dominant (50–56%) lower-mantle inclusions which presents a broad range in magnesium index (Mg# = Mg/(Mg+Fe)_at., 0.36–0.90). Brazil, the most productive for population of transition zone and lower mantle diamonds, however, possesses (Mg,Fe)O of severely Fe²⁺-rich, up to Mg# = 0.36 (Kaminsky, 2012). An origin from the D″ layer was proposed as a mechanism to explain this discrepancy (Kaminsky and Wirth, 2011), but it seemed at odds with the immiscibility gap in MgO-FeO system. Factors such as temperature (Palyanov et al., 2013) and oxygen fugacity (Martirosyan et al., 2015) may be more possible to account for the variations of (Mg,Fe)O composition.

In comparison with MgCO₃, CaCO₃ reacting with Fe⁰ proceeded much slower. For CaCO₃-Fe⁰ coupling in closed systems, Ca-wüstites were identified after reactions at 3–5.5 GPa and 1573–1673 K (5–48 h) with solid carbon in the form of graphite (Chepurov et al., 2011). Similarly for deeper subduction geotherms, interaction of CaCO₃-Fe⁰ system at 4–16 GPa and 923–1673 K was established as (Martirosyan et al., 2016):

${\text{CaCO}}_3\left(\text{aragonite}\right)+{\text{Fe}}^{\text{0}}\left(\text{metal}\right)\to {\text{Fe}}_7{\text{C}}_3+{\text{CaFe}}_2{\text{O}}_3\left(\text{Ca}-\text{wűstite}\right)$(R3-3)

The sluggish kinetics of established CaCO₃-Fe⁰ interaction suggested that significant amount of carbonates could survive and sink to the transition zone (Martirosyan et al., a 2016). At 6 GPa and 1273–1473 K, redox reaction between CaCO₃+MgCO₃ mixture and metallic iron resulted in formation of Ca-(Mg,Fe)O and Fe₃C (Martirosyan et al., 2015). But under higher pressure-temperature conditions of 51–113 GPa and 1800–2500 K, the CaCO₃ component in (Mg_0.38Ca_0.59Fe_0.03)CO₃ seemed unresponsive in the reactions but remained in the phase of post-aragonite (Dorfman et al., 2018). Model of carbonates-Fe⁰ interaction was applicable to Fe²⁺-carbonates self-reduction, where Fe²⁺ disproportionated and reacted with C⁴⁺ to form diamond and Fe³⁺-oxides at lower mantle pressures (see Section 2.2; Boulard et al., 2011; 2012; 2015; Chen et al., 2018). But diamonds nucleated from Fe²⁺-carbonates self-reduction were too tiny to be analyzed.

For all the carbon-phases during carbonates-Fe⁰ interaction, a closer look leads us to check the behaviors of iron carbide (e.g. cohenite Fe₃C, Fe₇C₃). Iron carbide has been discovered as inclusions in sublithospheric diamonds (Kaminsky and Wirth, 2011) and silicates which are believed to originate at various depths spanning from the upper mantle to possibly the CMB. That indicates an active participation of iron carbide in deep carbon cycle. Iron carbides could form by solid-solid reaction in portions of the Earth’s interior where metallic iron and elemental carbon or carbonateare in mutual contact (Palyanov et al., 2013). It dissolves carbon with bulk content as high as ∼300 ppm, and maintains stable in the reduced mantle environments with oxygen fugacity values close to or below IW buffer equilibrium. Besides, multiple lines of evidence suggest a mantle with sulfide-rich metallic melt (Stachel and Harris, 2008; Zhang Z. et al., 2018). Owing to strong non-ideal mixing between carbon and sulfur in Fe-bearing alloy, the presence of sulfur significantly lower the solubility of carbon and trigger carbon exsolution from the carbide (Zhang Z. et al., 2018). The discovery of a carbide+sulfide+metal assemblage in sublithospheric diamonds shed lights on mechanism of diamond formation from Fe-C-S melt (Smith et al., 2016). Here we evaluate the phase relations of Fe-C-S systems in the mantle focusing on exsolution of diamond. Chemical reactions involving:

