Tetracarbonate melts and the fate of primordial carbon in the deep Earth

Much of Earth’s carbon is thought to have been stripped away from the silicate mantle by dense metallic-iron to form the core 1 . However, recent studies 2,3 suggest that a considerable part of it could have remained stranded in the deep mantle due to a change in its anity to dissolve into iron metal-alloys at the extreme pressures and temperatures of the deep Earth. The underlying physical phenomena that would render carbon less siderophile at extreme conditions remain elusive. Here we describe the compaction mechanisms and structural evolution of a simple carbonate glass to deep mantle pressures by monitoring the evolution of the electronic state and atomic structure of the glass upon compression. Our new experiments demonstrate a pressure-induced change in hybridization of carbon from sp 2 to sp 3 starting at 40 GPa, due to the conversion of [3] CO 32- groups into [4] CO 44- units, which is completed at ~112 GPa. The pressure-induced increase of carbon coordination number from three to four increases possibilities for carbon-oxygen interactions with lower mantle silicates and increased compatibility 4,5 . Tetracarbonate melts provide a mechanism for changing the presumed siderophile nature of deep carbon and instead imply storage of carbon in the deep mantle as a possible source for carbon-rich emissions registered at the surface in intra-plate and near-ridge hot spots 6,7

mantle boundary (CMB). Our results provide experimental evidence for the change in D C in the Earth's lower mantle. We used a combination of synchrotron-based techniques, non-resonant inelastic X-ray scattering (XRS) at the oxygen K-edge and X-ray diffraction via pair distribution function (PDF) analysis, to monitor the evolution of the electronic structure and bond lengths upon compression of the glass (Methods). The interpretation of the experimental data is supported by ab initio Molecular Dynamic (AIMD) simulations and spectrum calculations based on solving the Bethe-Salpeter equation (Methods). Figure 1 shows the oxygen K-edge XRS spectra for K 2 Mg(CO 3 ) 2 at pressures ranging from 1 atm (Fig. 1A bottom) to 100 GPa (Fig. 1A top). The peak near 534 eV, more pronounced at lower pressures, re ects the antibonding molecular orbital of the resonant C=O double bond associated with the CO 3 planar unit, also known as π*-resonance 16 . At higher energies, above 538 eV, a second broader feature indicated as σ* corresponds to the unoccupied C-O σ* antibonding orbitals. The most prominent spectral change in the oxygen K-edge with pressure is the disappearance of the π*-resonance at 534 eV (Fig. 1A,B), unambiguous evidence of the breakage of the C=O double bond. The integrated intensity of the π*feature as a function of pressure in Fig. 1C shows that this process occurs gradually. After an initial decrease between ambient conditions and 10 GPa, we observe a plateau in π*-spectral intensity and a subsequent almost linear decay starting at ~40 GPa until the highest pressure, extrapolation suggests a complete disappearance at ca. 112 GPa. Besides these prominent changes in the spectral shape, we observe a shift of spectral weight visible by a smooth intensity increase in the σ*-resonance at pressures above 80 GPa, evidenced as the emergence of a peak at ~536 eV in the spectral-difference plot (Fig. 1B). We associate the shift to general compaction around the O atoms. Simulated XRS spectra based on atomic structural snapshots from AIMD simulations are shown as thin black lines in Fig. 1A, supporting our assessment of the formation of sp 3 -hybridized [4] CO 4 units at the expense of [3] CO 3 (Extended Data  Table 1). Above 85 GPa the C-O peak is also affected by a remarkable broadening at the FWHM, which shows that C-O distances at 104 GPa vary between 0.98 Å to 1.42 Å, ( at ~85 GPa more than 50% of the CO 3 -groups have been converted to CO 4 -groups, con rming the conclusions drawn from the XRS results and AIMD simulations (Extended Data Fig. 6). In the same pressure interval, the Mg-O bond expands to 1.98 Å (Extended Data Fig. 8, Table 1), which suggests an increase of coordination environment to 7+ at 104 GPa, from the initial ~4.5 at 1 bar 18 . K-O bonds show an almost linear compression, with a kink on the compression rate at 46 GPa, which coincides with the onset of formation of [4] CO 4 -groups. C-O, Mg-O and K-O distances extracted from the AIMD simulations con rm these compression trends (Extended Data Fig. 9, Supplementary Text 1).
Overall, the behavior of the carbonate glass at high pressures can be rationalized as follows (Supplementary Text 1): 1. 0 -40/45 GPa: lling of voids (mainly the rst 10 GPa), compaction with local atomic rearrangement and distortions.
3. >85 GPa : a fully polymerized framework structure with some local remnants of CO 3 -groups, up to ca.
