# CO2-induced destabilization of pyrite-structured FeO2Hx in the lower mantle

CO2-induced destabilization of pyrite-structured FeO2Hx in the lower mantle Abstract Volatiles, such as carbon and water, modulate the Earth's mantle rheology, partial melting and redox state, thereby playing a crucial role in the Earth's internal dynamics. We experimentally show the transformation of goethite FeOOH in the presence of CO2 into a tetrahedral carbonate phase, Fe4C3O12, at conditions above 107 GPa—2300 K. At temperatures below 2300 K, no interactions are evidenced between goethite and CO2, and instead a pyrite-structured FeO2Hx is formed as recently reported by Hu et al. (2016; 2017) and Nishi et al. (2017). The interpretation is that, above a critical temperature, FeO2Hx reacts with CO2 and H2, yielding Fe4C3O12 and H2O. Our findings provide strong support for the stability of carbon-oxygen-bearing phases at lower-mantle conditions. In both subducting slabs and lower-mantle lithologies, the tetrahedral carbonate Fe4C3O12 would replace the pyrite-structured FeO2Hx through carbonation of these phases. This reaction provides a new mechanism for hydrogen release as H2O within the deep lower mantle. Our study shows that the deep carbon and hydrogen cycles may be more complex than previously thought, as they strongly depend on the control exerted by local mineralogical and chemical environments on the CO2 and H2 thermodynamic activities. deep carbon cycle, FeOOH, high pressure INTRODUCTION Water (H2O) and carbon dioxide (CO2) both play an important role in the history of the Earth, as they strongly influence the chemical and physical properties of minerals, melts and fluids. Distribution and circulation of H2O and CO2 between the Earth's surface and the mantle have dominated the evolution of the crust, the oceans and the atmosphere, controlling several aspects of the Earth's habitability. It is therefore crucial to determine the stability and circulation of hydrous and CO2-bearing minerals in the Earth's interior. Sedimentary material together with altered mafic and ultramafic rocks that constitute the subducted slabs represents the main source for recycling of H2O and CO2 as well as other volatiles at great depth, possibly down to the core-–mantle boundary. The transport of H2O and CO2 via subducting slabs down to the transition zone and to the lower mantle has been the subject of many studies but is still under debate [1,2]. As for the carbon cycle, carbonates preserved during subduction are estimated to account for a flux of 3.6 × 1012 mol/year of carbon being returned into the deep mantle [3–5]. This quantity accounts for 10–30 wt % of the carbon reservoir in the deep mantle [6]. Regarding the water cycle, Van Keken et al. [2] suggested that 4–6 × 1013 mol/year of H2O are recycled into the mantle through slab subduction. Dehydration of the slab accounts for the loss of two-thirds of this amount of H2O, while one-third of the H2O remains bounded to the slab (i.e. ≈1.5 × 1013 mol/year) reaching depths exceeding 240 km. Although this amount of H2O entering the deep mantle may not appear very large, it provides a mechanism for having significant amounts of water in the deep mantle. In addition, part of the CO2 and H2O present in the deep mantle may also originate from primitive mantle reservoirs [7], leading potentially to fairly large amounts of these volatiles in the deep mantle. Because of its very low solubility in deep Earth's minerals [8,9], carbon is expected to be present as accessory phases in the mantle, either as oxidized phases such as carbonates or CO2 and carbonated fluids or melts, or as reduced phases such as diamonds or Fe-C alloys [10]. It is commonly considered that the lower mantle is too reducing to host carbonates [11,12]. However, the relatively high oxygen fugacities prevailing in subducting slabs might contribute to preserve oxidized carbon-bearing phases in the deep mantle [13,14]. Moreover, it has recently been demonstrated that carbonates at lower-mantle conditions adopt oxidized iron-bearing structures based on CO4 tetrahedra that are associated with reduced carbon phases [13,15–17]. Little is known about the stability of these new tetrahedral carbon-bearing phases but their systematic association with reduced carbon suggests the idea that the mineralogies of the lower mantle and D″region may be more complex than previously thought. Interestingly, carbonate-bearing inclusions have also been reported in diamonds formed in the lower mantle. This suggests again the presence of carbonates in the deep Earth and a possible coexistence of reduced and oxidized carbon species [15,18,19]. Decarbonation reactions of carbonates involving silicates (SiO2 and MgSiO3) were also reported to take place as shallow as ∼600 km in depth (20 GPa) [20,21]. Such reactions could produce CO2 in the lower mantle. Given the current uncertainties on the phase diagram of CO2 at high pressures, CO2 may be expressed as a solid CO2-V phase [21] or rather dissociate as C + O2 [22]. Carbonated fluids yet unknown at such conditions might also contribute to CO2 transfer at large depth in the mantle. In any case, large thermodynamic activities of CO2 are plausible in the lower mantle. A significant amount of water can be dissolved in nominally anhydrous minerals such as olivine, garnet and stishovite [23], as well as in high-pressure silicates such as wadsleyite and ringwoodite [24,25]. In addition, diverse dense hydrous silicates are stable in mafic and ultramafic assemblages at upper- and lower-mantle conditions, such as phase A, phase D, phase H and superhydrous phase B [26–30]. Finally, δ-AlOOH, a high-pressure form of diaspore (α-AlOOH) with an orthorhombic symmetry very close to that of the CaCl2-type polymorph of SiO2, is stable throughout the mantle and may be present in suitably aluminous and hydrated lithologies [31,32]. The high-pressure polymorph ε-FeOOH that shares the same structure with δ-AlOOH [33] might also store water in the mantle. Iron oxyhydroxides, including FeOOH and its polymorphs, are common at the surface of the Earth, where they are abundant in soils and sediments. The incorporation of hydrogen atoms in newly discovered iron oxyhydroxides with a pyrite structure [34–36] may thus contribute to the transfer of H to the deep Earth. In their recent work, Hu et al. [34] suggested that a phase of FeO2Hx composition