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We present thermodynamic and empirical calculations for the iron-wustite (IW) buffer applicable from 100 kPa to 100 GPa and from 1000 to 3000 K. The thermodynamic calculation self-consistently accounts for changing stoichiometry of iron-saturated wustite as a function of temperature and pressure. In contrast to some previous models for calculating IW at high pressure, the model incorporates a thermodynamically valid representation of the free energy of stoichiometric FeO at 100 kPa. Earlier high pressure models that relied on the JANAF thermochemical tables ( Chase, 1998 ) were compromised because JANAF has erroneous values for the properties of FeO. This resulted in predicted oxygen fugacities buffered by IW that are between 0.2 and 1.1 log units too reducing at 3000 and 1000 K, respectively. The revised thermodynamic calculations indicate that iron-saturated wustite becomes more nearly stoichiometric with increasing pressure, but that this shift depends on temperature. Near-stoichiometric FeO (y

To better understand the solution of volatile species in a reduced magma ocean, we identify via Raman spectroscopy the nature of C-O-H-N volatile species dissolved in a series of reduced basaltic glasses. The oxygen fugacity (ƒO2) during synthesis varied from highly reduced at two log units below the iron-wustite buffer (IW-2.1) to moderately reduced (IW-0.4), spanning much of the magmatic ƒO2 conditions during late stages of terrestrial accretion. Raman vibrational modes for H2, NH2–, NH3, CH4, CO, CN–, N2, and OH– species are inferred from band assignments in all reduced glasses. The integrated area of Raman bands assigned to N2, CH4, NH3 and H2 vibrations in glasses increases with increasing molar volume of the melt, whereas that of CO decreases. Additionally, with increasing ƒO2, CO band areas increase while those of N2 decrease, suggesting that the solubility of these neutral molecules is not solely determined by the melt molar volume under reduced conditions. Coexisting with these neutral molecules, other species as CN–, NH2– and OH– are chemically bonded within the silicate network. The observations indicate that, under reduced conditions, (1) H2, NH2–, NH3, CH4, CO, CN–, N2, and OH– species coexist in silicate glasses representative of silicate liquids in a magma ocean (2) their relative abundances dissolved in a magma ocean depend on melt composition, ƒO2 and the availability of H and, (3) metal-silicate partitioning or degassing reactions of those magmatic volatile species must involve changes in melt and vapor speciation, which in turn may influence isotopic fractionation.

To better understand the role of sulfide in C storage in the upper mantle, we construct a thermodynamic model for Fe-Ni-S-C sulfide melts and consider equilibrium between sulfide melts, mantle silicates, Fe-Ni alloy, and diamond. The sulfide melt model is based upon previous parameterization of Fe-Ni-S melts calibrated at 100 kPa, which we have extended to high pressure based on volumetric properties of end-member components. We calculate the behavior of C in the sulfide melt from empirical parameterization of experimental C solubility data. We calculate the continuous compositional evolution of Fe-Ni sulfide liquid and associated effects on carbon storage at pressure and redox conditions corresponding to mantle depths of 60 to 410 km. Equilibrium and mass balance conditions were solved for coexisting Fe-Ni-S melt and silicate minerals (olivine [(Mg,Fe,Ni)2SiO4], pyroxene [(Mg,Fe)SiO3]) in a mantle with 200 ppmw S. With increasing depth and decreasing oxygen fugacity (fO2), the calculated melt (Fe+Ni)/S atomic ratio increases from 0.8–1.5 in the shallow oxidized mantle to 2.0–10.5 in the reduced deep upper mantle (>8 GPa), with Fe-Ni alloy saturation occurring at >10 GPa. Compared to previous calculations for the reduced deep upper mantle, alloy saturation occurs at greater depth owing to the capacity of sulfide melt to dissolve metal species, thereby attenuating the rise of Fe and Ni metal activities. The corresponding carbon storage capacity in the metal-rich sulfide liquid rises from negligible below 6 GPa to 8-20 ppmw at 9 GPa, and thence increases sharply to 90-110 ppmw at the point of alloy saturation at 10-12 GPa. The combined C storage capacity of liquid and solid alloy reaches 110-170 ppmw at 14 GPa. Thus, in the deep upper mantle, all carbon in depleted sources (10–30 ppmw C) can be stored in the sulfide liquid, and alloy and sulfide liquids host a significant fraction of the C in enriched sources (30–500 ppmw C). Application of these results to the occurrences of inferred metal-rich sulfide melts in the Fe-Ni-S-C system and inclusions in diamonds from the mantle transition zone suggests that oxidization of a reduced metal-rich sulfide melt is an efficient mechanism for deep-mantle diamond precipitation, owing to the strong effect of (Fe+Ni)/S ratio on carbon solubility in Fe-Ni-S melts. This redox reaction likely occurs near the boundary between oxidized subducted slabs and the reduced ambient peridotitic mantle.

