Carbon-Bearing Magmas in the Earth's Deep Interior

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CARBON-BEARING MAGMAS IN THE EARTH’S DEEP INTERIOR

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Konstantin D. Litasov, Anton Shatskiy Sobolev Institute of Geology and Mineralogy SB RAS, Novosibirsk, Russia

CHAPTER OUTLINE 1. Introduction .....................................................................................................................................43 2. Problems of Experimental Techniques to Study Systems With Volatiles................................................45 2.1 Multianvil Experiments ....................................................................................................45 2.2 Diamond Anvil Cell Experiments .......................................................................................46 3. Mantle Temperature and Silicate Solidi .............................................................................................46 4. Redox State in the Deep Upper and Lower Mantle ..............................................................................49 5. Melting and Phase Relations in the Carbon-Bearing Systems ..............................................................50 5.1 Systems With CO2 ...........................................................................................................51 5.2 FeeCarbide and CarbonateeFe Relations ..........................................................................56 5.3 The Role of Hydrocarbons and H2O in the Deep Carbon Cycle .............................................59 6. Deep Carbon Cycle, Melting, and Material Transport in the Earth’s Mantle...........................................62 6.1 Carbon Sources and Cycling in the Early Earth...................................................................62 6.2 Subduction, Melting, Kinetics, and Probable Stability of Carbonates at the CoreeMantle Boundary........................................................................................................................63 6.3 The Inevitable Fate of Carbonates at the CoreeMantle Boundary .........................................64 6.4 The Model of Carbonatite or Hydrocarbon-Bearing Diapirs...................................................66 7. Concluding Remarks.........................................................................................................................71 Acknowledgments ..................................................................................................................................71 References ............................................................................................................................................71

1. INTRODUCTION The nature and composition of igneous rocks originating in the asthenosphere strongly depend on carbon and hydrogen abundance and cycling throughout the Earth’s history (Dasgupta, 2013; Hirschmann, 2006; Litasov, 2011). Magmatic activity at depths exceeding 100e200 km is probably almost impossible without the involvement of volatile species, which can be described mostly in the CeOeH system. The role of other volatile compounds including S, N, P, and halogens is less important, but should not be underestimated. The best subsurface examples of magmatism caused by volatiles in the asthenosphere are kimberlites, related alkaline rocks, and carbonatites (e.g., Mitchell, Magmas Under Pressure. https://doi.org/10.1016/B978-0-12-811301-1.00002-2 Copyright © 2018 Elsevier Inc. All rights reserved.

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1995; Woolley and Church, 2005). The concentrations of carbon and hydrogen/H2O in the source region of these rocks may locally exceed 1 wt%. Major rock types ranging from abundant midocean ridge basalts (MORB) to most ancient komatiites are also considered to originate under volatilebearing conditions (Arndt et al., 1998; Kelemen and Manning, 2015; Resing et al., 2004; Sobolev et al., 2016). However, in this case the volatile abundance in the source region is much smaller, in the range of hundreds ppm. Unfortunately, significant devolatilization and postmagmatic alteration of volcanic rocks can substantially befog identification of their original composition and nature. For example, some primary magmatic carbonatites can be recognized only by relics of Cr-spinel grains (Bailey, 1989; Brady and Moore, 2012), whereas alkali-bearing species in kimberlites may survive only as inclusions trapped in minerals (Kamenetsky et al., 2014). The transformation of newly erupted natrocarbonatites to calciocarbonatite takes only several months or even days, representing an endmember of rapid alteration processes (Keller and Zaitsev, 2006). This means that components such as alkalis can be easily removed from the original rocks, and the rock compositions in geological sequences may be totally modified. The observation of reduced fluid species like hydrogen or methane after volcanic eruptions or intrusive magmatic activity is even more difficult, since they are strongly oxidized during residence at subsurface conditions. Some evidence for reduced fluid activity and their importance in the deep mantle originates from the studies of inclusions in minerals. Mantle-derived methane and more complex hydrocarbons can participate in kimberliteecarbonatite magmatism (e.g., Garanin et al., 2011), observed in island arc magmas (e.g., Schmidt et al., 2002), alkaline rocks (Potter and KonnerupMadsen, 2003; Sephton, 2013), and specific ophiolite sections such as podiform chromitites with unusual reduced inclusions and diamond (Yang et al., 2007, 2015). Indeed, some reduced fluid occurrences can be affected by methanogenesis in serpentinites from ophiolitic and modern oceanic environments (e.g., McCollom and Seewald, 2001; Proskurowski et al., 2008). Direct information about the possible nature of melts at the deepest levels of the lower mantle is provided by the inclusions in superdeep diamonds. These inclusions reveal the possible activity of both oxidized CO2-bearing and reduced, presumably CH4-bearing, fluids, identified by the appearance of carbonates (Brenker et al., 2007), carbides, and native iron (Kaminsky, 2012; Smith et al., 2016) in the same diamonds, which contain the lower mantle phases bridgmanite and ferropericlase. Our knowledge about magmas and material transport at depths below the lithospheree asthenosphere boundary (LAB) is limited. Yet much has been achieved in recent years, revealing a role for carbon and other volatiles in magmatic processes and deep mantle and core dynamics based on the observation of natural samples, experimental and theoretical modeling at high pressure and high temperature, seismological data on deep mantle heterogeneities, and related numerical geodynamic simulations. Major reviews related to the topic of this chapter include works on the deep carbon cycle (Dasgupta, 2013; Dasgupta and Hirschmann, 2010; Hammouda and Keshav, 2015), the deep hydrogen cycle (Hirschmann, 2006; Litasov and Ohtani, 2007; Ohtani, 2015), and mutual carbon and hydrogen activity in the deep Earth’s mantle (Hirschmann and Dasgupta, 2009; Litasov, 2011; Litasov et al., 2013b; Luth, 2014). Most of these works include comprehensive literature analyses and some works additional to those covered here can be found in those reviews. In this chapter, we review current data on the modeling of carbon-bearing melts and solids at high pressures deep in the Earth’s interior. We begin with remarks on experimental techniques and their uncertainties, a brief description of the thermal structure and oxidation state of the mantle, followed by

2. PROBLEMS OF EXPERIMENTAL TECHNIQUES TO STUDY SYSTEMS

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an analysis of the phase equilibria and melting in systems containing oxidized carbon as carbonates, and reduced carbon as hydrocarbons, diamond, and carbide. Finally, we provide a cross check of current models of carbon cycling and mantle magmatism originating from as deep as the coreemantle boundary (CMB).

2. PROBLEMS OF EXPERIMENTAL TECHNIQUES TO STUDY SYSTEMS WITH VOLATILES Useful techniques for high-pressure experiments on systems with volatile components using pistoncylinder and multianvil (MA) apparatuses and diamond anvil cells (DACs) in the laboratory and at synchrotron radiation facilities have been carefully reviewed elsewhere (Dunn, 1993; Katsura et al., 2004; Mao and Mao, 2007; Shatskiy et al., 2011). Here, we emphasize the uncertainties and possible sources of error in the experimental results. These errors cannot be avoided at the present status of experimental and analytical techniques, but they should not be ignored as a source of misinterpretation of the data even if published in top-level scientific journals.

2.1 MULTIANVIL EXPERIMENTS The major challenges in piston-cylinder and MA techniques are the careful control of volatiles during experiments in presumably volatile-free systems and control of the oxidation state during experiments with added volatiles. One problem is related to the presence of volatiles in starting materials and cell assembly parts, which is difficult to recognize before and during experiments, even after careful preliminary drying. In fact, it is almost impossible to avoid the presence of minor H2O-, CO2-, or Cl-bearing fluid (where chloride is used). Therefore, careful checking of the hydrogen and CO2 content in the run products is important (Shatskiy et al., 2009). The problem of starting materials that contain some unintended volatiles is also important for DAC, because sometimes they are synthesized in pistoncylinder or MA apparatuses and can be contaminated by volatile components (Nomura et al., 2014). Another major problem is controlling the escape of hydrogen from the high-pressure cell in pistoncylinder and MA experiments with reduced volatiles. To control fO2 in the sample, buffering techniques are required. In conventional experiments, for example, with H2O or CO2, the fO2 is controlled either by the choice of starting composition (i.e., mineral equilibria with a fluid phase) or by the fO2 imposed by the parts of the cell assemblage (e.g., BN, LaCrO3, graphite, metals). Buffering techniques that control fO2 in the sample during experiments typically use metalemetal oxide pairs (Jakobsson and Holloway, 1986), and can be considered as most suitable for studies of volatile-bearing systems. The double capsule method with hydrogen-transmitting medium and outer buffer capsule was introduced for piston-cylinder and pressless split-sphere apparatuses (Taylor and Green, 1988; Taylor and Foley, 1989; Sokol et al., 2010). However, this requires extra space around the sample and is difficult to use in the small cells of MA experiments at pressures above 20 GPa. To study phase equilibria involving reduced CeOeHeN fluid species, we developed a modified double-capsule method (Litasov et al., 2013b, 2014), where the thick-walled outer capsule serves as the buffer material (e.g., molybdenum or iron). In this case, the sample is placed into an AuePd or Pt capsule and separated from the outer capsule by talc, which transforms to enstatite and H2O-fluid or H2O-bearing silicate melt upon heating, thus increasing hydrogen fugacity in the sample cell up to the level corresponding to the buffer material. The drawback of this method is achieving equilibrium, because the

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double-capsule system has a lifetime limited by the availability of fluid in the hydrogen-transmitting media (between the outer and inner capsules). Thus, the duration of the experiment must be consistent with the system lifetime. True equilibrium will never be achieved since the composition and amount of fluid gradually change during the experiment.

2.2 DIAMOND ANVIL CELL EXPERIMENTS DAC experiments have serious limitations related to the achievement of chemical equilibrium, thermal gradient, and temperature measurements in the sample (Dewaele et al., 1998; Kavner and Nugent, 2008; Rainey et al., 2013). An additional problem is that different scientific groups use different criteria to model temperature gradient and uncertainties. The problem of equilibrium is especially difficult in melting experiments, where a melt pool may be separated from the bulk sample and pressure medium and remains unequilibrated. In some studies, researchers avoid any corrections on thermal pressure, which can exceed 10% at high temperatures, and measure pressure by ruby fluorescence only. Uncertainties in the pressure scales of calibrants are also essential at pressures above 50e100 GPa. We recommend using the recent pressure scales from Sokolova et al. (2013, 2016), where internally consistent equations of state were calculated for most of the important pressure calibrants and can be used for pressure recalculations. Indeed, different diffusion rates of the components at high temperature, pressure, and stress distribution in DAC also affect the results (e.g., Sinmyo and Hirose, 2010; Lobanov et al., 2015b; Mezouar et al., 2017). The control of fluid composition and fO2 in laser-heated DAC is even more difficult. Fluid escape from the heating zone of the DAC may be a major problem, as observed, for example, in methane dissociation experiments at pressures to 100 GPa (Lobanov et al., 2013). Nevertheless, the data from DAC experiments are the only available source of information for phase relations and physical properties of materials at pressures exceeding 30e50 GPa: we note, however, that the reader should be careful in applying results from DAC experiments directly and widely for geochemical and geodynamic modeling.