might indeed deliver H2 instead of water when heated above a threshold temperature—a particularity due to valence changes of oxygen in this compound (from 2O2– to O22–) [37]. After H2O, CO2 is the second most important volatile compound in the deep Earth. To get a more complete understanding of the H and C cycles in the deep Earth, it is necessary to know how deep subducted materials can transport both C and H and identify the distinct species involved. In the present study, we shed light on this crucial issue by constraining experimentally the interactions of CO2 with potential carriers of H2O or H2 at great depths. We performed high-pressure and high-temperature diamond-anvil cell (DAC) experiments to investigate the effects of a CO2-rich medium on the transformations of FeOOH at pressures and temperatures of the lower mantle. RESULTS A natural sample of crystalline α-FeOOH (Supplementary Fig. 1) was loaded in CO2 and first pressurized up to 107 GPa in a DAC at ambient temperature. In situ X-Ray diffraction (XRD) patterns showed a significant broadening of α-FeOOH main diffraction reflections characteristic of incipient amorphization. After laser heating at 2000 K for a few minutes, several changes in the diffraction pattern were observed: the α-FeOOH phase disappeared and two distinct phases were identified (Fig. 1). The most intense diffraction peaks correspond to a cubic structure with extinctions of the two reflections 001 and 011 in agreement with a Pa-3 space group. Recently, Hu et al. [34] reported the transformation of FeOOH into a new Pa-3 cubic structure FeO2Hx at similar pressure and temperature conditions. This phase is directly related to the newly discovered pyrite-structured FeO2 peroxide but is characterized by a larger unit cell volume [34,37]. FeO2Hx can be interpreted as a solid solution between pyrite-structure FeO2 and FeOOH. In addition, a pyrite-structured FeOOH oxyhydroxide (i.e. FeO2Hx with x = 1) was recently observed experimentally by Nishi et al. [35]. It presents a structure close to the pyrite-type structure of AlOOH predicted above 170 GPa by Tsuchiya and Tsuchiya [38]. Here, we measured a unit cell parameter of a = 4.367 Å at 107 GPa, which is significantly larger than that reported for FeO2 (a = 4.363 Å at 76 GPa) by Hu et al. [37], but smaller than that reported for FeOOH (a = 4.386 Å at 109 GPa) in Nishi et al. [35] (Fig. 2). It is thus probable that FeOOH in our study underwent partial dehydrogenation. A similar unit cell volume based on a FeO2Hx formula was reported at the same P-T conditions by Hu et al. [34], for which, using their calibration, they deduced x = 0.66. Because this calibration is not only built using experimental data, but also incorporates theoretical results, which are known to either overestimate or underestimate unit cell volume, uncertainty on the exact amount of hydrogen ‘x’ present in FeO2Hx may be high. To account for this, we simply refer to this phase as FeO2Hx. Figure 1. View largeDownload slide XRD patterns collected at 2000 K and at 2300 K and LeBail unit cell refinement of the two phases FeO2Hx (red markers, space group Pa-3, a = 4.365(1)) and Fe4C3O12 (blue markers, space group P2, a = 9.697(2), b = 6.296(2), c = 5.726(1), beta = 92.94(2)). Red circles materialize the expected peak position of the cubic phase according the XRD collected at lower temperature. Figure 1. View largeDownload slide XRD patterns collected at 2000 K and at 2300 K and LeBail unit cell refinement of the two phases FeO2Hx (red markers, space group Pa-3, a = 4.365(1)) and Fe4C3O12 (blue markers, space group P2, a = 9.697(2), b = 6.296(2), c = 5.726(1), beta = 92.94(2)). Red circles materialize the expected peak position of the cubic phase according the XRD collected at lower temperature. Figure 2. View largeDownload slide (a) Rietveld refinement of the XRD pattern collected at 107 GPa and 300 K after laser heating with a FeO2H pyrite-structured (right hand). (b) Unit cell volume of the pyrite-type structure measured experimentally as a function of pressure and temperature for FeO2 [37], FeO2Hx ([31], from both Fe2O3+H2O and FeOOH in Ar experiments and this study) and FeOOH [35]. All the data reported here were collected after quenching the temperature. Figure 2. View largeDownload slide (a) Rietveld refinement of the XRD pattern collected at 107 GPa and 300 K after laser heating with a FeO2H pyrite-structured (right hand). (b) Unit cell volume of the pyrite-type structure measured experimentally as a function of pressure and temperature for FeO2 [37], FeO2Hx ([31], from both Fe2O3+H2O and FeOOH in Ar experiments and this study) and FeOOH [35]. All the data reported here were collected after quenching the temperature. In the XRD pattern, the less intense diffraction peaks can be assigned to an already discovered carbon-rich phase stable at these P-T conditions: Fe4C3O12 [15] (Fig. 1). Among the five structures proposed in literature [15,17,39,40], we found that only the monoclinic structure reported in [15] allowed us to assign all of the observed diffraction peaks. Although ex situ analyses of the hydrogen content of this phase would be necessary, the fact that we measured unit cell parameters in very good agreement with that reported in [15] for a hydrogen-free composition leads us to propose an Fe4C3O12 stoichiometry. Upon heating at higher temperature, diffraction peaks of Fe4C3O12 increase in intensity at the expense of the FeO2Hx cubic phase, which fully disappears above 2300 K (Fig. 1). Neither iron oxides (e.g. Fe3O4, Fe2O3, Fe4O5, Fe13O19) nor diamond were observed in these experiments. We note that Fe in Fe4C3O12 is ferric iron Fe(III) as in FeOOH. After laser heating, we collected a profile of diffraction patterns across the heated spot (Supplementary Fig. 2). The FeO2Hx cubic phase was observed at the edge of the 2300-K heated spot only. Although it is theoretically possible that the reaction is kinetically restricted at lower temperatures, the fact that a change occurred abruptly at around 2300 K (FeO2Hx disappears within a few seconds) more probably pinpoints to a thermodynamic boundary. The pyrite-structured FeO2Hx would be stable only at relatively low temperatures in the Fe-O-C-H system provided CO2 thermodynamic activities are high enough. Further studies should verify this point. Raman spectra were also collected at ambient temperature and high pressure in the 