International audience; “Cousin” worlds—slightly bigger or slightly hotter than Earth—can help us understand planetary habitability, but we need more lab and numerical experiments to make the most of this opportunity.

During the formation of terrestrial planets, volatile loss may occur through nebular processing, planetesimal differentiation, and planetary accretion. We investigate iron meteorites as an archive of volatile loss during planetesimal processing. The carbon contents of the parent bodies of magmatic iron meteorites are reconstructed by thermodynamic modelling. Calculated solid/molten alloy partitioning of C increases greatly with liquid S concentration and inferred parent body C concentrations range from 0.0004 to 0.11 wt.%. Parent bodies fall into 2 compositional clusters characterized by cores with medium, and low C/S. Both of these require significant planetesimal degassing, as metamorphic devolatilization on chondrite-like precursors is insufficient to account for their C depletions. Planetesimal core formation models, ranging from closed system extraction to degassing of a wholly molten body, show that significant open system silicate melting and volatile loss is required to match medium and low C/S parent body core compositions. Greater depletion in C relative to S is the hallmark of silicate degassing, indicating that parent body core compositions record processes that affect composite silicate/iron planetesimals. Degassing of bare cores stripped of their silicate mantles would deplete S with negligible C loss, and could not account for inferred parent body core compositions. Devolatilization during small-body differentiation is thus a key process in shaping the volatile inventory of terrestrial planets derived from planetesimals and planetary embryos., 50 pages including supplementary materials

Carbon is an essential element for life but its behavior during Earth's accretion is not well understood. Carbonaceous grains in meteoritic and cometary materials suggest that irreversible sublimation, and not condensation, governs carbon acquisition by terrestrial worlds. Through astronomical observations and modeling we show that the sublimation front of carbon carriers in the solar nebula, or the soot line, moved inward quickly so that carbon-rich ingredients would be available for accretion at 1 au after the first million years. On the other hand, geological constraints firmly establish a severe carbon deficit in Earth, requiring the destruction of inherited carbonaceous organics in the majority of its building blocks. The carbon-poor nature of the Earth thus implies carbon loss in its precursor material through sublimation within the first million years., Comment: 21 pages including main article and supplementary materials

We conducted experiments at high pressure (P) and temperature (T) to measure hydrogen solubility in plagioclase (Pl) with a range of compositions (An15 to An94). Experiments were run at 700–850 °C, 0.5 GPa, and f O 2 close to either the Ni-NiO (NNO) or iron-wustite (IW) oxygen buffers. Experiments at 700 °C on An15 (containing 0.03 wt% FeO) reveal no dependence of H solubility on f O 2 between IW and NNO, but experiments at 800–850 °C on other compositions (with 0.3–0.5 wt% FeO) demonstrate that H solubility is enhanced by a factor of ∼2–3 at IW compared to NNO, consistent with previous experiments by Yang (2012a) on An58. By analogy with synthetic hydrogen feldspar (HAlSi3O8), we infer that the predominant mechanism for H incorporation in Pl is through bonding to O atoms adjacent to M-site vacancies, and we propose likely O sites for H incorporation based on M O bond lengths in anhydrous Pl structures. Increased uptake of structurally bound H at low f O 2 is explained by the formation of defect associates resulting from the reduction of Fe3+ in tetrahedral sites to Fe2+, allowing additional H to be incorporated in adjacent M-site vacancies. This mechanism counteracts the expected effect of water fugacity on H solubility. We also speculate on possible substitutions of H on tetrahedral vacancies, as well as coupled H-F substitution. Enhanced incorporation of H in Pl at low f O 2 may have implications for estimating the water content of the lunar magma ocean. However, mechanisms unrelated to low f O 2 are needed to explain high H contents in terrestrial Pl xenocrysts, such as those found in basalts from the Basin and Range.