3. MANTLE TEMPERATURE AND SILICATE SOLIDI The assignment of temperature profiles in the mantle is very important to constrain melting models (see also Fiquet, 2017). The thermal regimes are different in the lithosphere, asthenosphere, and subduction zones. In the lithosphere, temperature increases rapidly with depth, whereas temperature profiles in the asthenosphere are nearly adiabatic and correspond to an average mantle potential temperature of about 1600 K (e.g., Wyllie, 1988). The depth of the transition from mainly conductive to convective thermal regimes varies from 10 to 20 km beneath midocean ridges to >250 km beneath the ancient cratons with thick lithosphere (Fig. 2.1). The temperature profiles in subducting oceanic plates are significantly lower than in the surrounding mantle, but it is difficult to constrain them at depths below 250e300 km. Reasonable profiles are indicated by lines 1e3 in Fig. 2.1 (Syracuse et al., 2010; van Keken et al., 2011), which show that the PT profiles of the hottest modern slabs may be similar to the cratonic geotherms. Some calculations predict the existence of very cold slabs (profile 4 in Fig. 2.1) (e.g., Kirby et al., 1996), which should stagnate in the transition zone and not penetrate into the lower mantle (King et al., 2015). According to the PT estimations for Precambrian metamorphic rocks, the PT profiles of ancient subduction might be significantly hotter than those of the modern

3. MANTLE TEMPERATURE AND SILICATE SOLIDI

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FIGURE 2.1 PT profiles in the Earth’s mantle to 900 km depth. The gray field shows the range of mantle adiabats with potential temperatures of 1590e1690 K. Note that olivineewadsleyite (410 km) and post-ringwoodite (660 km) transitions correspond to the lower bound of these adiabats. The dashed line K10 shows the 1673 K adiabat after Katsura et al. (2010). The dotted line SD-08 corresponds to an average mantle PT profile (Stacey and Davis, 2008). The Archean adiabat corresponds to the potential temperature estimated for komatiite w1973 K (e.g., Herzberg et al., 2010). Mantle plume temperatures correspond to the interval between Archean and modern adiabats. Numbers show the PT profiles of hot (1), medium (2), cold (3), and coldest (4) subduction slabs stagnant in the transition zone and penetrating into the lower mantle (e.g., 2a and 3a) based on estimates for depths of 50 and 250 km (bars correspond to slab surface and slab Moho level; Syracuse et al., 2010) and for depth of 660 km (King et al., 2015; Kirby et al., 1996). K, parameters for peak metamorphism in the high-pressure Kokchetav UHP-massif (Korsakov and Hermann, 2006). Brg, bridgmanite; Cpv, Ca-perovskite; Dia, diamond; Fpc, ferropericlase; Gr, graphite; Gt, garnet; Ol, olivine; Px, pyroxene; Rw, ringwoodite; Sp, spinel; Wd, wadsleyite. Modified from Litasov, K.D., Shatskiy, A., Ohtani, E., 2013b. Earth’s mantle melting in the presence of C-O-H-bearing fluid. In: Karato, S. (Ed.), Physics and Chemistry of the Deep Earth. Wiley, New York, pp. 38e65.

subduction and correspond to the modern oceanic geotherm, whereas most estimations for Phanerozoic rocks are consistent with modern hot subduction (Litasov et al., 2013b; Maruyama and Liou, 2005). Temperature estimations in the lower mantle have significant uncertainties. The mantle adiabat calculated using thermoelastic properties of minerals (Katsura et al., 2010; Stacey and Davis, 2008) corresponds to about 2750 K near the top of the D00 discontinuity (Fig. 2.2). Inside D00 there is a strong temperature gradient and temperature at the CMB is in the range 3800e4200 K (e.g., Litasov and Shatskiy, 2016; Fiquet, 2017).

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FIGURE 2.2 Mantle adiabat compared to solidi of nominally volatile-free silicate rocks including fertile peridotite (pyrolite) (Fiquet et al., 2010; Nomura et al., 2014), MORB-type eclogite (Andrault et al., 2014; Pradhan et al., 2015) and model volatile-free chondrite (Andrault et al., 2011). The data at pressures below 25 GPa are from Yasuda et al. (1994) and Zhang and Herzberg (1994). Lower line is solidus; upper line of the same color is liquidus (if present). CMB, coreemantle boundary. Differences between adiabat and solidi at the top and bottom of the lower mantle are indicated. Data below 25 GPa based on multianvil (MA) experiments and at higher pressures are from diamond anvil cell (DAC) experiments.

The melting temperatures of model peridotite and eclogite systems in the lower mantle vary significantly according to different works (Fig. 2.2). Most likely, the low melting temperature of the peridotite mantle in Nomura et al. (2014) can be explained by the effect of volatiles, since their starting composition contained at least 0.15 wt% H2O. The solidus of eclogite from Pradhan et al. (2015) corresponds better to low-pressure data (Yasuda et al., 1994) relative to that of eclogite from Andrault et al. (2014). Accordingly, estimations of the thermal structure of the Earth’s mantle reveal that melting may not occur in the solid silicate matrix without the addition of fusible components. The melting temperatures of silicate rocks are much higher than possible temperatures of the present-day mantle and exceed temperatures of the mantle adiabat by at least 600 K at depths of 800 km and 1200 K at 2700 km (Fig. 2.2). The solidus of peridotite intersects the mantle adiabat only at shallow levels beneath midocean ridges (Fig. 2.1). Abundant silicate melting could happen during the very early history of the Earth during and after core segregation or at local conditions such as the strong thermal gradients in mantle plumes. The potential temperatures of Archean basalt and komatiites were estimated as 1800e2000 K (Fig. 2.1) (e.g., Herzberg et al., 2010; Lee et al., 2010). Nevertheless, the possibility of temperature increase in plumes originating at the CMB during most of the geological

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history crossing the silicate solidi in the lower mantle is very low. Therefore, models of thermochemical plumes involving fusible components are usually considered (Lay et al., 1998; Lin and van Keken, 2005). Typically, the components of CeOeH fluids are assumed to be the most reliable candidates for the fusible component in thermochemical plumes (Belonoshko et al., 2015; Frost and McCammon, 2008).

4. REDOX STATE IN THE DEEP UPPER AND LOWER MANTLE The oxidation state may be as important as temperature for melting and freezing of the mantle rocks (Foley, 2011; Litasov, 2011; Rohrbach and Schmidt, 2011). Most estimates of fO2 for basaltic rocks and shallow mantle peridotites plot close to the FMQ (fayaliteemagnetiteequartz) oxygen buffer at fO2 versus temperature diagrams. However, data for garnet peridotites from kimberlites (e.g., Woodland and Koch, 2003; Yaxley et al., 2012) indicate a significant decrease of the fO2 with depth (Fig. 2.3). As is shown for the example of Siberian mantle peridotites, this pattern is complicated and fO2 trends scatter due to different metasomatic effects, which oxidize or reduce original mantle assemblages. The data for natural samples combined with experimental studies (Frost et al., 2004; Rohrbach et al., 2007) show that the Fe3þ contents of silicates increase with pressure even in equilibrium with Fe-metal. This increase is due to the stabilization of Fe3þ-bearing endmembers in solid solutions of high-pressure phases, particularly majorite garnet, pyroxene, and bridgmanite. At 8e10 GPa (250e300 km depth), the curve of the average fO2 from mantle peridotites would cross the stability line of FeeNi metal (Fig. 2.3). In the presence of a small amount of this alloy (0.1e0.5 wt% in the upper mantle and transition zone and w1.0 wt% in the lower mantle), the system will be buffered near the IW (ironewustite) equilibrium (or 1e2 log units below this buffer) and the average fO2 corresponds to the bold curve in Fig. 2.3. Under these fO2 conditions the oxidized forms of carbondthat is carbonates or carbonated meltsdcannot be in equilibrium with mantle rocks and are reduced to diamond, reduced fluids, or Fe-carbides (Rohrbach and Schmidt, 2011). It is difficult to estimate mantle fO2 below the intersection with the FeeNi alloy line, because it depends not only on the balance of Fe in silicates, but also on the heterogeneity of oxygen and carbon distribution in the bulk Earth. This heterogeneity can be primordial, caused by inefficient separation of the core-forming melt, by oxygen pump operation across the solidification front of bridgmanite in the magma ocean (e.g., Wade and Wood, 2005; Wood et al., 2013), or superimposeddcreated by subduction of the oxidized material. The oxidation state of subducting slabs should also evolve through time since subsurface conditions became more oxidized, especially after the great oxidation event in the early Proterozoic. However, evidence for more reduced compositions of ancient subduction-related rocks and basalts is still lacking (e.g., Frost and McCammon, 2008; Galliard et al., 2015). Oxidized subducting slab or mantle wedge materials can be entrained in the convective mixing, but can preserve heterogeneity for a long time due to the sluggish kinetics of exchange reactions, even in partially molten matrices. In this case, carbonate or carbonatite melt can survive in the deep mantle even when surrounded by reduced material of the ambient mantle. This possibility is clearly indicated by carbonate inclusions in superdeep diamonds originating from the uppermost lower mantle (e.g., Brenker et al., 2007; Kaminsky, 2012). The drastic change of the oxidation state across a redox front can trigger melting and this type of process would be very important in the deep mantle. Examples of such redox anomalies can be

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FIGURE 2.3 Oxygen fugacity ( fO2; relative to the IW buffer) for garnet peridotite along a cratonic geotherm as a function of pressure (black line is average for Kaapvaal). The data for Kaapvaal Craton (South Africa) after Woodland and Koch (2003) and for Siberian Craton after Yaxley et al. (2012). The FeeNi alloy line was calculated for peridotite after Frost and McCammon (2008). Positions of buffers FMQ (Fayalite þ O2 ¼ Magnetite þ Quartz) and EMOG/D (Enstatite þ Magnesite ¼ Olivine þ Graphite/Diamond þ O2) after Stagno and Frost (2010) and Stagno et al. (2013). Dashed lines show MCO3 2 molar concentration in the melt in equilibrium with diamond or graphite (Mda divalent cation) under reduced conditions below EMOG/D buffer (Stagno et al., 2013). The majority of cratonic peridotites would coexist with melts characterized by dilute carbonate contents (1%e10%). The bending of average fO2 line across the FeeNi alloy line indicates the change of redox regime from Fe2þ/Fe3þ control to Fe/Fe2þ control. This can be drastic (as shown here) or gradual (if transition to FeeNi metal-bearing mantle is scattered over depth interval).

associated with compositional contrast (slabemantle or plumeemantle interface), with a phase boundary (olivineewadsleyite) in the mantle, or with the depth of a prominent redox reaction, such as the FeeNi alloy boundary noted above.