300–1300 cm−1 range. As presented in Fig. 3, when collected in areas where the CO2 loading gas was pure, the spectra reveal only one low intensity mode at ∼1123 cm−1, which corresponds to the most intense mode A1g of CO2-VI [41]. In the solid sample area, an additional mode was detected at ∼930 cm−1 assigned to the T2g mode from the high-pressure phase of H2O ice-X [42,43]. No Raman active modes associated with Fe4C3O12 could be observed, which may be due to high fluorescence background of the diamond from the DAC. Figure 3. View largeDownload slide Raman spectra collected after laser heating in the CO2 area (in black) and sample area (in red). Figure 3. View largeDownload slide Raman spectra collected after laser heating in the CO2 area (in black) and sample area (in red). Transmission electron microscopy (TEM) analyses of a thin section extracted from the recovered sample are reported in Fig. 4. Semi-quantitative chemical analyses (XEDS) showed a homogenous composition with carbon, iron and oxygen with Fe-O atomic proportions consistently with Fe4C3O12 (Fig. 4A). The sample was unstable under electron beam, and selective area electron diffraction (SAED) revealed the presence of two patterns (Fig. 4B): γ-Fe2O3 maghemite coexisting with a phase characterized by a 6-fold symmetry diffraction pattern that could not be indexed. It is probable that, under the electron beam, Fe4C3O12 underwent carbon loss to form γ-Fe2O3 together with a second phase still containing carbon. Note that the observed texture, often observed in cases of irradiation damages, is in agreement with amorphization and devolatilization of the sample under the electron beam (Fig. 4C). Figure 4. View largeDownload slide (a) Semi-quantitative chemical analyses (EDX); (b) electron diffraction of Fe2O3 maghemite (white markers) together with an unknown phase (yellow markers); and (c) TEM picture of the sample after analyses. Figure 4. View largeDownload slide (a) Semi-quantitative chemical analyses (EDX); (b) electron diffraction of Fe2O3 maghemite (white markers) together with an unknown phase (yellow markers); and (c) TEM picture of the sample after analyses. DISCUSSION This study demonstrates that, at pressures of about 110 GPa and upon laser heating, a chemical reaction occurs between FeO2Hx and CO2 yielding a tetrahedral carbonate Fe4C3O12. The transformation from the initial goethite FeOOH with increasing temperature can be schematized as: $${\rm{FeOOH}} = > {\rm{Fe}}{{\rm{O}}_2}{{\rm{H}}_{\rm{x}}} + 1/2\left( {1 - {\rm{x}}} \right){{\rm{H}}_2}$$ (1) \begin{eqnarray} 4{\rm{Fe}}{{\rm{O}}_2}{{\rm{H}}_{\rm{x}}} + 3{\rm{C}}{{\rm{O}}_2} &+& 2\left( {1 - {\rm{x}}} \right){{\rm{H}}_2} = > {\rm{F}}{{\rm{e}}_4}{{\rm{C}}_3}{{\rm{O}}_{12}}\nonumber\\ &&+ 2{{\rm{H}}_2}{\rm{O}} \end{eqnarray} (2) which can be summed up as: $$4{\rm{FeOOH}} + 3{\rm{C}}{{\rm{O}}_2} = > {\rm{F}}{{\rm{e}}_4}{{\rm{C}}_3}{{\rm{O}}_{12}} + 2{{\rm{H}}_2}{\rm{O}}{\rm{.}}$$ (3) If the local thermodynamic activity of H2 is too low for reaction (2) to proceed, other reactions are possible, such as: \begin{eqnarray} 4{\rm{Fe}}{{\rm{O}}_2}{{\rm{H}}_{\rm{x}}} + 3{\rm{C}}{{\rm{O}}_2} &= >& {\rm{F}}{{\rm{e}}_4}{{\rm{C}}_3}{{\rm{O}}_{12}} + 2{\rm{x}}{{\rm{H}}_2}{\rm{O}}\nonumber\\ &&+ \left( {1 - {\rm{x}}} \right){{\rm{O}}_2}, \end{eqnarray} (4) which might have interesting consequences for oxygen fugacity at large depth and by consequence at the Earth’s surface. The P-T conditions at which FeO2Hx and Fe4C3O12 have been observed are presented in Fig. 5 along with mantle geotherms and hypothetical slab geotherms [44,45]. The exact chemistry and stability of the high-pressure pyrite-structured FeO2Hx are still controversial: Nishi et al. [35] propose a pyrite-structured oxyhydroxide FeOOH that is stable down to the core--mantle boundary and might undergo dehydration in the D″layer, whereas Hu et al. [34] and Liu et al. [46] suggest a pyrite-structured peroxide/hydride FeO2Hx that would undergo progressive dehydrogenation from about 1800-km depth down to the core–mantle boundary. However, our present study demonstrates that the presence of CO2, produced for example by decarbonation reactions involving silicate phases, could completely alter these interpretations. Indeed, the pyrite-structured FeO2Hx would react with CO2 to form a high-pressure carbon-bearing phase Fe4C3O12 at P-T conditions of the lower-mantle geotherm, as well as of those of a ‘hot’ slab path (such as Central America slabs [47]), and even on geotherms of cold slabs close to the core--mantle boundary. Unfortunately, we currently lack thermodynamic constraints to evaluate the activity of CO2 in the mantle and its stability relative to carbonates or C-reduced species. This should be addressed in the future to confirm that Fe4C3O12-forming reaction actually takes place in the mantle. Although the thermodynamic stability of tetrahedral carbonates with respect to reduced carbon phases is still unknown, it appears that Fe4C3O12 tetrahedral carbonate is an excellent candidate for a stable carbon host in the lower mantle [15,40]. The carbonation reaction (R3) is associated with release of H2O. Therefore, the carbonation reaction provides a new mechanism for releasing hydrogen into the deep mantle as H2O. It adds up to dehydration reactions that take place at shallower depths in subduction settings and to the progressive dehydrogenation of FeOOH at about 1800-km depth [34,46]. Similarly to carbon [12,48], H2 would be oxidized to produce OH or H2O through the reduction of Fe3+ in silicate minerals during mantle upwelling. Such release of OH or H2O could trigger partial melting, since H2O is much more soluble in silicate melts than H2 [49,50]. In hot subducting slabs, the carbonation reaction from oxyhydroxide may take place as shallow as 1200 km [15], before any transformation of α -FeOOH into FeO2Hx. In environments rich in iron oxides such as hematite Fe2O3 (e.g. in banded iron formation lithology), Fe2O3 may directly react with CO2 [15] without implication of FeOOH in the chemical reaction. In this latter scenario, carbon and hydrogen would be both transported in the deep mantle without dehydrogenation due to carbonation (although it is possible that slow dehydrogenation of FeO2Hx takes place [34]). The degree of coupling between the deep carbon and hydrogen cycles is therefore strongly dependent on the local mineralogical and chemical