${2\text{Fe}}_3\text{C}+3{\text{S}}_2\to 6\text{FeS}+{2\text{C}}^{\text{0}}$(R3-4)

${\text{Fe}}_3\text{C}+3{\text{FeS}}_2\to 6\text{FeS}+{\text{C}}^{\text{0}}$(R3-5)

Most studies on Fe-C-S systems were performed at pressures ≤8 GPa (Bataleva et al., 2017). These studies reported efficient capabilities of Fe₃C to generate graphite (1473–1873 K), diamond (1573–1873 K) and pyrrhotite (FeS) in the presence of sulfur-bearing metasomatic agents; further temperature increase led to higher efficiency of diamond crystallization. The established Fe-C-S model was applicable to Fe-Ni-C-S system, where Fe-Ni alloy showed a greater accommodation of dissolved carbon with temperature up to 2600 K but at lower pressures (≤20 GPa) (Narygina et al., 2011).

Redox reactions between subducted carbonates and Fe²⁺-peridotites/eclogites can be taken as a simplified model of slab-mantle interactions processing in the deep zones. It was mainly controlled by low concentrations of carbonates and metallic iron at the CMB as well as low diffusion of chemical components (e.g. O²⁻, Fe²⁺) among silicates and carbonates. This is a limited process and set constraints for carbon speciation and protections for carbonates from reduction (Figure 2). Likely that a significant amount of (Ca,Mg,Fe)-carbonates may survive from Fe⁰-saturation boundary (Martirosyan et al., a 2016) at the upper mantle and the transition zone. But at the top of the lower mantle due to stagnation of subducting slabs, carbonates accumulate and are heated up. Carbonates-Fe⁰ interaction would be more probable for diamond production. The rate of graphite-to-diamond transformation was higher in a FeNi system. Higher pressure appeared to favor growth of cubo-octahedral diamonds, whereas octahedral diamonds grow at lower pressure conditions (≤10 GPa in MgCO₃ and ≤15 GPa in FeNi). The run products lacked the slip planes and lamination lines commonly associated with superdeep diamonds, and we may suggest that these deformation features are due to transport instead of growth conditions (Tomlinson et al., 2011). Accompanying with carbon freezing in the Earth, carbon isotope fractionation occurs. The carbon isotope composition in the mantle displays a well-defined peak around δ¹³C = −5‰ in an extensive range to the lowest of ∼−40‰. Subduction of organic carbon was advocated for the mantle regions with ¹³C depletion. Relatively, reduction of carbonates extracted ¹²C to a greater degree and it was more likely to contribute to the broad distribution of carbon in the mantle (Horita and Polyakov, 2015).

Carbonates Decarbonization by Silicates

In consideration of major minerals in basaltic plates and the mantle, carbonate stability should also be inspected by carbonates-silicates interactions under mantle conditions. This reaction sets fundamental control not merely on subduction limit of carbonate or CO₂ release in the mantle, but on physicochemical properties of the mantle such as initiation of melting by lowering the mantle solidus, generation of alkalic magma in oceanic islands and carbonate metasomatism in cratonic mantle. Two typical chemical reactions are appraised here:

${\text{CaCO}}_3+{\text{SiO}}_2\to {\text{CaSiO}}_3+{\text{CO}}_2$(R4-1)

${\text{MgCO}}_3+{\text{SiO}}_2\to {\text{MgSiO}}_3+{\text{CO}}_2$(R4-2)