5. MELTING AND PHASE RELATIONS IN THE CARBON-BEARING SYSTEMS The phase relations of carbon-bearing systems were considered in detail in Dasgupta (2013), Hammouda and Keshav (2015), Litasov (2011), and Litasov et al. (2013b). In this section, we focus on new data on phase relations in carbon-bearing systems, and briefly describe the data reported in the above reviews with additional analysis of Fe-metalecarbonateesilicate systems.

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5.1 SYSTEMS WITH CO2 The petrologically important peridotiteeCO2 and eclogiteeCO2 systems have been studied in simplified (CaOeMgOeAl2O3eSiO2eNa2OeCO2) (Keshav and Gudfinnsson, 2010; Litasov and Ohtani, 2009, 2010) and complex systems, replicating natural compositions (Brey et al., 2008, 2011; Dasgupta and Hirschmann, 2006; Dasgupta et al., 2004; Ghosh et al., 2009; Kiseeva et al., 2012, 2013; Shirasaka and Takahashi, 2003) (Fig. 2.4). Recently, data in simplified systems with carbonates have become available from DAC experiments at pressures exceeding 50e100 GPa (Maeda et al., 2017; Solopova et al., 2015; Thomson et al., 2014) (Fig. 2.5), but at present these data are still strongly controversial (see below). The understanding of melting in complex systems is hardly possible without careful consideration of the simplest and simplified compositions. First of all, the physical properties of CO2 in solid and fluid states change drastically with pressure. At about 35 GPa, the molecular phase IV is transformed to covalent CO2eV (Fig. 2.5), which can be harder than the closely related stishovite SiO2 structure. At high temperature phase V does not melt, but dissociates into C and O2 with a negative Clapeyron slope (Litasov et al., 2011a; Tschauner et al., 2001) (Fig. 2.5). This dissociation reaction plays a significant role for the high-pressure behavior of all carbonate and orthocarbonate compounds at pressures corresponding to the Earth’s mantle. Possible release of O2 by dissociation of solid CO2 or CO3 2 compounds in the melt may create strongly oxidized conditions in the deep lower mantle, changing local fO2 by up to 10 orders of magnitude (Litasov et al., 2011a). The melting curves of MgCO3 and CaCO3 are located at higher temperature relative to the CO2 dissociation line, but MgCO3 and CaCO3 melts may also dissociate to oxide, diamond, and O2 with further temperature increase for 200e400 K above the melting line (Solopova et al., 2015; Spivak et al., 2012) (Fig. 2.5). Similarly, carbonateesilicate melt is also transformed to produce diamond; for example, Seto et al. (2008) and Maeda et al. (2017) reported that the products of the MgCO3 þ SiO2 reaction are bridgmanite þ CO2 or bridgmanite þ C þ O2 at lower mantle pressures (Fig. 2.5). At pressures below 6 GPa the reaction of magnesite with quartz or coesite produces MgSiO3 and fluid CO2. However, in the pressure range from 8 to 25 GPa MgCO3 þ SiO2 melts, producing MgSiO3 þ melt or SiO2 þ melt assemblages (Kakizawa et al., 2015; Litasov, 2011) (Fig. 2.4). This melt should also be produced at higher pressures, but was not reported by Maeda et al. (2017). Melting of this assemblage is hard to detect in DAC experiments without the study of recovered samples. Therefore, additional study of this reaction is required. Nevertheless, Maeda et al. (2017) observed a transition in the reaction products from bridgmanite þ CO2 to bridgmanite þ C þ O2 near 70e75 GPa and 1800 K (Fig. 2.5). If the coldest subducting slabs penetrate into the lower mantle, they can produce diamond by magnesite decomposition reaction at this level. The CaCO3 þ SiO2 reaction was studied at pressures up to 30 GPa, and is located at slightly higher temperature than MgCO3 þ SiO2. It also produces CaSiO3 þ CO2 fluid at pressures below 6 GPa and CaSiO3 þ melt at pressures above 6 GPa (Litasov, 2011). Thermodynamic estimations indicate that at 25e30 GPa aragonite reacts with stishovite at much lower temperatures of 1700e1800 K (Litasov et al., 2017), producing CaSiO3 perovskite and CO2 and indicating decarbonation of aragonite in eclogitic rocks subducting to the lower mantle. However, this should be confirmed from experiments. Most solidi or decarbonation reaction lines change the PT slope at pressures above 6e10 GPa. This may be related to the transition from CO2-producing reactions to melting reactions and also to the significant change of partial molar volume of CO2 in carbonateesilicate melt from a rapid decrease to

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FIGURE 2.4 Carbonate stability in selected CO2-bearing systems at pressures to 30 GPa from MA experiments. Black dashed lines indicate magnesite decarbonation and melting (Irving and Wyllie, 1975; Katsura and Ito, 1990). The MgCO3 þ SiO2 ¼ MgSiO3 þ CO2 (or CO2-bearing melt) reaction is after Kakizawa et al. (2015)dbold orange line (light gray in print version), and after Litasov (2011)ddashed orange line (light gray in print version). Note at higher temperatures the reaction products can change to SiO2 þ melt. CaCO3 melting curve after Li et al. (2017). The CaCO3 þ SiO2 ¼ CaSiO3 þ melt reaction after Litasov (2011). Pink lines (dark gray in print version) P þ CO2dsolidus of model peridodite (Litasov and Ohtani, 2009) and E þ CO2declogite (Litasov and Ohtani, 2010) in the Na2OeCaOeMgOeAl2O3eSiO2 system þ 5 wt% CO2. Black line E þ CO2dsolidus of the MORB-eclogite (5 wt% CO2) after Dasgupta et al. (2004) and Shirasaka and Takahashi (2003) with decarbonation reaction bending at 3e4 GPa. Na-eclogite þ CO2dis solidus of Na-bearing eclogite (2.5 wt% CO2) and carbonatite (42 wt% CO2) after Thomson et al. (2016). PeliteeCO2 is the solidus of carbonated pelite (5 wt% CO2) (Grassi and Schmidt, 2011). Blue dotted lines (light gray in print version) show stability of Na-aragonite and solidus in the Na- and K-bearing carbonatite systems after Litasov et al. (2013c). Na2Ca, Na2Ca2, and (K, Na)2Ca4 are subsolidus double carbonates stable at different pressures, respectively. In addition, K2Mg(CO3)2 is stable slightly above the solidus temperature (not shown). The Dolomite (Dol) ¼ aragonite (Arag) þ magnesite (Mst) line after Litasov et al. (2011b). Eutectic temperature in the FeeFe3C system after Nakajima et al. (2009). Mantle adiabat, coldest subduction, and stagnant slab PT profiles are from Fig. 2.1. Three major fields can be recognized: (1) stability of magnesite, (2) stability of magnesite and CaCO3 with free SiO2 phases, and (3) stability of alkali-bearing carbonates. The boundary between (1) and (2) roughly coincides with the mantle geotherm and K-free carbonated peridotite and eclogite solidi, indicating that Mg- and Ca-carbonates can be transported to the deep mantle by subducting slabs, whereas alkali-bearing carbonates can survive only along the coldest subduction PT paths and are unlikely to be transported to the deep lower mantle. The boundary between (2) and (3) corresponds to the solidus of MORB-eclogite or pelite.

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FIGURE 2.5 Carbonate stability in simplified systems at lower mantle pressures based on DAC experiments. Data for MgCO3 melting and MgCO3-melt decomposition to MgO þ C þ O2 from Solopova et al. (2015). The data for MgCO3 þ MgSiO3 and MgCO3 þ CaCO3 melting are from Thomson et al. (2014). Phase boundaries in the CO2 system are after Litasov et al. (2011a). Reaction boundaries for the MgCO3 þ SiO2 system are after Kakizawa et al. (2015) at 0e25 GPa and after Maeda et al. (2017) at 25e140 GPa: (1) MgCO3 þ SiO2 ¼ MgSiO3 þ CO2 fluid; (2) majorite/stishovite þ melt; (3) bridgmanite þ CO2; (4) bridgmanite þ C þ O2; (5) MgSiO3postperovskite þ C þ O2. Reaction (6) shows assumed boundary between CO2 and C þ O2 coexisting with bridgmanite. Eutectic temperature in the FeeFe3C system is after Liu et al. (2016). Mantle adiabat, subduction PT profiles, and peridotite and eclogite solidi are from Figs. 2.1 and 2.2. Brg, bridgmanite; CS, CaCl2-type SiO2 phase; Fl, fluid CO2; Mst, magnesite; Ppv, MgSiO3-postperovskite; Sf, seifertite; St, stishovite.

relatively slow change (Ghosh et al., 2007; Sakamaki et al., 2011). It was found that Na2O and K2O play a key role in the melting of carbonate-bearing peridotite and eclogite in the complex systems: the addition of 0.1 wt% K2O reduces the solidus temperature by 500 K at 20 GPa in both systems. There are three major thermal regimes for carbonate stability under oxidized conditions (Fig. 2.4) marked by (1) magnesite-bearing silicate systems without a free silica phase; (2) magnesite/aragonite þ SiO2 assemblages (e.g., in eclogite); and (3) the stability of fusible alkali carbonates. In the first regime, the stability of magnesite is limited by decarbonation and melting reactions involving silicates, such as MgCO3 þ MgSiO3 ¼ Mg2SiO4 þ CO2. In the second regime, magnesite or aragonite stability is controlled by melting reactions with silica phasesdfor example, MgCO3 þ SiO2 ¼ MgSiO3 þ CO2. The third regime is limited by melting and reactions involving alkali carbonates. This emphasizes that the presence of alkali components has a major effect on carbonated peridotite and eclogite solidi. It was observed that near-solidus melts in the Na- and K-bearing carbonated peridotite and eclogite systems show strong enrichments in alkalis. However, the precise determination of the compositions of