environment. Because carbonates are also potential oxidized carbon carriers, additional studies on the interaction between carbonates and FeOOH should be carried out in order to provide a comprehensive model for the deep-mantle carbon and water cycles. Nevertheless, the transformation reported here would prevent the production of FeHx, which is expected by the reaction of iron alloy from the core and hydrous phase at the core–mantle boundary [28,35,51]. This might have favored transfer of carbon to the core rather than of hydrogen during early Earth differentiation and therefore provide a mechanism for high amounts of C in an O-rich core [52]. Figure 5. View largeDownload slide P-T conditions at which the different phases have been observed experimentally: pyrite-structured FeO2 from Hu et al. [37], pyrite-structured FeO2Hx from this study as well as from Hu et al. [34] from both Fe2O3+H2O and FeOOH in Ar/Ne experiments, pyrite-structured FeOOH from Nishi et al. [35] and Fe4C3O12 from the present study as well as from FeO + CO2 experiments in Boulard et al. [15]. The mantle geotherms from [59] (O(08)) and [60] (A(82)), as well as hypothetical P-T paths for a ‘very cold slab’, ‘cold slab’ and a ‘hot slab’ [45,61] are also represented for comparison. Figure 5. View largeDownload slide P-T conditions at which the different phases have been observed experimentally: pyrite-structured FeO2 from Hu et al. [37], pyrite-structured FeO2Hx from this study as well as from Hu et al. [34] from both Fe2O3+H2O and FeOOH in Ar/Ne experiments, pyrite-structured FeOOH from Nishi et al. [35] and Fe4C3O12 from the present study as well as from FeO + CO2 experiments in Boulard et al. [15]. The mantle geotherms from [59] (O(08)) and [60] (A(82)), as well as hypothetical P-T paths for a ‘very cold slab’, ‘cold slab’ and a ‘hot slab’ [45,61] are also represented for comparison. MATERIALS AND METHODS Experiments were conducted using symmetric Mao-Bell-type DAC equipped with 300/100-μm beveled culet diamonds and rhenium gasket with 30-μm starting diameter hole and 25-μm starting thickness. Natural sample of crystalline goethite (α-FeOOH) from lateritic soil in Central African Republic was provided by the collection of University Pierre et Marie Curie. FeOOH was loaded in CO2 together with a ruby ball (see Supplementary Fig. 1 for an XRD characterization before the experiment). The CO2 gas was loaded using high-pressure gas-loading apparatus at room temperature and 600 bars. FeOOH was isolated from diamonds by CO2, preventing reactions with diamonds. In situ angle-dispersive XRD measurements were performed on the high-pressure beamline ID-27 at the European Synchrotron Radiation Facility (ESRF) using a monochromatic incident X-ray beam of 0.3738-Å wavelength. Before the experiments, the X-ray spot, spectrometer entrance and the heating laser spot were carefully aligned. The sample was first pressurized to its target pressure (107 GPa) before laser heating. Pressure was measured using ruby fluorescence before and after laser heating at room temperature by the use of a blue laser [53]. XRD peaks from the Re-gasket collected across the diamond culet gave pressures between 101 and 107 GPa. Two YAG lasers with excellent power stability were aligned on both sides of the sample, which produce hot spots larger than 20-μm (FWHM) diameter. Temperatures were obtained by fitting the sample thermal emission spectrum from the central 2 × 2 μm2 of the hotspot to the Planck's function using the wavelength range 600–900 nm. Reflective lenses were used for measurement in order to prevent any chromatic aberration [54]. The monochromatic X-ray beam was focused to 3 × 3 micron. This is smaller than the laser heating spot in order to reduce both the radial and axial temperature gradients. Typical exposure time was 30 s at high pressures and high temperatures. The diffraction images were integrated with the Fit2d software [57]. The 1D diffraction patterns were treated with the General Structure Analysis System (GSAS) software package [58] using the Rietveld or Lebail methods to identify the different phases and refine lattice parameters. During heating and XRD acquisition, temperature was measured continuously. Temperature uncertainties are estimated to be of about 150 K [55]. At high temperatures, thermal pressure corrections are of the order of +10–15% of the initial pressure [56]. Raman spectra were collected at high pressure and ambient temperature after transformation of the sample. We used a Jobin–Yvon® HR-460 spectrometer with monochromator with 1500 gratings/mm, equipped with an Andor® CCD camera. Raman signal was excited using the 514.5-nm wavelength of an Ar+ laser, delivering 300 mW focused into a 2-μm spot by a long-working distance Mitutoyo® x20 objective. A focused ion beam (FIB) thin section was extracted from the recovered sample at the center of laser-heated spot and thinned to electron transparency (∼100-nm thickness). FIB milling was performed using a FEI STRATA DB 235 at IEMN (Lille, France) with a focused Ga+ ion beam operating at 30 kV and currents from 20 nA down to 1 pA for final surfacing. Analytical transmission electron microscopy (ATEM) was carried out on the FIB thin section with a JEOL 2100-F operating at 200 keV, equipped with a field emission gun. Semi-quantitative chemical analyses on the individual phases was obtained by X-ray energy dispersive spectrometry (XEDS) and SAED patterns were used for phase identification. SUPPLEMENTARY DATA Supplementary data are available at NSR online. Acknowledgements We acknowledge the European Synchrotron Radiation Facility for the allocation of beam time. A.C. acknowledges a research sabbatical leave support from the French National Center for Scientific Research (CNRS). The transmission electron microscopy facility at Institut de Minéralogie, Physique des Matériaux et Cosmochimie is supported by Région Ile de France grant SESAME 2000 E 1435. The manuscript was significantly improved by the constructive comments from three anonymous reviewers. REFERENCES 1. Kelemen PB , Manning CE . 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# CO2-induced destabilization of pyrite-structured FeO2Hx in the lower mantle

8 pages

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Oxford University Press
© The Author(s) 2018. Published by Oxford University Press on behalf of China Science Publishing & Media Ltd.