Carbon distribution in the deep Earth is expected to depend fundamentally on the physicochemical properties of carbon-complexes in melts along with their buoyancy. While most of the primordial carbon has been stripped away from the primordial mantle during core formation, new estimates predict carbon concentration between 0.09(4) and 0.20 (10) wt% C in the core 3 , much lower than the 1 wt% previously assessed from cosmochemical and geochemical considerations 19 . Even with the lowest estimates, the core would still account for more than half of Earth's total carbon budget 3,20 , the other half being in the mantle. Today's reservoirs re ect the mobility of carbon during early Earth history, when a hotter and more vigorous mantle mixing was possible 21 . Our results indicate that a large amount of the C present at depths > 1200 km could have stayed there, forming C-rich silicate melts in virtue of its ability to polymerize with (CO 4 ) 4molecules or higher order oxy-carbon complexes 4 . The presence of oxidized Cspecies such as (CO 3 ) 2or (CO 4 ) 4groups in silicate melts even at reducing oxygen fugacity (fO 2 ) conditions is predicted by simulations 4,5 with carbon atoms having the ability to bond directly to Si (Si-C), C (C-C) or with O forming a number of different ionic complexes 4 . Carbonated-silicate melts are also predicted as a result of redox melting processes in the reduced mantle at the CMB 22 . Moreover, the reducing fO 2 in the primordial MO was shown to increase over time as a consequence of delivering increased oxidized material from the outer disk and the redox reaction SiO 2 + 2Fe à 2FeO + Si tending towards the right as pressures and temperatures increase (4 and refs in therein). Carbon could have also participated as a redox buffer and stabilized oxy-carbon complexes.
The increased polymerization has a strong effect on density and melt rheology. Density estimates indicate that adding up to 5 wt% of CO 2 in a MgSiO 3 melt at CMB conditions does not change the melt density su ciently to make it buoyant with respect to the surrounding mantle 5 . Hence, silicate melts could e ciently sequester signi cant amounts of carbon and simultaneously preserve the local mechanical stability since they have been found to be denser than the surrounding environment 5,14 . Thus, it is conceivable that the carbon present in the lowermost part of the lower mantle would have reached mechanical stability as a constituent of dense silicate melts and remained con ned there.
Deep primordial carbon reservoirs could explain distinct geochemical and geophysical signatures observed at depths and at the surface. Seismic images reveal the presence of two massive anomalous zones above the core, commonly known as large low velocity provinces (LLVPs), where seismic waves travel slowly. These provinces occupy an area of ~30 % of the CMB and are considered to be chemically distinct from the surrounding mantle, hence potentially formed by primordial or subducted material 23 (Supplementary Text 2, 3). LLVPs exhibit geographical correlation with hotspot volcanism at the surface, which commonly overlie the edges of the LLVPs, in locations known as ultralow-velocity zones (ULVZ) [23][24][25] (Fig. 3). Recent tomographic work suggests the presence of broad conduits of reduced wave velocities beneath some prominent hotspots, extending from ULVZ at the CMB to roughly ~1000 km depth 26 . The total carbon emissions (e.g. magmatic eruptions and degassing) at these hotspots, e.g. Hawaii, Iceland, and Samoa, are considered to have parental magmatic CO 2 content of the source ranging between 3000 and 10000 ppm, whereas mid-ocean-ridge (MORBs), which principally sample the convecting upper mantle, around 600 ppm 6 . Despite the high possible uncertainties in these estimates, hotspot mantle CO 2 concentrations are regarded as distinctly higher than mantle concentrations in all MORB segments located > 1000 km from hotspots 7 . Thus, we suggest that carbon enrichment in the deep lower mantle is possible as localized viscous carbonated silicate melts-and carbonatitic melts. These melts might be stripped away by deep upwelling material to become incorporated in rising plumes 23,27 , from as deep as the CMB, possibly constituting the parental source of intraplate oceanic volcanism and related ocean island basalts (OIBs). Upon decompression and depolymerization these C-rich oxidized melts become increasingly less viscous and more buoyant. Once exhumed to depths < 1200 km mechanical factors, such as low melt density, high buoyancy and mantle convection, contribute to ascent these carbon-rich melts into the upper mantle and ultimately their degassing into the atmosphere (Fig. 3).
The change in the speciation of the oxy-carbon complexes that we measured at extreme pressure affects carbon partitioning in the Earth. Here, we bring compelling evidence for the mechanism behind the enhanced lithophile character of carbon at extreme conditions. This suggests that the deep lower mantle may be the largest reservoir for the carbon cycle connected to well-known surface processes.