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these melts and position of the solidus is difficult due to the low modal abundance of alkaline melt and alkali carbonates, and to the poor stability of quench products during polishing and electron probe microanalysis (Dasgupta and Hirschmann, 2006; Ghosh et al., 2009). The PT stability of alkali-bearing carbonates has been discussed in several papers (Grassi and Schmidt, 2011; Litasov, 2011; Litasov et al., 2013b,c; Thomson et al., 2016). The solidus temperatures of Na- and K-rich (Na2O ¼ 7 wt% and K2O ¼ 2 wt% and vice versa) MgeCa carbonatite with minor SiO2 and FeO increase from 1000 to 1400 K at 3e10 GPa, after which the solidus remains flat at 1400 K up to at least 21 GPa (Litasov et al., 2013c) (Fig. 2.4). The major carbonate phases in these systems are aragonite and magnesite. Aragonite contains significant amounts of Na2O (up to 7 wt%), K2O (up to 1 wt%), and MgO (up to 8 wt%) in the Naecarbonatite system. Na-bearing aragonite is the highest temperature alkali-bearing carbonate stable at pressures up to 20 GPa (Fig. 2.4), but at higher pressures it can be replaced by (K, Na)2Mg(CO3)2 carbonate. The second important double carbonate is (K,Na)2Ca4(CO3)5, which is stable under subsolidus conditions in the K-bearing system and slightly above the solidus in the Na-bearing system. In the Naecarbonatite system the subsolidus phases include Na2Mg2(CO3)3 at 3 GPa, Na2Ca2(CO3)3 (shortite composition) at 10 GPa, and Na2Ca(CO3)2 (nyerereite composition) at 16 GPa (Litasov et al., 2013c). NaeCa-carbonates control the solidus in the carbonated pelite (Grassi and Schmidt, 2011) and Na-rich eclogite (Thomson et al., 2016) systems. Both systems have a solidus bending near 13e16 GPa (Fig. 2.4), which is explained by the limited pressure stability of Na-bearing clinopyroxene. Imbalanced solubility of Naepyroxene in majorite garnet causes formation of fusible Na2Ca4(CO3)5 carbonate, resulting in solidus depression by 100e200 K. Grassi and Schmidt (2011) found Naecarbonate to be stable at much higher temperatures than Litasov et al. (2013c) or Thomson et al. (2016) (Fig. 2.4). Unfortunately, only one composition of Naecarbonate, intermediate between Naearagonite and Na2Ca4(CO3)5, was reported for the carbonated pelite system. Detailed phase diagrams for almost all binary and ternary alkali and alkali earth carbonate systems have been reported at 6 GPa. These diagrams, reviewed by Shatskiy et al. (2015a), help to constrain incipient melting conditions for carbonated rocks in the lithospheric mantle, but are not considered here in detail. We emphasize careful determination of composition and identification of crystal structures of subsolidus double carbonates with intermediate compositions (e.g., (K, Na)2(Mg, Fe)(CO3)2, (K, Na)2Ca3(CO3)4 and Na2Ca4(CO3)5) and close consistency of eutectic melts with those formed by partial melting of the peridotite, eclogite, and kimberlite-like systems (Gavryushkin et al., 2014, 2016a,b; Golubkova et al., 2015; Shatskiy et al., 2014, 2015a,b, 2016a,b). Important natural systems with carbonates studied at pressures up to 12 GPa include kimberlite and kimberlite-related rocks with variable amounts of volatiles including also H2O and halogens (Litasov et al., 2010; Sharygin et al., 2015; Shatskiy et al., 2017; Sokol et al., 2013, 2015; Ulmer and Sweeney, 2002). Shatskiy et al. (2017) show that parental kimberlite magmas were most likely close to carbonatite and could originate from a carbonated peridotite mantle source. They show that addition of 20 wt% CO2 to erupted kimberlite composition of the Udachnaya pipe (Siberia) equilibrates carbonatite melt with the four-phase garnet lherzolite assemblage at 6.5 GPa and 1700e1750 K. The carbonatitic nature of low-degree partial melts is a characteristic feature of all compositions studied to date in peridotiteeCO2, eclogiteeCO2, and kimberlite-like systems. There is a fundamental difference between these and systems with H2O. Low-SiO2 (<5 wt%) MgeCa-bearing alkali carbonatite melt has a very wide temperature stability field and is stable up to temperatures of the mantle adiabat at transition zone depths (Fig. 2.6). In the system with H2O, low-degree melting produces silicate-bearing melt (depending on total H2O in the system) and at temperatures of mantle adiabat this

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FIGURE 2.6 Solidi and melt compositions in the peridotiteeH2O (A) and peridotiteeCO2 (B) systems. Wet solidusdH2Osaturated solidus of peridotiteeH2O system (Iwamori, 2004). Hydrous minerals curve shows upper temperature limit of hydrous mineral stability (Litasov, 2011; Litasov and Ohtani, 2007). Dotted curves with numbers represent approximate H2O contents in the hydrous melt (A) and CO2/SiO2 contents in carbonated melt (B) based on the experimental data for hydrous (Inoue, 1994; Litasov and Ohtani, 2002; Stalder et al., 2001) and CO2bearing systems (Dasgupta et al., 2013b; Ghosh et al., 2014; Shatskiy et al., 2013), respectively. Carbonatite solidus after Litasov et al. (2013c). Mantle adiabat, cratonic, and selected subduction PT profiles and volatile-free solidus of peridotite are shown in Fig. 2.1.

hydrous melt contains 70e80 wt% of silicates. This explains the different mobility of hydrous silicate melt and carbonatite melt in the mantle, which is considered in detail below. The physical properties of carbonatite melt (low density, low viscosity, and low dihedral angles in silicate matrix) are poorly constrained at lower mantle depths. A recent theoretical study indicates stabilization of tetrahedrally coordinated carbon in carbonated MgSiO3 melt (Ghosh et al., 2017). The CO4 groups become dominant at pressures above 40e50 GPa, and partial reduction of carbonate to carbon was observed in calculation snapshots. These authors also argued for a minor effect of CO2 on the density of melt, which indicates much smaller gravitational instability of carbonated melt in the lower mantle. Indeed, transformation to tetrahedral symmetry of carbon was observed in solid carbonates at high pressure, confirmed in experiments and computations for CaCO3 near 100 GPa (Oganov et al., 2006; Ono et al., 2007), MgCO3 near 80 GPa (Isshiki et al., 2004; Maeda et al., 2017; Oganov et al., 2008), and FeCO3 near 60 GPa (Liu et al., 2015). Finally, it should be noticed that siderite undergoes high-spin to low-spin transition near 45 GPa (Lobanov et al., 2015a). The optical absorption data indicate that low-spin MgeFe-carbonate will be strongly enriched by Fe relative to coexisting MgeFe bridgmanite. This feature of FeeMg partitioning across the spin transition may be important for Fe-bearing carbonated melt in the lower mantle.

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The major element chemistry of carbonatite and carbonated silicate melts at 5e25 GPa is characterized by a different Mg/Ca-ratio, which is higher in the peridotite relative to eclogite systems. The lowest degree melting produces alkali-rich melts, which may contain up to 10e15 wt% K2O þ Na2O (Ghosh et al., 2009; Grassi and Schmidt, 2011; Litasov et al., 2013c; Thomson et al., 2016). Carbonates and carbonatite melt are important concentrators of rare earth elements (REE), high-fieldstrength elements (HFSE), and some other trace elements, especially in high-Ca varieties. Therefore, they can create recognizable signatures in trace element and isotopic compositions of coexisting minerals and affect the chemical characteristics of basaltic and alkaline magmas. We briefly review data on mineral/carbonated melt partitioning, summarized in Fig. 2.7, in this section. The direct comparison of element partitioning between minerals and carbonated and silicate melts is complicated, because temperatures for measurements at the same pressure differ appreciably. In general, partition coefficients between minerals and carbonatite melt (CL) (Dmineral/CL) do not change with pressure and temperature and remain 0.5e1.0 order of magnitude lower relative to Dmineral/SL (SLdsilicate melt). Some anomalies were detected as low DHFSE cpx=CL in melting experiments with the hydrous carbonated peridotite at 6e10 GPa (Girnis et al., 2013) and high DHFSE gt=CL and DPb gt=CL in experiments on carbonated pelite at 22 GPa (Grassi et al., 2012) (Fig. 2.7). For Ca-perovskite, D values are instead slightly higher for CL relative to SL (Fig. 2.7). The data for D between other minerals and CL are limited. Grassi et al. (2012) noticed that at 13 and 22 GPa K-hollandite has high D for Rb, Ba, and Pb, whereas Ca-Al phase at 22 GPa concentrates REE, Th, U, Pb, and Ti. D values for magnesite were obtained by Dasgupta et al. (2009) and Girnis et al. (2013), according to whom magnesite concentrates compatible elements such as Co and Ni, and has low D for most of other elements in the range of 103 to 102.

5.2 FEeCARBIDE AND CARBONATEeFE RELATIONS Feecarbide can be a minor phase throughout the deep upper and lower mantle as revealed by findings of different carbide and carbonitride inclusions in superdeep diamonds (Jacob et al., 2004; Kaminsky and Wirth, 2011, 2017; Smith et al., 2016) and significant abundance of Fe-carbides in iron meteorites (e.g., Rubin, 1997), representing the building blocks of small planetesimals. The ironecarbon system was studied in detail at pressures below 25 GPa and also in the DAC at pressures up to about 160 GPa (Fei and Brosh, 2014; Liu et al., 2016; Lord et al., 2009; Nakajima et al., 2009). It was shown that the FeeFe3C eutectic is shifted toward Fe-bearing compositions, and is replaced by an FeeFe7C3 eutectic near 150 GPa (Fig. 2.8). Liu et al. (2016) reported a very wide melting interval for Fe7C3 þ melt, which exceeds 1000 K at 100e150 GPa. The eutectic temperature of the FeeC system is located slightly above the mantle geotherm at lower mantle depths (Fig. 2.5), which indicates that Fe3C or Fe7C3 carbide can be stable throughout the lower mantle, but can be easily remobilized by heating or reactions in mantle plumes. Phase relations and redox reactions in the carbonateeFe systems have been studied at pressures of 6.0e7.5 GPa (Bataleva et al., 2015; Martirosyan et al., 2015a,b; Palyanov et al., 2013). The interaction below the solidus follows the schematic reaction: ðMg; CaÞCO3 þ Fe ¼ ðFe; Mg; CaÞO þ Fe3 C ðor Fe7 C3 Þ þ C

(2.1)

and above the solidus: ðMg; CaÞCO3 þ Fe ¼ ðFe; Mg; CaÞO þ Fe  CðmeltÞ þ ðMg; Ca; FeÞCO3 ðmeltÞ þ C.