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2095-5138
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2053-714X
D.O.I.
10.1093/nsr/nwy032
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### Abstract

Abstract Volatiles, such as carbon and water, modulate the Earth's mantle rheology, partial melting and redox state, thereby playing a crucial role in the Earth's internal dynamics. We experimentally show the transformation of goethite FeOOH in the presence of CO2 into a tetrahedral carbonate phase, Fe4C3O12, at conditions above 107 GPa—2300 K. At temperatures below 2300 K, no interactions are evidenced between goethite and CO2, and instead a pyrite-structured FeO2Hx is formed as recently reported by Hu et al. (2016; 2017) and Nishi et al. (2017). The interpretation is that, above a critical temperature, FeO2Hx reacts with CO2 and H2, yielding Fe4C3O12 and H2O. Our findings provide strong support for the stability of carbon-oxygen-bearing phases at lower-mantle conditions. In both subducting slabs and lower-mantle lithologies, the tetrahedral carbonate Fe4C3O12 would replace the pyrite-structured FeO2Hx through carbonation of these phases. This reaction provides a new mechanism for hydrogen release as H2O within the deep lower mantle. Our study shows that the deep carbon and hydrogen cycles may be more complex than previously thought, as they strongly depend on the control exerted by local mineralogical and chemical environments on the CO2 and H2 thermodynamic activities. deep carbon cycle, FeOOH, high pressure INTRODUCTION Water (H2O) and carbon dioxide (CO2) both play an important role in the history of the Earth, as they strongly influence the chemical and physical properties of minerals, melts and fluids. Distribution and circulation of H2O and CO2 between the Earth's surface and the mantle have dominated the evolution of the crust, the oceans and the atmosphere, controlling several aspects of the Earth's habitability. It is therefore crucial to determine the stability and circulation of hydrous and CO2-bearing minerals in the Earth's interior. Sedimentary material together with altered mafic and ultramafic rocks that constitute the subducted slabs represents the main source for recycling of H2O and CO2 as well as other volatiles at great depth, possibly down to the core-–mantle boundary. The transport of H2O and CO2 via subducting slabs down to the transition zone and to the lower mantle has been the subject of many studies but is still under debate [1,2]. As for the carbon cycle, carbonates preserved during subduction are estimated to account for a flux of 3.6 × 1012 mol/year of carbon being returned into the deep mantle [3–5]. This quantity accounts for 10–30 wt % of the carbon reservoir in the deep mantle [6]. Regarding the water cycle, Van Keken et al. [2] suggested that 4–6 × 1013 mol/year of H2O are recycled into the mantle through slab subduction. Dehydration of the slab accounts for the loss of two-thirds of this amount of H2O, while one-third of the H2O remains bounded to the slab (i.e. ≈1.5 × 1013 mol/year) reaching depths exceeding 240 km. Although this amount of H2O entering the deep mantle may not appear very large, it provides a mechanism for having significant amounts of water in the deep mantle. In addition, part of the CO2 and H2O present in the deep mantle may also originate from primitive mantle reservoirs [7], leading potentially to fairly large amounts of these volatiles in the deep mantle. Because of its very low solubility in deep Earth's minerals [8,9], carbon is expected to be present as accessory phases in the mantle, either as oxidized phases such as carbonates or CO2 and carbonated fluids or melts, or as reduced phases such as diamonds or Fe-C alloys [10]. It is commonly considered that the lower mantle is too reducing to host carbonates [11,12]. However, the relatively high oxygen fugacities prevailing in subducting slabs might contribute to preserve oxidized carbon-bearing phases in the deep mantle [13,14]. Moreover, it has recently been demonstrated that carbonates at lower-mantle conditions adopt oxidized iron-bearing structures based on CO4 tetrahedra that are associated with reduced carbon phases [13,15–17]. Little is known about the stability of these new tetrahedral carbon-bearing phases but their systematic association with reduced carbon suggests the idea that the mineralogies of the lower mantle and D″region may be more complex than previously thought. Interestingly, carbonate-bearing inclusions have also been reported in diamonds formed in the lower mantle. This suggests again the presence of carbonates in the deep Earth and a possible coexistence of reduced and oxidized carbon species [15,18,19]. Decarbonation reactions of carbonates involving silicates (SiO2 and MgSiO3) were also reported to take place as shallow as ∼600 km in depth (20 GPa) [20,21]. Such reactions could produce CO2 in the lower mantle. Given the current uncertainties on the phase diagram of CO2 at high pressures, CO2 may be expressed as a solid CO2-V phase [21] or rather dissociate as C + O2 [22]. Carbonated fluids yet unknown at such conditions might also contribute to CO2 transfer at large depth in the mantle. In any case, large thermodynamic activities of CO2 are plausible in the lower mantle. A significant amount of water can be dissolved in nominally anhydrous minerals such as olivine, garnet and stishovite [23], as well as in high-pressure silicates such as wadsleyite and ringwoodite [24,25]. In addition, diverse dense hydrous silicates are stable in mafic and ultramafic assemblages at upper- and lower-mantle conditions, such as phase A, phase D, phase H and superhydrous phase B [26–30]. Finally, δ-AlOOH, a high-pressure form of diaspore (α-AlOOH) with an orthorhombic symmetry very close to that of the CaCl2-type polymorph of SiO2, is stable throughout the mantle and may be present in suitably aluminous and hydrated lithologies [31,32]. The high-pressure polymorph ε-FeOOH that shares the same structure with δ-AlOOH [33] might also store water in the mantle. Iron oxyhydroxides, including FeOOH and its polymorphs, are common at the surface of the Earth, where they are abundant in soils and sediments. The incorporation of hydrogen atoms in newly discovered iron oxyhydroxides with a pyrite structure [34–36] may thus contribute to the transfer of H to the deep Earth. In