SiO₂ is one of the abundant components in MORB. Along subduction geotherms, SiO₂ changes to its high-pressure polymorphs as coesite, stishovite, CaCl₂-type and seifertite successively (Murakami et al., 2003; Grocholski et al., 2013). According to experimental and theoretical results, even for the temperature profiles in cold subduction slabs, CaCO₃-VII reacted with SiO₂-stishovite above 23 GPa and 1500 K to produce CaSiO₃-perovskite and possibly free CO₂. High temperature greatly increased the tendency of partitioning Ca²⁺ from carbonates CaCO₃ to silicates CaSiO₃ (Zhang ZG. et al., 2018; Li et al., 2018), accounting for the observation of abundance of CaSiO₃-perovskite in sublithospheric diamonds compared to typical mantle mineralogy. Decarbonation of MgCO₃ by SiO₂-coesite was confirmed beginning at 6–7 GPa and above 1600 K, and the decarbonatation temperature increased sharply with pressure (Palyanov et al., 2002; Kakizawa et al., 2015). An addition of alkali-component to MgCO₃-SiO₂ system negligibly changed the reaction path, and the output diamond was close to the initial MgCO₃ in terms of carbon isotopic composition (Palyanov et al., 2002). At 28–62 GPa and 1490–2000 K (770–1500 km), MgSiO₃-perovskite and solid CO₂ (V or VI) formed from interactions between MgCO₃ and SiO₂-stishovite (Takafuji et al., 2006). With this regard, both CaCO₃ and MgCO₃ was unstable and would release CO₂ to the lower mantle for SiO₂-rich subducted basalts.

Calculations carried out at higher pressures of 50–200 GPa under static conditions (0 K) showed that CaCO₃ (post-aragonite) and MgCO₃ (R-3c) maintained stable but react with SiO₂ at closely the CMB (Oganov et al., 2008; Pickard and Needs, 2015; Santos et al., 2019). This was contradictory to the experimental observations (Maeda et al., 2017) where the reaction of MgCO₃-SiO₂ occurred at mantle depth beginning from ∼1000 km and that held true for all the mantle regions below. Different results could be due to different conditions in these studies. MgCO₃-R-3c reacted with SiO₂-stishovite for formation of MgSiO₃-bridgmanite and CO₂-VI (six-fold coordinated of carbon) at depths around 1700 km (83 GPa; 1780 K). Internal production of CO₂ from MgCO₃-SiO₂ system was thus, possible in the middle of the lower mantle (1700–1900 km) and at greater depths CO₂-VI would decompose to diamond. If survived, MgCO₃-R-3c changed to C2/m and reacted with SiO₂-seifertite to generate MgSiO₃-post-perovskite and diamond at the lowermost mantle.

Figure 3 summarizes the phase relations of carbonate-silicate system at mantle conditions. The solidus of the carbonated MORB intersected most subduction geotherms at mantle depths of 300–700 km, creating a focused region of melt at the top of the transition zone and producing a barrier to carbonate recycling into the deep mantle (Thomson et al., 2016). But in super-cold slabs, carbonated MORB would survive from the melting process and sink further. Carbonation reactions involving olivine or pyroxene tended to react with CO₂ to form carbonate without dissociation into a CO₂-bearing assemblage, so present-day subduction of oceanic crust likely introduced carbon into the mantle in the form of carbonated eclogite. Phase relations of CaO-MgO-Al₂O₃-SiO₂-CO₂ system exhibited a substantial increase in MgCO₃ component in melt compositions from lower pressures (12–16 GPa) to the higher (20–25 GPa). That may be due to CaSiO₃-perovskite formation at 20 GPa but MgSiO₃-perovskite required higher pressures (Keshav and Gudfinnsson, 2010). For carbonates with high bulk alkali contents, the alkali-rich carbonates tended to form a solid solution with Ca²⁺ but left Mg²⁺ incorporating into majorite. What is striking was that most carbon in SiO₂-rich melts partitioned into alkali-rich aragonite and diamond, whereas little diamond formed in SiO₂-deficient melts. It may be accounted for by more progressive Fe²⁺-Fe³⁺ transformation with pressure in SiO₂-rich melts (Kiseeva et al., 2013). Along subduction geotherms from the upper mantle to the transition zone, carbon-phases in alkali-rich carbonated eclogites evolve following the route of “MgCO₃+CaCO₃ (aragonite) or MgCO₃-CaCO₃ (calcite) solid solution → MgCO₃+alkali-rich CaCO₃ (aragonite) → alkali-rich CaCO₃ (aragonite) → graphite and diamond.”