(2.2)

FIGURE 2.7 Partition coefficients of trace elements between (A) clinopyroxene, (B) garnet, and (C) CaSiO3-perovskite and melts from the high-pressure experiments. Carbonatite melts (Dalou et al., 2009; Dasgupta et al., 2009; Grassi et al., 2012; Klemme et al., 1995); carbonated silicate melt (Keshav et al., 2005); hydrous carbonated silicate melt (Girnis et al., 2013); silicate melts (Corgne et al., 2005; Corgne and Wood, 2004; Klemme et al., 2002). Gray field represents other D values for volatile free and hydrous silicate melt (Bennett et al., 2004; Green et al., 2000; Pertermann and Hirschmann, 2002; Pertermann et al., 2004; Tuff and Gibson, 2007; Yaxley and Sobolev, 2007).

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FIGURE 2.8 Ironecarbon phase diagram at 20, 70, and 150 GPa calculated from experimental and thermodynamic data (Fei and Brosh, 2014; Litasov et al., 2013a; Liu et al., 2016; Nakajima et al., 2009). Note the shift of eutectic composition toward Fe and change of eutectics near 150 GPa from FeeFe3C to FeeFe7C3. Ironecarbon phase diagram at 20 (A), 70 (B), and 150 (C) GPa.

The reactions of iron with CaCO3 and dolomitic mixture (CaMg(CO3)2) produce nearly pure CaCO3 melt, magnesiowustite, and FeeC alloy at 1473e1673 Kdthat is, at temperatures significantly lower than the melting curve of CaCO3. The reaction kinetics for the CaCO3eFe system at 4e16 GPa reveal slow reduction of Ca-carbonate on geological timescales. In the MgCO3eFe system, the reaction rate is much faster. However, if (Mg, Ca)-carbonate does not melt in the slab during subduction, it can easily be transported to the CMB without significant reduction. To illustrate the rate of kinetic interaction, in a 50 m thick layer of CaCO3 carbonate (equivalent to CaCO3 in average 2e3 km thick metabasalt) subducting in direct contact with metallic Fe down to 660 km, only 0.5%e20% of the carbonate will be reduced by Fe depending on the slab temperature (Martirosyan et al., 2015b, 2016b). In real mantle environments most subducted carbonates will be isolated from direct contact with Feemetal even in the D00 layer near the CMB (see Sections 6.2 and 6.3). Depending on the balance between carbonate and Fe in the local area of the D00 layer or elsewhere at the slabemantle interface, the reduction of carbonate will produce diamond or Feecarbide. However, carbonate can even coexist with iron at high pressures if pressure shifts IW buffer equilibria toward oxidized conditions (higher f O2) faster than the carbonate (or CCO) buffers. Some preliminary data are available on the carbonateeFe interaction at lower mantle and CMB pressures: Martirosyan et al. (2016a) reported results on MgCO3eFe interactions at 70e149 GPa and

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1800e2600 K, showing formation of carbide and wustite. The reaction is complicated by the presence of two immiscible wustite phases: 6MgCO3 þ 19Fe ¼ 8FeO þ 10ðMg0:6 Fe0:4 ÞO þ Fe7 C3 þ 3C.

(2.3)

A similar reaction of dolomite (Ca0.6Mg0.4)CO3 with Fe was studied by Dorfman et al. (2015) at 49e110 GPa and 1800e2500 K. At the highest pressures they observed formation of (Fe, Mg)O, Fe-carbide, and diamond, and emphasized the formation of nearly pure CaCO3 solid or melt (as in experiments at 6 GPa). This indicates that a CaCO3 phase can be another host of carbon at the CMB and is an excellent candidate to trigger melting and initiate mantle plumes.

5.3 THE ROLE OF HYDROCARBONS AND H2O IN THE DEEP CARBON CYCLE The deep carbon cycle and material transport in the mantle are closely related to the cycling of hydrogen. Both elements, even in very small amounts, can govern major geodynamic processes and enhance melting, convection, and plume and plate motion. Indeed, they can be segregated and accumulated at specific places in the mantle, likely along the compositional, phase, or redox boundaries. In the presence of hydrogen, hydrocarbons in reduced systems are additional carbon-bearing compounds in the deep mantle. H2O- and CO2-(carbonate)-bearing melts coexist with silicates under oxidized conditions. In contrast to the CO2-bearing systems, the solidi of hydrous peridotite and eclogite strongly depend on the total H2O content in the system. As an example, the solidus of peridotite þ 0.1 wt% H2O (Fig. 2.9) reveals a complicated pattern according to the high solubility of water in the transition zone minerals. The low-pressure data (up to 5 GPa) on the silicate systems with H2O and CO2 are reviewed in detail in Litasov (2011) and Litasov et al. (2013b). The key observations include (1) the formation of carbonatite melts at the peridotite solidus in the amphiboleephlogopite stability field and (2) decarbonation reactions between carbonates and silicates resulting in the extraction of CO2. Green and Falloon (1998) argued that all types of mantle magmas, from basalts to kimberlites and carbonatites, may be produced by melting in the peridotiteeH2OeCO2 system at pressures of 2e7 GPa. Addition of H2O to the carbonated systems causes a reduction in the thermal stability limit of magnesite. At pressures above 6e7 GPa in peridotite with 3 wt% CO2 and 3 wt% H2O magnesite stability falls to the level of the solidi of the K2O-containing carbonatite systems (Fig. 2.9), whereas the solidus is determined by the stability of dense hydrous magnesium silicates (DHMSs). In the eclogite system, the solidus is located below the lowest temperatures in the experiments, 1273 K at 10e30 GPa, and magnesite is stable up to 1400 K in the same pressure interval. The reason for such a drastic reduction in magnesite stability is poorly clarified thermodynamically. Both H2O and CO2 reduce mutual activity in the system and reduce the thermal stability of carbonates and DHMS (Foley et al., 2009; Foley and Pinter, 2017; Litasov et al., 2011b). Low fraction melts in these systems depend on the total H2O and CO2 contents. If the H2O content exceeds the solubility in the nominally anhydrous minerals at 10e30 GPa, incipient melts contain 15e25 wt% SiO2 (or up to 40e50 wt% SiO2 in the volatile-free residue) since carbonates are more stable than DHMS. However, at lower pressures (below 6 GPa) the incipient melt fraction can be carbonatitic (Foley et al., 2009) due to enhanced amphibole and phlogopite stability.

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FIGURE 2.9 (A) Approximate variations of fO2 relative to the ironewustite (IW) buffer (black line) and Feemetal fraction (gray line) through the mantle section to 700 km depth after Rohrbach and Schmidt (2011). (B) A comparison of solidi in different peridotite systems. FeeP and MoePdsolidi for peridotite with reduced CeOeH fluid at fO2 control by IW (more precisely by olivineeironeenstatite buffer for pyrolite composition) and MoeMoO2 (MMO) buffers, respectively (Litasov et al., 2014). P þ CO2dperidotite þ 5 wt% CO2 (Dasgupta and Hirschmann, 2006; Ghosh et al., 2009). Magnesite stability in P þ H2O þ CO2 coincides with this line. P þ H2Odwater-saturated solidus of peridotite (at pressures above 10 GPa defined by the stability of dense hydrous magnesium silicate phases; Litasov and Ohtani, 2003, 2007). P þ H2O þ CO2dperidotite þ 3 wt% H2O þ 3 wt% CO2 (Litasov, 2011; Litasov et al., 2011b). The solidi for all related eclogite systems are generally slightly lower than those for peridotite and are not shown for simplification. P þ 0.1 wt% H2Odis an example of complex hydrous solidus of peridotite defined by solubility of hydrogen in olivine polymorphs (Litasov, 2011). Volatile-free peridotite solidus, mantle adiabat, oceanic and cratonic geotherms, and selected subduction PT profiles for stagnant and downgoing slabs are from Fig. 2.1 and eutectics in the FeeC system are from Fig. 2.4. Brg, bridgmanite; Fpv, ferropericlase; Ol, olivine; Rw, ringwoodite; Wd, wadsleyite.

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Under reduced conditions corresponding to the deep mantle (at f O2 below carbonate stability), the CeOeH fluid can be represented by a mixture of H2O and hydrocarbons. The approximate composition of CeOeH fluid at various f O2 can be calculated using equations of state of components in the gas/fluid phase. The pressure dependence of the fluid composition along mantle geotherm and f O2 calculated for average Kaapvaal peridotite (black line in Fig. 2.3) using equations of state from Zhang and Duan (2009) indicates a gradual change from CO2eH2O-bearing fluid to H2OeCH4 composition (Fig. 2.10). H2O becomes predominant in the fluid with increasing pressure and under reduced conditions close to the IW buffer (Fig. 2.10B), while methane or other hydrocarbons remain as subordinate components. The relative stability of different hydrocarbon compounds and their relation to H2O have not been studied yet at very high pressures. The available experiments on high-pressure behavior of methane in DAC to 100 GPa reveal intervals of (1) melting up to about 40 GPa and 1000 K, (2) partial dissociation with formation of diamond and H2 at pressures up to 100 GPa and 1500 K (this observation is complicated by continuous migration of fluid around the laser-heating spot in DAC), and (3) formation of heavy alkanes and unsaturated hydrocarbons at higher temperatures (Kolesnikov et al., 2009; Lobanov et al., 2013). Computation of H2OeCO2 or CeOeH fluid mixtures coexisting with metallic iron at CMB conditions indicates formation of CeC and CeH bonds (Belonoshko et al., 2015), which is consistent with the experimental results on methane.

FIGURE 2.10 Speciation of CeOeH fluid in equilibrium with graphite/diamond calculated (A) as a function of pressure along the mantle adiabat with potential temperature of 1590 K and fO2 defined by black line in Fig. 2.3, and (B) as a function of fO2 relative to the IW buffer in the lower mantle at pressures of 25 and 50 GPa and 1873 K (Frost and McCammon, 2008). Other calculated species in (A) are H2 and CO, but their fraction is close to zero. CRM and HRM define conditions for carbonated redox melting and HRM at the slopes of “water maximum,” respectively (Foley, 2011). The calculations were performed using an equation of state model by Zhang and Duan (2009) in (A) and Belonoshko and Saxena (1992) in (B).