their recent work, Hu et al. [34] suggested that a phase of FeO2Hx composition might indeed deliver H2 instead of water when heated above a threshold temperature—a particularity due to valence changes of oxygen in this compound (from 2O2– to O22–) [37]. After H2O, CO2 is the second most important volatile compound in the deep Earth. To get a more complete understanding of the H and C cycles in the deep Earth, it is necessary to know how deep subducted materials can transport both C and H and identify the distinct species involved. In the present study, we shed light on this crucial issue by constraining experimentally the interactions of CO2 with potential carriers of H2O or H2 at great depths. We performed high-pressure and high-temperature diamond-anvil cell (DAC) experiments to investigate the effects of a CO2-rich medium on the transformations of FeOOH at pressures and temperatures of the lower mantle. RESULTS A natural sample of crystalline α-FeOOH (Supplementary Fig. 1) was loaded in CO2 and first pressurized up to 107 GPa in a DAC at ambient temperature. In situ X-Ray diffraction (XRD) patterns showed a significant broadening of α-FeOOH main diffraction reflections characteristic of incipient amorphization. After laser heating at 2000 K for a few minutes, several changes in the diffraction pattern were observed: the α-FeOOH phase disappeared and two distinct phases were identified (Fig. 1). The most intense diffraction peaks correspond to a cubic structure with extinctions of the two reflections 001 and 011 in agreement with a Pa-3 space group. Recently, Hu et al. [34] reported the transformation of FeOOH into a new Pa-3 cubic structure FeO2Hx at similar pressure and temperature conditions. This phase is directly related to the newly discovered pyrite-structured FeO2 peroxide but is characterized by a larger unit cell volume [34,37]. FeO2Hx can be interpreted as a solid solution between pyrite-structure FeO2 and FeOOH. In addition, a pyrite-structured FeOOH oxyhydroxide (i.e. FeO2Hx with x = 1) was recently observed experimentally by Nishi et al. [35]. It presents a structure close to the pyrite-type structure of AlOOH predicted above 170 GPa by Tsuchiya and Tsuchiya [38]. Here, we measured a unit cell parameter of a = 4.367 Å at 107 GPa, which is significantly larger than that reported for FeO2 (a = 4.363 Å at 76 GPa) by Hu et al. [37], but smaller than that reported for FeOOH (a = 4.386 Å at 109 GPa) in Nishi et al. [35] (Fig. 2). It is thus probable that FeOOH in our study underwent partial dehydrogenation. A similar unit cell volume based on a FeO2Hx formula was reported at the same P-T conditions by Hu et al. [34], for which, using their calibration, they deduced x = 0.66. Because this calibration is not only built using experimental data, but also incorporates theoretical results, which are known to either overestimate or underestimate unit cell volume, uncertainty on the exact amount of hydrogen ‘x’ present in FeO2Hx may be high. To account for this, we simply refer to this phase as FeO2Hx. Figure 1. View largeDownload slide XRD patterns collected at 2000 K and at 2300 K and LeBail unit cell refinement of the two phases FeO2Hx (red markers, space group Pa-3, a = 4.365(1)) and Fe4C3O12 (blue markers, space group P2, a = 9.697(2), b = 6.296(2), c = 5.726(1), beta = 92.94(2)). Red circles materialize the expected peak position of the cubic phase according the XRD collected at lower temperature. Figure 1. View largeDownload slide XRD patterns collected at 2000 K and at 2300 K and LeBail unit cell refinement of the two phases FeO2Hx (red markers, space group Pa-3, a = 4.365(1)) and Fe4C3O12 (blue markers, space group P2, a = 9.697(2), b = 6.296(2), c = 5.726(1), beta = 92.94(2)). Red circles materialize the expected peak position of the cubic phase according the XRD collected at lower temperature. Figure 2. View largeDownload slide (a) Rietveld refinement of the XRD pattern collected at 107 GPa and 300 K after laser heating with a FeO2H pyrite-structured (right hand). (b) Unit cell volume of the pyrite-type structure measured experimentally as a function of pressure and temperature for FeO2 [37], FeO2Hx ([31], from both Fe2O3+H2O and FeOOH in Ar experiments and this study) and FeOOH [35]. All the data reported here were collected after quenching the temperature. Figure 2. View largeDownload slide (a) Rietveld refinement of the XRD pattern collected at 107 GPa and 300 K after laser heating with a FeO2H pyrite-structured (right hand). (b) Unit cell volume of the pyrite-type structure measured experimentally as a function of pressure and temperature for FeO2 [37], FeO2Hx ([31], from both Fe2O3+H2O and FeOOH in Ar experiments and this study) and FeOOH [35]. All the data reported here were collected after quenching the temperature. In the XRD pattern, the less intense diffraction peaks can be assigned to an already discovered carbon-rich phase stable at these P-T conditions: Fe4C3O12 [15] (Fig. 1). Among the five structures proposed in literature [15,17,39,40], we found that only the monoclinic structure reported in [15] allowed us to assign all of the observed diffraction peaks. Although ex situ analyses of the hydrogen content of this phase would be necessary, the fact that we measured unit cell parameters in very good agreement with that reported in [15] for a hydrogen-free composition leads us to propose an Fe4C3O12 stoichiometry. Upon heating at higher temperature, diffraction peaks of Fe4C3O12 increase in intensity at the expense of the FeO2Hx cubic phase, which fully disappears above 2300 K (Fig. 1). Neither iron oxides (e.g. Fe3O4, Fe2O3, Fe4O5, Fe13O19) nor diamond were observed in these experiments. We note that Fe in Fe4C3O12 is ferric iron Fe(III) as in FeOOH. After laser heating, we collected a profile of diffraction patterns across the heated spot (Supplementary Fig. 2). The FeO2Hx cubic phase was observed at the edge of the 2300-K heated spot only. Although it is theoretically possible that the reaction is kinetically restricted at lower temperatures, the fact that a change occurred abruptly at around 2300 K (FeO2Hx disappears within a few seconds) more probably pinpoints to a thermodynamic boundary. The pyrite-structured FeO2Hx would be stable only at relatively low temperatures in the Fe-O-C-H system provided CO2 thermodynamic activities are high enough. Further studies