${\text{CaCO}}_3+\text{MgO}\to \text{CaO}+{\text{MgCO}}_3$(R4-3)

${\text{CaCO}}_3+{\text{MgSiO}}_3\to {\text{CaSiO}}_3+{\text{MgCO}}_3$(R4-4)

FIGURE 3

FIGURE 3. Schematic reaction line of carbonate-silicate at the Earth’s mantle conditions. The black lines representing CaCO₃-SiO₂ reaction is measured according to Li et al. (2018) and Zhang and Liu (2015). The dark-blue lines indicate phase boundaries for MgCO₃-SiO₂ system (solid: Maeda et al., 2017; dotted [1]: Takafuji et al., 2006). Partial melting of MgCO₃-SiO₂ series produces a simple eutectic system (Kakizawa et al., 2015). Curve [2] represents solidus curve of carbonated peridotite (Litasov and Ohtani, 2009). Solidus curves of carbonated eclogites are plotted according to Litasov and Ohtani (2010) (Curve [3]) and Kiseeva et al. (2013) (Curve [4], [5]). The melting curve of carbonated MORB (Curve [6]) according to Thomson et al. (2016). Phase diagram for the stability of MgCO₃+CaSiO₃ versus CaCO₃+MgSiO₃ is referred to Santos et al. (2019). Phase boundaries of CaCO₃- and MgCO₃-polymorphs follow the references in Section 2.1. The phase diagrams of SiO₂ (Murakami et al., 2003; Grocholski et al., 2013), CO₂ (Litasov et al., 2011) and MgSiO₃ (Komabayashi et al., 2008) are included with gray solid lines. Abbreviations: Gr., graphite; Dm., diamond; St., stishovite; CS., CaCl₂-type SiO₂; Se., seifertite; Brd., MgSiO₃ bridgmanite; Ppv., MgSiO₃ post-perovskite.

Calculations indicated that the reaction R4-3 was energetically unfavorable over the entire mantle. That is, CaCO₃ is stable under thermodynamic conditions if there is MgO in excess as expected in a pyrolitic mantle (Oganov et al., 2008; Pickard and Needs, 2015; Santos et al., 2019). For the conditions of MgSiO₃ in excess (R4-4), which was the main mineral in the lower mantle, CaCO₃ would be consumed and the MgCO₃+CaSiO₃ association was more favorable to the mid-mantle, hereafter CaCO₃ became restabilized over MgCO₃ (Oganov et al., 2008; Pickard and Needs, 2015; Zhang ZG. et al., 2018; Santos et al., 2019). The higher densities of the CaCO₃-polymorphs relative to the MgCO₃-polymorphs could be responsible for the free energy change of the reaction. However, temperature and chemical potentials need to be taken into account further as they can shift the phase relations conceivably. Santos et al. (2019) believed that an increase in temperature increased the pressure at which the reaction becomes unfavorable. That is, MgCO₃ was the main host of oxidized carbon over almost the mantle. But for super-cold subduction geotherms, CaCO₃ was more stable than MgCO₃.

As is highlighted in Figure 3, during deep-subduction processes, MgCO₃ exhibits much higher stability relative to CaCO₃ in consideration of existence of subducted SiO₂ and MgSiO₃ in a pyrolitic mantle model. Only for locally regions where MgO is in excess, CaCO₃ could be preserved to the CMB. Carbonatitic melts reacting with SiO₂ would release CO₂ (solid) into the lower mantle, which will eventually decompose to diamond.

Conclusions and Prospects

Carbonates play a key role in the recycling of carbon from the Earth’s crust to the interior via subduction processes, producing one of the carbon budgets in the entire Earth’s system. In light of the striking discoveries above, we now have a much clearer picture of behaviors of carbonates in the deep Earth.