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It is important to emphasize that the solubility of silicates in reduced CH4eH2O fluid is very limited according to experiments at 0.9e3.6 GPa (Taylor and Green, 1988). However, at higher pressures this is not the case: in experiments on the peridotite and eclogite systems with reduced CeOeH fluid buffered at MMO (MoeMoO2) and IW at 10e16 GPa and 1473e1873 K, we observed significant precipitation of silicate micrograins in a diamond trap, which formed after breakdown of the stearic acid fluid source. This was interpreted as silicates dissolved in reduced CH4eH2O fluid. The exact proportion of silicates in the fluid was not determined, but estimated to be at least 5e10 vol%. In experiments at 3e6 GPa we did not observe silicate inclusions in the carbon aggregates (Litasov et al., 2014). The solidus temperature in these systems (for both buffers) was substantially higher than the solidi in systems with H2O and CO2 (Fig. 2.10), but still 400e500 K below the melting curve of volatile-free peridotite at 16 GPa. The eclogite solidi are 50e100 K below the peridotite solidus. The first melt generated near the solidus has 40e50 wt% SiO2 in the dry residue recalculated to 100% (Jakobsson and Holloway, 2008; Litasov et al., 2014). The data on melting in volatile-bearing systems constrain the conditions for redox melting in the Earth’s mantle. Foley (2011) distinguished two types of redox melting: hydrous redox melting (HRM) and carbonate redox melting (CRM). The HRM is a classic mechanism and corresponds to the transition from hydrocarbon-bearing to hydrous fluids (at f O2 ¼ IW ¼ 0  IW þ 1.5di.e., 0e1.5 log units above IW buffer) (Fig. 2.10). This type of melting is distinctive for ancient subduction, lithosphere erosion, and metasomatism. The CRM corresponds to transition from hydrous to CO2-rich fluids (at f O2 ¼ IW þ 4.5  IW þ 5.5 or FMQ  1.5  FMQ  0.5) (Fig. 2.10) and is typical for Phanerozoic rifting and subduction. Due to possible expansion of the fO2 interval for H2O maximum (Fig. 2.10) the role of redox melting in the deep mantle would be diminished.

6. DEEP CARBON CYCLE, MELTING, AND MATERIAL TRANSPORT IN THE EARTH’S MANTLE 6.1 CARBON SOURCES AND CYCLING IN THE EARLY EARTH The origin of volatile elements including carbon on Earth is debated. The major problems are related to: (1) the basic model of the Earth composition: carbonaceous versus enstatite chondrite; (2) the nature of late delivery of 10% or even 20% of bulk mass and possible loss of volatiles during Moonforming giant impact; and (3) efficiency and timing of metallic melt and the partial segregation of light elements into the core. Furthermore, in the carbonaceous chondrite model, it is unclear how abundant volatiles could be accreted. Their abundance may not correspond to average carbonaceous chondrite (3.5 wt% C and 2.1 wt% Hde.g., Lodders, 2003) since volatile accretion at the distance of 1 a.u. from the Sun is unlikely due to very low condensation temperatures. This means that Earth-forming carbonaceous chondrites could have been essentially “dry.” Thus, the estimations of the bulk Earth and BSE (bulk silicate Earth) composition are highly uncertain with regard to the light elements. Dasgupta (2013) examined carbon cycling in the magma ocean and during core formation in detail and the reader is referred to this comprehensive review. Here we outline a brief sketch of the suggested models. Carbon is strongly partitioned into the metallic phase in metalesilicate liquid (ML/SL) equilibria, depending on the PeTeXefO2 conditions DC ML=SL ¼ 500  5500 (Dasgupta et al., 2013a). The partition coefficient decreases with increasing silicate melt NBO/T, fO2, temperature, and decreasing pressure. Rough estimations of carbon partitioning in the magma ocean indicate that

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whatever the present-day BSE content of carbon, it exceeds the equilibrium concentration in the FeeNi metal during core formation and, thus, carbon concentrations in the present-day mantle and core are not equilibrated. For example, if the primordial bulk Earth contained 1 wt% carbon, the core would contain 3 wt%, whereas BSE will contain 6e60 ppm. This may correspond to the depleted mantle source and is much lower than reliable estimations for average BSE, which lie in the range 120e530 ppm C (Marty, 2012; McDonough and Sun, 1995). The estimations in Dasgupta (2013) indicate that only 1e20 wt% of core melt was equilibrated with silicate magma during core formation. In these scenarios, the calculated carbon content in BSE and core will be more reliable. Another possibility for carbon storage in the mantle after core formation is inefficient metal extraction, whereby some part of the metal is retained in the mantle after solidification of the magma ocean. Again, in this scenario it is difficult to estimate the amount of metal and carbon stored in the mantle. Without a doubt, a significant amount of carbon was accreted to the Earth’s mantle, most likely as a late veneer or as residual metal that was not effectively extracted to the core. This is evident from geochemical characteristics and fluid inclusions in volcanic and sedimentary rocks from Archean time.

6.2 SUBDUCTION, MELTING, KINETICS, AND PROBABLE STABILITY OF CARBONATES AT THE COREeMANTLE BOUNDARY The establishment of plate tectonics 3.5e4.0 Ga ago (e.g., Hopkins et al., 2010; Komiya et al., 1999) created the possibility for ingassing of carbon and hydrogen into the deep mantle via subduction and delamination. Progressive cooling of the Earth may have transformed the volatile cycles from dominant outgassing when carbon and hydrogen preferentially accumulated in the subsurface reservoirs, to a regime of prevalent ingassing when more carbon and hydrogen submerged to depths (Maruyama et al., 2014; Ru¨pke et al., 2006). Indeed, the cycles of carbon and hydrogen degassing are connected with fluctuations of the thermal regime of the mantle. During supercontinent formation and lower magmatic activity outbound fluxes diminish, whereas during breakup of supercontinents the outgassing rate increases. The major issue affecting carbon and hydrogen subduction is their devolatilization at island arc depths and the question of the existence of very cold slabs. Carbon and hydrogen subduction modeled separately can be estimated with reasonable uncertainty, but behavior in the mixed system is still poorly constrained. The estimations of carbon subduction beyond mantle wedge depths range between 20% and 80% of total subducted carbon (Johnston et al., 2011; Sano and Williams, 1996). Ancient Archean subduction allowed essentially volatile-free rocks to be subducted beyond mantle wedge depths, since the subduction PT profiles recorded from ophiolite and UHP complexes corresponded to the modern oceanic geotherms (Fig. 2.1). Proterozoic PT profiles are intermediate and most Phanerozoic UHP complexes and ophiolites are consistent with the PT profiles of modern hot slabs (Fig. 2.1) (Dasgupta, 2013; Litasov et al., 2013b; Maruyama and Liou, 2005). This means that starting from Proterozoic time (ca. 1.5 Ga), deep subduction of carbonates could be initiated, but even for modern subduction roughly half of the plates can transport a significant amount of carbonates beyond the mantle wedge zone (Dasgupta, 2013; Litasov et al., 2013b). If we take hydrogen into account, this amount will be smaller. Some scenarios of subduction (e.g., ophicarbonates over thick peridotite slab) have allowed carbonate penetration into the deep mantle since Archean times. Despite huge uncertainties and approximations, it is clear that up to 1e2 wt% CO2 in the upper kilometers of the slabs can be subducted to the transition zone and deeper into the lower mantle. Thus, a potentially

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significant amount of carbon can be delivered to the CMB if subducted slabs can penetrate to those depths. It should not be forgotten that some amount of reduced organic carbon as graphite can be subducted beyond arc depths without loss, but this carbon is just added to the primordial reduced reservoir and involved in melting only by the addition of oxidizing agents. The major issue is to deliver oxidized carbon as carbonates to deep levels. The carbonate submerged in the slab to the transition zone and lower mantle must overcome other dangers to survive to the CMB. A second important barrier for carbon and hydrogen subduction is the transition zone and its boundaries. The transition zone has significant potential to store hydrogen in wadsleyite and ringwoodite, thus creating a dehydration zone above the 410 and below the 660 km discontinuities (Bercovici and Karato, 2003; Litasov and Ohtani, 2007; Schmandt et al., 2014). Some carbonates in subducting slabs can be involved in redox melting with the appearance of CH-bearing liquids, forming a reservoir similar to that described in Section 6.4. Most melting episodes during slab subduction are limited by the disappearance of alkali-bearing carbonates, since (Mg, Ca)-carbonates are stable in mantle lithologies along the subduction PT profiles. Redox freezing zones may be limited to a narrow area along the slabemantle interface (Fig. 2.11). As was discussed above, significant reduction of slab carbonates by surrounding reduced mantle may not be possible due to sluggish kinetics of material exchange between slab and mantle. Self-reduction may be the most efficient process to destroy carbonates in the subducted slab because major minerals are in direct contact with carbonates. Partial transition of Fe2þ to Fe3þ during subduction of garnet and pyroxene causes reduction of coexisting carbonates coupled with the formation of diamond. If MORB eclogite contains 1 wt% CO2 (0.83 mol% C) and 5 mol% Fe2þ and about 20% of this Fe is oxidized to Fe3þ, then 30% (Ca, Mg)-carbonate will be reduced to diamond during subduction to 660 km depth. If MORB eclogite is subducted into lower mantle and bridgmanite and Al-phases contain 50% Fe3þ, then about 75% of carbonates will be reduced to diamond in the lower mantle. This approximation indicates that even in direct contact with a reducing agent, a significant amount of carbonates can be subducted to the CMB. Self-reduction of carbonates in the subducted basalt was shown in experiments by Kiseeva et al. (2013). A comparison of solidi for Na- and K-bearing carbonatite with the subduction PT profiles indicates that if the oxygen fugacity of the subducting slab remains sufficiently high to stabilize carbonate, melting of alkali carbonate may occur at a range of depths depending on the PT profile of subduction. In some cases, they may survive to lower mantle depths, but in most cases, a likely region for melting of subducted alkali-bearing carbonated eclogite is the transition zone as the solidi of carbonate-bearing systems become flat above w10 GPa (Fig. 2.4).