should verify this point. Raman spectra were also collected at ambient temperature and high pressure in the 300–1300 cm−1 range. As presented in Fig. 3, when collected in areas where the CO2 loading gas was pure, the spectra reveal only one low intensity mode at ∼1123 cm−1, which corresponds to the most intense mode A1g of CO2-VI [41]. In the solid sample area, an additional mode was detected at ∼930 cm−1 assigned to the T2g mode from the high-pressure phase of H2O ice-X [42,43]. No Raman active modes associated with Fe4C3O12 could be observed, which may be due to high fluorescence background of the diamond from the DAC. Figure 3. View largeDownload slide Raman spectra collected after laser heating in the CO2 area (in black) and sample area (in red). Figure 3. View largeDownload slide Raman spectra collected after laser heating in the CO2 area (in black) and sample area (in red). Transmission electron microscopy (TEM) analyses of a thin section extracted from the recovered sample are reported in Fig. 4. Semi-quantitative chemical analyses (XEDS) showed a homogenous composition with carbon, iron and oxygen with Fe-O atomic proportions consistently with Fe4C3O12 (Fig. 4A). The sample was unstable under electron beam, and selective area electron diffraction (SAED) revealed the presence of two patterns (Fig. 4B): γ-Fe2O3 maghemite coexisting with a phase characterized by a 6-fold symmetry diffraction pattern that could not be indexed. It is probable that, under the electron beam, Fe4C3O12 underwent carbon loss to form γ-Fe2O3 together with a second phase still containing carbon. Note that the observed texture, often observed in cases of irradiation damages, is in agreement with amorphization and devolatilization of the sample under the electron beam (Fig. 4C). Figure 4. View largeDownload slide (a) Semi-quantitative chemical analyses (EDX); (b) electron diffraction of Fe2O3 maghemite (white markers) together with an unknown phase (yellow markers); and (c) TEM picture of the sample after analyses. Figure 4. View largeDownload slide (a) Semi-quantitative chemical analyses (EDX); (b) electron diffraction of Fe2O3 maghemite (white markers) together with an unknown phase (yellow markers); and (c) TEM picture of the sample after analyses. DISCUSSION This study demonstrates that, at pressures of about 110 GPa and upon laser heating, a chemical reaction occurs between FeO2Hx and CO2 yielding a tetrahedral carbonate Fe4C3O12. The transformation from the initial goethite FeOOH with increasing temperature can be schematized as: $${\rm{FeOOH}} = > {\rm{Fe}}{{\rm{O}}_2}{{\rm{H}}_{\rm{x}}} + 1/2\left( {1 - {\rm{x}}} \right){{\rm{H}}_2}$$ (1) \begin{eqnarray} 4{\rm{Fe}}{{\rm{O}}_2}{{\rm{H}}_{\rm{x}}} + 3{\rm{C}}{{\rm{O}}_2} &+& 2\left( {1 - {\rm{x}}} \right){{\rm{H}}_2} = > {\rm{F}}{{\rm{e}}_4}{{\rm{C}}_3}{{\rm{O}}_{12}}\nonumber\\ &&+ 2{{\rm{H}}_2}{\rm{O}} \end{eqnarray} (2) which can be summed up as: $$4{\rm{FeOOH}} + 3{\rm{C}}{{\rm{O}}_2} = > {\rm{F}}{{\rm{e}}_4}{{\rm{C}}_3}{{\rm{O}}_{12}} + 2{{\rm{H}}_2}{\rm{O}}{\rm{.}}$$ (3) If the local thermodynamic activity of H2 is too low for reaction (2) to proceed, other reactions are possible, such as: \begin{eqnarray} 4{\rm{Fe}}{{\rm{O}}_2}{{\rm{H}}_{\rm{x}}} + 3{\rm{C}}{{\rm{O}}_2} &= >& {\rm{F}}{{\rm{e}}_4}{{\rm{C}}_3}{{\rm{O}}_{12}} + 2{\rm{x}}{{\rm{H}}_2}{\rm{O}}\nonumber\\ &&+ \left( {1 - {\rm{x}}} \right){{\rm{O}}_2}, \end{eqnarray} (4) which might have interesting consequences for oxygen fugacity at large depth and by consequence at the Earth’s surface. The P-T conditions at which FeO2Hx and Fe4C3O12 have been observed are presented in Fig. 5 along with mantle geotherms and hypothetical slab geotherms [44,45]. The exact chemistry and stability of the high-pressure pyrite-structured FeO2Hx are still controversial: Nishi et al. [35] propose a pyrite-structured oxyhydroxide FeOOH that is stable down to the core--mantle boundary and might undergo dehydration in the D″layer, whereas Hu et al. [34] and Liu et al. [46] suggest a pyrite-structured peroxide/hydride FeO2Hx that would undergo progressive dehydrogenation from about 1800-km depth down to the core–mantle boundary. However, our present study demonstrates that the presence of CO2, produced for example by decarbonation reactions involving silicate phases, could completely alter these interpretations. Indeed, the pyrite-structured FeO2Hx would react with CO2 to form a high-pressure carbon-bearing phase Fe4C3O12 at P-T conditions of the lower-mantle geotherm, as well as of those of a ‘hot’ slab path (such as Central America slabs [47]), and even on geotherms of cold slabs close to the core--mantle boundary. Unfortunately, we currently lack thermodynamic constraints to evaluate the activity of CO2 in the mantle and its stability relative to carbonates or C-reduced species. This should be addressed in the future to confirm that Fe4C3O12-forming reaction actually takes place in the mantle. Although the thermodynamic stability of tetrahedral carbonates with respect to reduced carbon phases is still unknown, it appears that Fe4C3O12 tetrahedral carbonate is an excellent candidate for a stable carbon host in the lower mantle [15,40]. The carbonation reaction (R3) is associated with release of H2O. Therefore, the carbonation reaction provides a new mechanism for releasing hydrogen into the deep mantle as H2O. It adds up to dehydration reactions that take place at shallower depths in subduction settings and to the progressive dehydrogenation of FeOOH at about 1800-km depth [34,46]. Similarly to carbon [12,48], H2 would be oxidized to produce OH or H2O through the reduction of Fe3+ in silicate minerals during mantle upwelling. Such release of OH or H2O could trigger partial melting, since H2O is much more soluble in silicate melts than H2 [49,50]. In hot subducting slabs, the carbonation reaction from oxyhydroxide may take place as shallow as 1200 km [15], before any transformation of α -FeOOH into FeO2Hx. In environments rich in iron oxides such as hematite Fe2O3 (e.g. in banded iron formation lithology), Fe2O3 may directly react with CO2 [15] without implication of FeOOH in the chemical reaction. In this latter scenario, carbon and hydrogen would be both transported in the deep mantle without dehydrogenation due to carbonation (although it is possible that slow dehydrogenation of FeO2Hx takes place [34]). The degree of