Additional information
Extended data is available for this paper.
Supplementary information is available for this paper.
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Sample selection and preparation
Pure carbonate glass (Extended Data Fig. 1) with composition K 2 Mg(CO 3 ) 2 was prepared at the Bayerisches Geoinstitut (BGI) by melting a homogenous and stoichiometric (50:50) mixture of K 2 CO 3 and MgCO 3 powder in an externally heated rapid-quench cold-seal pressure vessel, at 1 kbar and 700 ˚C for 12 h. The experiment was terminated by rapid in-situ quenching. High purity (> 99.9 %) K 2 CO 3 was purchased from Alfa Aesar GmbH, dried prior to preparation of the mixture at 110 ˚C and stored in a drying oven. High purity MgCO 3 was kindly provided by the Natural History Museum of London (UK), and dried at 200 ˚C to ensure dehydration of the Mg carbonate hydrates.

X-ray Raman Scattering
The high pressure X-ray Raman scattering (XRS) spectroscopy experiments (non-resonant inelastic x-ray scattering) were performed at beamline ID20 of the ESRF -The European Synchrotron Radiation Facility (Grenoble, France) using the large solid angle spectrometer 28 . We used radiation from four consecutive U26 undulators that was monochromatized by a succession of a high heat-load Si(111) pre-and a Si(311) channel-cut post-monochromator. The monochromatic x-rays were focused onto a spot size of 10x20 μm 2 at the sample position using a Kirkpatrick-Baez (KB) mirror system. The opening angles of the DAC allowed us to use spherically bent Si(660) analyzer crystals: six in the forward and six in the backscattering geometry, respectively. Signals from the forward and backscattering arrays were checked for consistency and averaged for momentum transfers of 2.3±0.1 Å -1 in forward scattering and 9.5±0.1 Å -1 in backscattering. The incident energy was scanned between values of 10.20 and 10.24 keV in order to obtain energy losses in the vicinity of the oxygen K-edge. All averaged signals were area normalized in the loss-region between 528.0 and 557.0 eV. In all experiments we used BX90 DACs 29 equipped with 0.5 mm thick miniature diamonds 30 and 200 μm and 120 μm diamonds culet size. All diamonds had a 15 μm deep recess drilled in the middle of the culet in order to maximize the x-ray path length through the sample inside the pressure chamber for an increased signal. All data analysis was performed using the XRStools program package 31,32 as described in Weis et al. 33 . The experimental pressure conditions were estimated o ine before and after each XRS measurement based on the Raman signal of the diamonds 34 , with the reported pressures being an average of the two values, and their difference was used as the error estimate.