6.3 THE INEVITABLE FATE OF CARBONATES AT THE COREeMANTLE BOUNDARY There are several possible scenarios for the interaction of subducted material with rocks in the D00 layer and the CMB (Fig. 2.11). In the first model, subducted rocks will never appear in contact with the CMB and metallic iron, but interact only with uppermost layer of dense D00 rocks before being incorporated in mantle convection or plumes. In this case, subducted carbonates will be mostly transformed to diamond without a significant role for reduced components (Maeda et al., 2017). Depending on the rheological properties of the D00 layer, subducted materials may not interact with the CMB due to formation of a dense thermochemical pile (Li et al., 2014; Tackley, 2012). A second scenario invokes active reaction of subduction materials with the D00 and even with metal from the core. In this case, we

6. DEEP CARBON CYCLE, MELTING, AND MATERIAL TRANSPORT

65

FIGURE 2.11 Schematic model for subduction of carbonates down to the transition zone (stagnant slab) and coreemantle boundary combined with the major redox processes associated with burial and remobilization of carbon-bearing compounds. We emphasize five important processes related to carbonate or carbon-bearing compound activity in the post-Archean deep mantle. (1) MORB generation induced by deep melting of carbonates (Stagno et al., 2013). The model implies redox melting by oxidation of diamond and graphite in upwelling mantle at depths of 150 km with the formation of incipient carbonated melts. This melt will enhance large-scale melting at the shallow depths. Circle (1) shows field of the redox reaction 2Fe2O3 þ C ¼ 4FeO þ CO2. (2) Carbonate subduction inside crustal layer (sediments þ basalts) of the slab showing partial reduction at the slabemantle interface, formation of the local redox freezing zones, and self-reduction by Fe-disproportionation in slab silicates. (3) Mantle transition zone (MTZ) plume generation showing carbonatite melt diapir formed by melting of carbonates in the stagnant slab (see also Fig. 2.12). Carbonatite diapirs form an oxidized channel that is reused until carbonate supply by slab is shut down. (4) Carbonate reduction at the coreemantle boundary (CMB) and generation of the CMB plume. In hydrogen-free conditions carbonates are reduced to diamond in contact with reduced lithologies of the D00 layer and to carbide if they encounter metallic Fe. In this case, remobilization of carbon from the CMB is difficult. In hydrous conditions (H2O þ CO2) carbon and hydrogen react with metallic iron or reduced rocks to form hydrocarbons (shown as CeC and CeH bonds), which can enhance formation of mantle plumes (Belonoshko et al., 2015; Martirosyan et al., 2016a). (5) Carbonatite or hydrocarbon-bearing hydrous melts approach the lithosphereeasthenosphere boundary (LAB), where they can solidify to form the source for later magmatism or, in case of high enough capacity, immediately cause lithosphere erosion, deformation, and metasomatism with the formation of carbonatite, kimberlite, or other alkaline magmas. The examples for kimberlite magmas are shown in Pokhilenko et al. (2015), Sharygin et al. (2015), and Shatskiy et al. (2017). Brg, bridgmanite; CC, continental crust; Dia, diamond; Gr, graphite; OC, oceanic crust; Ppv, postperovskite (Mg, Fe)SiO3; SCLM, subcontinental lithospheric mantle.

may expect formation of diamond in the top portions of the slab/D00 contacts and Fe-carbide if carbonate or diamond occurs in direct contact with metallic Fe (Martirosyan et al., 2016a) (Fig. 2.11). The reactions of Fe with mixed carbon and hydrogen have been modeled only theoretically. The study by Belonoshko et al. (2015) shows formation of hydrocarbons by reaction of Fe with CeOeH or

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CHAPTER 2 CARBON-BEARING MAGMAS

H2OeCO2 fluid. These hydrocarbons can serve as fusible components for the formation and movement of mantle plumes. It is important to note that Fe is preferentially bonded to carbon and oxygen so that these elements may be consumed by the core. In contrast, hydrogen does not form bonds with Fe and is mostly involved in plume recycling. This observation has yet to be tested experimentally.

6.4 THE MODEL OF CARBONATITE OR HYDROCARBON-BEARING DIAPIRS Depending on the balance between Fe and C we will have three possible assemblages in the deep mantle: Feecarbide þ Fe, Feecarbide (balanced Fe/C ratio), and Feecarbide þ diamond. In reality, based on evidence from inclusions in superdeep diamonds, we have strong local redox disequilibria, which allow the stability of all of carbonate, diamond, and Feecarbide along with some amount of metallic Fe. This is readily explained by sluggish solid-state kinetics and even melting reactions in the mantle, which cannot obliterate compositional and redox heterogeneities over geological time scales. In Litasov et al. (2013b,c) we developed a simple numerical model for a carbonatite diapir generated in the transition zone by melting of a stagnant or downgoing subducting slab. This model can be tested further by more robust computational approaches. We reproduce this model here with updated input and variable parameters. First of all, we recall that hydrous melt has very low mobility in the mantle due to the high solubility of silicates in the hydrous fluid and significant solubility of water in the nominally anhydrous mantle minerals (e.g., Litasov, 2011; Litasov and Ohtani, 2007). In contrast, carbonatite melt with minimal solubility of silicates (<5 wt% SiO2) is stable in a wide temperature range including temperatures of the mantle adiabat (Fig. 2.6). We assume the following input parameters to model the stagnant slab, 70 km thick and 700 km long with simplified geometry, dip angle 45 , and equivalent thickness before and after bending above 660 km (slab top at 590 km). The CO2 content of the upper 500 m of the oceanic crust varies from 1 to 3 wt% (Staudigel, 2014). A total of 20%e80% of these carbonates can be transported to depths beyond 150 km, so we assumed the CO2 content in the top 500 m portion of the slab to be 0.2e2.4 wt% and ignore all CO2 possibly transported in the deeper part of the slab. In the slab stagnating in the transition zone melting of carbonates produces carbonatite melt with 40 wt% CO2 and a volume fraction (4) of 0.05e0.10, which is significantly higher than the limit of carbonatite melt interconnectivity in solid silicate matrix (Hunter and McKenzie, 1989; Minarik and Watson, 1995). Extreme wetting properties of carbonatite melt in a peridotite matrix are confirmed by electrical conductivity measurements at 1.5e3.0 GPa (Yoshino et al., 2010, 2012). The density of CaeMg carbonatite melt (rCL) (with SiO2 < 5 wt%) at the transition zone conditions is 2900e3000 kg/m3 based on the partial molar volume of CO2 and the thermal expansion of carbonated melts (Ghosh et al., 2007; Liu and Lange, 2003; Sakamaki et al., 2011). This density is lower than PREM (Dziewonski and Anderson, 1981) at transition zone depths (3600e4000 kg/m3). Melting of carbonates initiates buoyancy-driven porous flow and the migration of carbonatite melt toward the slabemantle interface (Fig. 2.12). In the hydraulically limited regime of compaction-driven melt flow through a creeping matrix, when the rate of fluid flow is controlled by the matrix permeability, the melt ascent velocity (yp) can be expressed as (Connolly et al., 2009): yp ¼

kDrg ; fhCL

(2.4)

where Dr is the density contrast between melt and surrounding silicate matrix (w700 kg/m3), g is the gravitational acceleration (10.0 m/s2 in the transition zone), hCL is the viscosity of carbonatite melt

6. DEEP CARBON CYCLE, MELTING, AND MATERIAL TRANSPORT

67

FIGURE 2.12 Formation of a carbonatite diapir by melting of carbonates in the uppermost 500 m layer of a model subducting slab stagnating in the transition zone. Three mechanisms of melt transport are involved (Litasov et al., 2013b): (A) buoyancy-driven porous flow within the partially molten slab layer; (B) surface tension-driven infiltration of “dry” overlying mantle and accumulation of melt at the slabemantle interface; (C) buoyancy-driven diapir ascent accompanied by pressure-solution creep of the infiltrated layer.

(0.006e0.009 Pa s) (Dobson et al., 1996; Genge et al., 1995; Kono et al., 2014), and k is the permeability, which depends on the melt fraction (f), grain size (d), and geometrical factor (C): fn d 2 : (2.5) C The values of n ¼ 3 and C ¼ 10e270 were suggested from experimental data at low pressures (McKenzie, 1989; Wark and Watson, 1998; Zhu et al., 2011). The porous flow regimes (including the rheologically limited regime) provide sufficiently fast velocities (>100 m/year) to allow efficient partial melt extraction from the uppermost layer of the subducting slab (Fig. 2.13A). The corresponding CO2 flux toward the slabemantle interface will be: k¼

fp ¼

wt% LWfyp rCL CCO 2 %; 100

(2.6)

where L ¼ 7 km is the width of partially molten zone above the solidus temperature (Currie et al., 2004) (Fig. 2.12), W ¼ 700 km is the average length of subduction zone segment (e.g., van Keken

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CHAPTER 2 CARBON-BEARING MAGMAS

FIGURE 2.13 Carbonatite melt velocity (A) and fluxes of CO2 (B). “Subduction” indicates CO2 delivery by slab with a velocity of 1e11 cm/year (average rate of 6 cm is marked as a dashed line in [A]). “Porous flow” indicates rate of melt migration toward the slabemantle interface. “Infiltration” provides the rate and flux of CO2 loss from the slabemantle interface to the surrounding mantle. “Diapir” shows carbonatite melt velocity and flux by diapir with a 1 km radius. (C) Carbon loss during diapir ascent due to carbonate reduction to diamond and carbide. Black and gray lines indicate the amount of carbon lost from the diapir and crystallized as reduced diamond or carbide for two Fe concentrations of 0.1 and 0.2 wt%. Dotted line shows approximate loss of CO2 by silicate crystallization according to the experimental data shown in Fig. 2.6B. wt% ¼ 40 wt% is the CO content in the carbonatite melt (Fig. 2.6). With C values at the et al., 2011), CCO 2 2 higher bound of 200e270 the applied values yield log( fp) ¼ 13.6e15.1 kg CO2/year (Fig. 2.13B). The experimental data indicate rapid impregnation and dispersion of low-viscosity melt in the silicate matrix by wetting the grain boundaries. Carbonatite melt has very low dihedral angles (q < 60 ) in the silicate matrix (Hunter and McKenzie, 1989; Minarik and Watson, 1995; Yoshino et al., 2010). At first glance carbonatite melt from the subducted slab should impregnate surrounding mantle and be consumed by redox freezing to create a metasomatic aureole (Fig. 2.11). However, the rate of the melt percolation will diminish rapidly if the infiltration distance exceeds the kilometer scale because of the increase of diffusion distance and the blurring of interfacial energy difference at the interface with the nonwetted rock (Fig. 2.12). To balance the flux of the equilibrium melt into the nonporous solid aggregate, an equal diffusive counterflux of solid through the melt must exist (Hammouda and Laporte, 2000). The rate of melt infiltration can be expressed from the characteristic diffusion distance (x) (Crank, 1975) as: rffiffiffiffiffi dx D yi ¼ ¼ : (2.7) dt t The silicate diffusivity (D) in carbonatite melt was estimated as 2  109 m2/s at the PT conditions of the transition zone (Shatskiy et al., 2013). The CO2 flux by carbonatite melt infiltration into surrounding mantle can be expressed as:

fi ¼

wt% yi Wxfe rCL CCO 2 %; 100

(2.8)