coupling between the deep carbon and hydrogen cycles is therefore strongly dependent on the local mineralogical and chemical environment. Because carbonates are also potential oxidized carbon carriers, additional studies on the interaction between carbonates and FeOOH should be carried out in order to provide a comprehensive model for the deep-mantle carbon and water cycles. Nevertheless, the transformation reported here would prevent the production of FeHx, which is expected by the reaction of iron alloy from the core and hydrous phase at the core–mantle boundary [28,35,51]. This might have favored transfer of carbon to the core rather than of hydrogen during early Earth differentiation and therefore provide a mechanism for high amounts of C in an O-rich core [52]. Figure 5. View largeDownload slide P-T conditions at which the different phases have been observed experimentally: pyrite-structured FeO2 from Hu et al. [37], pyrite-structured FeO2Hx from this study as well as from Hu et al. [34] from both Fe2O3+H2O and FeOOH in Ar/Ne experiments, pyrite-structured FeOOH from Nishi et al. [35] and Fe4C3O12 from the present study as well as from FeO + CO2 experiments in Boulard et al. [15]. The mantle geotherms from [59] (O(08)) and [60] (A(82)), as well as hypothetical P-T paths for a ‘very cold slab’, ‘cold slab’ and a ‘hot slab’ [45,61] are also represented for comparison. Figure 5. View largeDownload slide P-T conditions at which the different phases have been observed experimentally: pyrite-structured FeO2 from Hu et al. [37], pyrite-structured FeO2Hx from this study as well as from Hu et al. [34] from both Fe2O3+H2O and FeOOH in Ar/Ne experiments, pyrite-structured FeOOH from Nishi et al. [35] and Fe4C3O12 from the present study as well as from FeO + CO2 experiments in Boulard et al. [15]. The mantle geotherms from [59] (O(08)) and [60] (A(82)), as well as hypothetical P-T paths for a ‘very cold slab’, ‘cold slab’ and a ‘hot slab’ [45,61] are also represented for comparison. MATERIALS AND METHODS Experiments were conducted using symmetric Mao-Bell-type DAC equipped with 300/100-μm beveled culet diamonds and rhenium gasket with 30-μm starting diameter hole and 25-μm starting thickness. Natural sample of crystalline goethite (α-FeOOH) from lateritic soil in Central African Republic was provided by the collection of University Pierre et Marie Curie. FeOOH was loaded in CO2 together with a ruby ball (see Supplementary Fig. 1 for an XRD characterization before the experiment). The CO2 gas was loaded using high-pressure gas-loading apparatus at room temperature and 600 bars. FeOOH was isolated from diamonds by CO2, preventing reactions with diamonds. In situ angle-dispersive XRD measurements were performed on the high-pressure beamline ID-27 at the European Synchrotron Radiation Facility (ESRF) using a monochromatic incident X-ray beam of 0.3738-Å wavelength. Before the experiments, the X-ray spot, spectrometer entrance and the heating laser spot were carefully aligned. The sample was first pressurized to its target pressure (107 GPa) before laser heating. Pressure was measured using ruby fluorescence before and after laser heating at room temperature by the use of a blue laser [53]. XRD peaks from the Re-gasket collected across the diamond culet gave pressures between 101 and 107 GPa. Two YAG lasers with excellent power stability were aligned on both sides of the sample, which produce hot spots larger than 20-μm (FWHM) diameter. Temperatures were obtained by fitting the sample thermal emission spectrum from the central 2 × 2 μm2 of the hotspot to the Planck's function using the wavelength range 600–900 nm. Reflective lenses were used for measurement in order to prevent any chromatic aberration [54]. The monochromatic X-ray beam was focused to 3 × 3 micron. This is smaller than the laser heating spot in order to reduce both the radial and axial temperature gradients. Typical exposure time was 30 s at high pressures and high temperatures. The diffraction images were integrated with the Fit2d software [57]. The 1D diffraction patterns were treated with the General Structure Analysis System (GSAS) software package [58] using the Rietveld or Lebail methods to identify the different phases and refine lattice parameters. During heating and XRD acquisition, temperature was measured continuously. Temperature uncertainties are estimated to be of about 150 K [55]. At high temperatures, thermal pressure corrections are of the order of +10–15% of the initial pressure [56]. Raman spectra were collected at high pressure and ambient temperature after transformation of the sample. We used a Jobin–Yvon® HR-460 spectrometer with monochromator with 1500 gratings/mm, equipped with an Andor® CCD camera. Raman signal was excited using the 514.5-nm wavelength of an Ar+ laser, delivering 300 mW focused into a 2-μm spot by a long-working distance Mitutoyo® x20 objective. A focused ion beam (FIB) thin section was extracted from the recovered sample at the center of laser-heated spot and thinned to electron transparency (∼100-nm thickness). FIB milling was performed using a FEI STRATA DB 235 at IEMN (Lille, France) with a focused Ga+ ion beam operating at 30 kV and currents from 20 nA down to 1 pA for final surfacing. Analytical transmission electron microscopy (ATEM) was carried out on the FIB thin section with a JEOL 2100-F operating at 200 keV, equipped with a field emission gun. Semi-quantitative chemical analyses on the individual phases was obtained by X-ray energy dispersive spectrometry (XEDS) and SAED patterns were used for phase identification. SUPPLEMENTARY DATA Supplementary data are available at NSR online. Acknowledgements We acknowledge the European Synchrotron Radiation Facility for the allocation of beam time. A.C. acknowledges a research sabbatical leave support from the French National Center for Scientific Research (CNRS). The transmission electron microscopy facility at Institut de Minéralogie, Physique des Matériaux et Cosmochimie is supported by Région Ile de France grant SESAME 2000 E 1435. The manuscript was significantly improved by the constructive comments from three anonymous reviewers. REFERENCES 1. Kelemen PB , Manning CE . 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National Science ReviewOxford University Press

Published: Mar 15, 2018

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