X-ray Diffraction and Pair Distribution Function
The high-pressure x-ray diffraction experiments were performed at beamlines ID15a 35 and ID15b of the ESRF. Membrane cells with 120 μm culets size provided by the ESRF Sample Environment Service-HP lab were used in both experiments. For each pressure point a new ~50 µm-diameter bead of K 2 Mg(CO 3 ) 2 was placed in a wide-aperture membrane cell, tted with Re gaskets and no pressure medium. For each measurement, both sample and background signals were measured separately and used to calculate the structure factors S(Q), which were obtained by subtraction of the background from the sample patterns (Extended Data Fig. 2). Background measurements were taken on the DAC containing only the empty gasket after decompression back to ambient conditions, thus maintaining the same gasket thickness as the sample measurement at high pressure. Incident monochromatic x-ray beams with energies of 60 keV (λ=0.2066 Å) and 33 keV (λ=0.4126 Å) were used for the experiments performed at ID15a and ID15b, respectively. The x-ray beam was focused down to 3x3 μm 2 at ID15a using a Kirkpatrick-Baez mirror system (KB) and to 10x10 μm 2 at ID15b using Compound Refractive Lenses (CRLs). Diffraction data were collected with a Dectris Pilatus3 X CdTe 2M detector at ID15a and a MAR555 image plate detector at ID15b. Pilatus3X and MAR555 detectors were placed at 250 mm and 260 mm from the sample. A 20 µm pinhole was mounted at the entrance of the DAC to cut background air scattering. The detectorsample distance and the detector tilt angle were calibrated using LaB 6 and silicon standards for ID15a and ID15b, respectively. Contrary to the MAR555 detector, which was centered with respect to the transmitted direct beam, the Pilatus detector was off-centered, allowing a higher Q-range to be detected. The typical collection time was of ~8 minutes at ID15a and 30 minutes at ID15b. A comparison example between structure factors measured at ID15a and ID15b at ~30 GPa are displayed in Extended Data Fig.   3. As one can notice, the two curves nearly coincide, and most importantly, the onset of the rst and second sharp diffraction peaks, FSDP and SSDP respectively, coincide.
Because of the larger high quality Q-range obtained from structure factors measured at ID15a (larger detector placed off of the sample center, higher energy), we decided to calculate the x-ray Pair Distribution Functions (PDFs) only from those data. The structure factors collected at ID15b were used, together with those from ID15a, to perform analyses on the FSDP and SSDP, of the S(Q) (Extended Data Fig. 4 and 5). Data for PDF analysis were collected at each pressure in 128 exposures of 4 s each. Particular care was taken during cell loadings to ensure reproducible positions of the diamond Bragg re ections on the detector between sample and background measurements. For each sample dataset the median image was calculated and we subtracted the corresponding median background image. Azimuthal integration to yield I(Q) diffraction patterns was done using pyFAI 36 . From each I(Q) the PDF, g(r), was calculated using PdfGetX3 37 . The range of Q for g(r) calculations was 0.8≤Q≤12 Å -1 . In order to obtain a reliable PDF from weak, noisy signal, ad-hoc intensity weighting and smoothing of the Q-space input functions were used to remove artifacts such as e.g. ripples, associated with parasitic scatting, particularly generated by the tails of diamond peaks. While peak intensities might not be very precise due to the di culties of estimating against an unknown background, i.e. consequence of DAC setup, the peak positions are considered to be correct. Thus, we explain the distinguished intensity drop of the C-O peak of the PDF at 104 GPa (Fig. 2) caused by the PDF generation process.

Ab initio molecular dynamics simulations
First-principles molecular dynamics (FPMD) calculations were performed using the Vienna Ab initio Simulation Package (VASP) 38 based on the density functional theory. Effects of the core electrons were replaced by atomic projected augmented wave (PAW) potentials 39 . Valence electronic con gurations of 3s 2 3p 6 4s 1 , 3s2 3p 6 4s 2 , 2s 2 2p 2 and 2s 2 2p 4 were used for K, Mg, C and O atoms, respectively. The Perdew-Burke-Ernzerhof (PBE) exchange-correlation functional 40 was used. The wave functions were represented by a plane-wave basis set with an energy cutoff of 495 eV. To generate the glass structure of K 2 Mg(CO 3 ) 2 at ambient pressure and ambient temperature, a supercell containing 176 atoms (16 K 2 Mg(CO 3 ) 2 formula units) was initially melted at 2000 K for 30 ps with a time step of 1 fs in the canonical ensemble (NVT). The Nose-Hoover thermostat was used to control the temperature. A single point was used for Brillouin zone sampling. Then the molten structure was gradually cooled to 300 K in 10 ps and followed by a NVT MD run at 300 K for at least another 30 ps. Higher pressure glass structures were obtained by isotropically reducing the cell size step by step. Due to the limited cell size and time duration of the MD calculation, over-heating is usually needed for the occurrence of the structural transformation in such closed system. Thus, the high pressure structure (above 28 GPa) were annealed at 4000 K for 30 ps and then quenched to 300 K. To eliminate anisotropy in the stress tensors from the NVT calculation, the isothermal-isobaric ensemble (NPT calculation) was employed on the quenched glass structures for an additional 20 ps. In agreement with experimental observations, structural analysis of the atomic con gurations show 4-fold C-O coordination starting to appear at pressures higher than 58 GPa. The amount of 3-and 4-coordinated C-O are almost equal at 106 GPa (Extended Data Fig. 6). At 145 GPa, the highest pressure considered here, 4-coordinated C-O becomes dominant. Snapshots extracted from the NPT MD trajectories were used in the computation of the XRS spectra (Fig. 1A).   Distribution and cycling of carbon in the deep mantle during core formation (left) and in today's Earth (right). Carbon becomes less siderophile with depth and as a consequence is stranded in the lowermost lower mantle as tetracoordinated carbon in silicate melts. With the solidi cation and convective motions, primordial carbon was accumulated in speci c regions at the CMB26 known as LLVPs and ULVZs. The latter are believed to be the original source of hotspot volcanisms23,26, thus primordial carbon trapped at the CMB could be the primary source of carbon emissions in these locations (Hawaii, Samoa, Pitcairn and many others). Image credit: Josh Wood.

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