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69

where x ¼ 0.06 m/year is the average subduction rate (van Keken et al., 2011), and fe ¼ 0.2 is the equilibrium (maximum) carbonatite melt fraction in the silicate matrix near the interface with the melt pool (Hammouda and Laporte, 2000). From this equation the maximum CO2 flux into surrounding mantle is log( fi) ¼ 6.4 kg CO2/year during the first year and diminishes with time (Fig. 2.13B). This is 2e3 orders of magnitude lower than the annual CO2 supply from the model subducting slab, which is log( fs) ¼ 8.2e9.3 kg CO2/year at CO2 ¼ 0.2e2.4 wt% in the top 500 m of the 700 km slab. Note that all fluxes can be easily recalculated to the total subduction flux using, for example, the total length of modern subduction zones w38,500 km (van Keken et al., 2011). The possibility of the buoyant ascent of the melt diapir is strongly related to the viscosity of the country rocks (Anderson, 1981). For a spherical diapir moving under gravity in a viscous medium, the ascent velocity (yD) will be: yD ¼

2gr 2 Dr ; 9hdm

(2.9)

where r is the radius of the sphere and hdm is the viscosity of the surrounding “dry” mantle. If we assume a mantle viscosity of 1021 Pa s (e.g., Stacey and Davis, 2008) the ascent rate will be negligibly slow. However, a low-viscosity silicate layer impregnated by carbonatite melt must surround the melt diapir (Fig. 2.12). This layer will have low viscosity due to the pressure-solution creep deformation regime for a rock wetted by a solvent, which is different from the dislocation creep regime limited by the grain-boundary diffusion in a melt-free silicate mantle. Indeed, the pressure-solution creep is much faster than other creep regimes. The viscosity of wetted mantle rock is expressed as: hwm ¼

d 3 RT ; ADC0 Mw

(2.10)

where A is the constant of about 10 for equiaxial polycrystals (Kruzhanov and Sto¨ckhert, 1998), C0 is the silicate solubility in carbonatite melt (we applied a silicate fraction of 0.1e0.2 at the temperature of the mantle adiabat, 1873 K), M is the molar volume of silicate (3.91e3.97  105 m3/mol for wadsleyite in the transition zone; Katsura et al., 2009), and w is the effective grain boundary width, which is of the order of 1e10 nm (Dysthe et al., 2003). The estimated viscosity is w1017 Pa s, which corresponds to a diapir ascent rate of 0.2e0.4 m/year for r ¼ 1 km (Fig. 2.13A). This viscosity may correspond to a maximum, and consequently, the estimated rate is the minimum bound. The maximum diapir CO2 flux estimated from the slab and diapir geometry and diapir ascent rate is log( fd) ¼ 11.4e11.7 kg CO2/year, which is 2e3 orders of magnitude higher than the subduction CO2 flux. Therefore, some time is needed for accumulation of CO2 at the slabemantle interface between the last diapir departure and the next diapir formation. The exact time for replenishment of the CO2 reservoir depends on the efficiency of carbonatite melt withdrawal by the diapir and needs more detailed calculations. The important issue is the reaction of the carbonatite diapir with the overlying, presumably reduced, mantle. Rohrbach and Schmidt (2011) suggested a priori that carbonatite melt will be involved in redox freezing and will be immobilized by transforming to diamond and other species. Simple mass-balance calculations indicate that carbonatite diapirs can oxidize all reduced components on the way to the LAB (Fig. 2.11). A first carbonatite diapir with r ¼ 1 km formed at the slabemantle interface in the transition zone will be reduced by 40% during migration from 590 to 200 km depth if the Fe0 content in the transition zone and upper mantle is 0.1 wt% and by 80% in the case of 0.2 wt%

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Fe0 (Fig. 2.13C). Every later diapir will use the oxidized channel and will not be reduced significantly. The minimum radius of diapirs, which approach 200 km depth at Fe0 ¼ 0.1 wt%, is 400 m. In these calculations, we also accounted for the solubility of elemental carbon in carbonatite melt and approximated reduction of the CO2 content due to increasing silicate solubility for 40 to 33 wt%. Accounting for the role of H2O in the carbonatite diapir movement is more difficult. If the transition zone and deep upper mantle contain 0.1 wt% H2O (which would be the maximum concentration) then nearly all this water (according to partitioning into carbonatite melt, e.g., Shatskiy et al., 2009) should be drained by the ascending carbonatite diapir. For simplicity, we assumed a homogeneous concentration of 0.1 wt% H2O through the transition zone and asthenospheric upper mantle and 70% reduction of this amount by carbonatite melt. Then, for a diapir with r ¼ 1 km the maximum H2O concentration can be calculated as: CL CH ¼ 100 2O

MHCL2 O MHCL2 O þ Md

;

(2.11)

where MHCL2 O ¼

m pr 2 hrm CH 2O 100

(2.12)

and Md is the mass of the diapir in kg. Applying the height of the mantle column h ¼ 590  200 ¼ 390 km, the density of the mantle rm ¼ 4000  3400 kg/m3 (with decreasing m ¼ 0.07 wt%, one can obtain C CL ¼ 19.4 wt% in depth), and concentration of H2O in the mantle CH H2 O 2O the diapir ascending from the transition zone to the lithosphereeasthenosphere boundary. No or minor hydration is expected for every subsequent diapir traveling though the dried conduit. Hydration of carbonatite melt will increase the rate of segregation and the diapir ascent rate by up to an order of magnitude due to the increased diffusivity and solubility of silicate components in the H2O-bearing melt. Increasing silicate solubility in carbonatite melt can be critical for its ability to migrate. However, in the model we applied maximum H2O concentrations. Based on our unpublished data on hydrous carbonatite melts at 3e6 GPa we do not expect a significant increase of the silicate content in the carbonatite melt until H2O concentrations exceed 10 wt% (which is more suitable for real mantle conditions). This, however, should additionally be tested with experiments. The interaction of the carbonatite melt diapir with the LAB may be complicated. However, without a doubt this boundary represents a strong barrier to the movement of the diapir toward the surface. The accumulation of carbonated melts at the LAB can create variable source regions of kimberlite, carbonatite, and other alkaline rocks, which return CO2 back to the subsurface reservoirs (Fig. 2.11). Although it was suggested that silicates become more soluble in reduced H2OeCH4-bearing fluids with increasing pressure, we need an approach to better estimate the exact values of this solubility. Based on the experimental data at 10e16 GPa (Litasov et al., 2014) we suggest that the silicate content of reduced fluid along the mantle adiabat is appropriate for diapiric melt movement by the dissolutioneprecipitation mechanismd5e10 wt%. This means that hydrocarbon-bearing hydrous melts can play the same role in the lower mantle as do carbonatite melts in the transition zone. Moreover, this reduced melt behaves in the same way through the whole depth interval from the CMB to LAB (Fig. 2.11). However, there are no experimental data for quantitative estimations. For example,

REFERENCES

71

we cannot extrapolate data at 10e16 GPa to CMB pressures, because of the increasing stability of heavier and unsaturated hydrocarbons (Lobanov et al., 2013), which may have a different effect on the solubility of silicates in the hydrocarboneH2O mixture. Alternatively, H2O may be the principal volatile in the reduced fluid (Fig. 2.10) in the lower mantle coexisting with diamond or carbide. Again, we do not know the details of the H2O interaction with silicates at very high pressures as a function of redox conditions. They may be quite different from what we expect based on the low-pressure data.

7. CONCLUDING REMARKS In conclusion, we emphasize fundamental differences between melting in the petrologically important systems containing H2O, CO2, and with a reduced CeOeH fluid. Melting in H2O-bearing systems is controlled by hydrogen solubility in nominally anhydrous minerals and occurs when silicates are supersaturated with H2O at defined PeTeXef O2 conditions. Melting in CO2-bearing systems is determined by the stability of alkali carbonates and controlled mainly by the amount of Na2O and K2O in the system. Significant melting of subducted carbonates is expected at transition zone depths, especially if the slab stagnates above the 660 km discontinuity. Melting in systems with reduced CeOeH fluid occurs at intermediate temperatures between H2O- and CO2-bearing and volatile-free systems. Mantle melting in the presence of volatiles depends strongly on the redox state. An increase or decrease of the fO2 causes redox melting or freezing in defined parts of the mantle. The role of sluggish kinetics of solidesolid and even melting reactions involving carbonates in the deep mantle is also of great important. Slow reaction rates allow subducted carbonates to penetrate as deep as the CMB. We focus on the possible compositional properties of melts, which can enhance material transport by mantle plumes originating from the transition zone or the CMB. An important requirement for plume motion is stress-induced melting and dissolutioneprecipitation of the fusible components. Carbonated or carbonatite melt is the best candidate for the fusible component of the plumes, especially for the upper mantle and transition zone. Hydrocarbon-bearing hydrous melt can serve as the liquid component in mantle plumes arising from the CMB.

ACKNOWLEDGMENTS We thank S. Foley and two reviewers for valuable comments and Y. Kono and C. Sanloup for editorial handling of this chapter. The work was supported by Russian Science Foundation (No 14-17-00601, 14-17-00609P, and 17-17-01177).

REFERENCES Anderson, D.L., 1981. Rise of deep diapirs. Geology 9, 7e9. Andrault, D., Bolfan-Casanova, N., Nigro, G.L., Bouhifd, M.A., Garbarino, G., Mezouar, M., 2011. Solidus and liquidus profiles of chondritic mantle: implication for melting of the earth across its history. Earth Planet. Sci. Lett. 304, 251e259. Andrault, D., Pesce, G., Bouhifd, M.A., Bolfan-Casanova, N., He´not, J.-M., Mezouar, M., 2014. Melting of subducted basalt at the core-mantle boundary. Science 344, 892e895.

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