First, several new polymorphs have been identified as stable of CaCO₃, MgCO₃ and (Ca,Mg)CO₃ and their ferrous series (Figure 1), which improve our expectations for their possibilities in the mantle. Existence of carbonates and the occurrence of the phase transitions may correlate with the discontinuities in the Earth’s seismic profile. Carbonates of CaCO₃-MgCO₃-FeCO₃ series remain in solid phases along most modern subduction geotherms. So liberation of carbonatitic melts or CO₂ owing to carbonate decomposition seems unlikely, with such possibility restricted merely for conditions in hot zones. However, decarbonization of CaCO₃ and MgCO₃ takes place in SiO₂-rich subducted basalts under the lower-mantle conditions (Figure 3). CO₂ decomposing to diamond is energetically favored in the region from the middle of the lower mantle to the CMB. Meanwhile, carbonates in slabs tend to be reduced to diamond at metal-saturation boundary at the topmost of the lower mantle (Figure 2). Diamonds produced from slab-mantle interactions are similar to natural stones in terms of chemical composition and inclusion characteristics. So far the existing data have covered almost the entire mantle conditions for decarbonization and reduction of carbonates and documented diamond nucleation. In this regard, carbonates reduction and decarbonization are effective mechanisms of carbon sequestration at the top-mid region of the lower mantle.

Researches on carbon cycle throughout the Earth’s system have long represented an interesting frontier. One fundamental goal of carbon cycle research is to know the balance between income and outcome global carbon fluxes as well as the carbon fluxes change in space and time. But many details of the processes remain unclear for the lack of intensive monitoring. Three issues may be worth of particular considerations in the future work:

(1) Besides carbonates, CO₂ and iron carbides mentioned above, there could exist other chemical forms of carbon in the mantle that has been uncovered, for example, a new carbon phase of Fe₄(CO₄)₃ derived from decomposition of Fe-bearing carbonates (Boulard et al., 2015). So a critical and pressing task is to search for these undiscovered carbon. Given the scarcity of natural samples, we may gain insights by thermodynamic calculations and experiments. These efforts will shed new lights on our view on deep carbon reservoirs, for example, how they form and under what conditions are they stable for long.

(2) Existing experiments have confirmed the possibility of redox reactions between subducted carbonates and metallic mantle to produce sublithospheric diamonds and other reduced carbon species. However, it is not yet clear whether carbonates-Fe⁰ processes generate significant amounts of diamonds. To what extent do carbonates turn into diamonds? And what role has the remaining carbonatitic melts played in deep carbon cycle and besides, in mantle metasomatism. These questions are basically related to the kinetics of carbonates-Fe⁰ interactions that set constraints on sequestration efficiency directly of deep carbon. Besides, carbon isotope fractionation during slab-mantle interactions is also an essential part in the study of deep carbon cycle. However, previous insights have been gained from theoretical calculations (Horita and Polyakov, 2015), as natural samples are too scarce. Experiments conducted at high-pressure and high-temperature conditions may be a favor to synthesize large samples.

(3) Now we have a better knowledge of carbon recycle in the carbonated MORB during deep subduction processes. Yet there is an increasing recognition that continental crust carrying sedimentary carbonate rocks can subduct to the upper mantle (Zhu et al., 2009). With depth, a certain amount of CO₂ is released from carbonatitic components via progressive metamorphism. Moreover, higher metamorphic temperatures enhance the activity of fluids in subduction zones, which significantly promotes decarbonization of carbonates (Zhang ZG. et al., 2018; Li et al., 2018). We believe that such reactions are a widespread process and exert an assignable effect upon carbon flux between the Earth’s surface and the interior. So equally important are efforts to be made on processes and rates by which carbon cycle in continental subduction zones.

Author Contributions

JG wrote the original draft. XW, XY, and WS were responsible for the review and modification.

Conflict of Interest

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Publisher’s Note

All claims expressed in this article are solely those of the authors and do not necessarily represent those of their affiliated organizations, or those of the publisher, the editors and the reviewers. Any product that may be evaluated in this article, or claim that may be made by its manufacturer, is not guaranteed or endorsed by the publisher.

Acknowledgments

The authors are thankful to the reviewers.

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