Crucial monitors of operation of deep Earth volatile cycles are the accumulated volatile masses in the exosphere, particularly when these are compared to outgassing fluxes and to the masses stored in the mantle. Further insight can be gained by examining the relative mantle/exosphere fractionations of the major (H2O, C, and N) volatiles. New estimates for the H2O and C contents of the convecting mantle (290 ± 80 and 110 ± 40 ppm, respectively) are derived from H2O/Ce and CO2/Ba ratios of oceanic basalts combined with convecting mantle Ce and Ba concentrations. Together with an earlier estimate of convecting mantle N (1.1 ± 0.55 ppm) and with surface inventories (including a new estimate for total exosphere C, 10.6 ± 1.8 × 10 22 g ), these allow construction of exosphere-normalized mantle masses of major volatiles. Values for H2O, C, and N are, respectively 0.75 ± 0.2, 4.2 ± 2.0 and 0.7 ± 0.35, meaning subequal amounts of H2O and N are in the mantle and exosphere, but most of the bulk silicate Earth (BSE) C is in the mantle. Outfluxes of major volatiles from the mantle suggest exosphere replenishment times of 7.5 and 1 Ga for H2O and C. Previous estimates for outfluxes of N indicate a replenishment time of 80 Ga, but an alternative based on the C/N ratio of the depleted mantle is 16 Ga. Importantly, the normalized flux of C from the mantle to the exosphere is much greater than those for H2O and N, even though the exosphere C reservoir is by far the smallest of the three. This is owing to more efficient recycling of C relative to H2O and N, where “recycling” means return of materials from near-surface reservoirs to the convecting mantle, and/or to large surface reservoirs of H2O and N (but not C) inherited from Earth's early history. H2O/Ce, CO2/Nb, and CO2/Ba ratios are little-fractionated from one another during mantle melting, but the H2O/Ce ratio of the exosphere (1540 ± 360) is much greater than mantle ratios (200 ± 50), indicating that H2O is not recycled as efficiently as Ce or partial preservation of a large primordial ocean. In contrast, the exosphere CO2/Ba ratio (40 ± 14) is far smaller than the convecting mantle ratio (100 ± 20), indicating that C is recycled more efficiently than Ba. The small exosphere CO2/Ba ratio reflects substantial long-term Ba storage in the continental crust combined with significant recycling to the mantle of C. Monte Carlo simulations of C and Ba exchange between the interior and exterior reservoirs, taking into account uncertainties in total C and Ba in the BSE and in the exosphere, suggest that for 0–40% Ba recycling to the mantle, 35–80% of the C that has outgassed to the surface through time has been returned to the deep mantle. Some former surface C may be stored in the continental lithosphere, but observed xenolith C concentrations indicate that the lithosphere accommodates only a small fraction of C apparently returned to the mantle. Consideration of models in which the continental lithosphere stores substantial C has only small effect on the quantitative conclusions of this exercise.

To investigate the effect of metal/sulfide and Ni/Fe ratio on the C storage capacity of sulfide melts, we determine carbon solubility in Fe-Ni-S melts with various (Fe + Ni)/S and Ni/Fe via graphite-saturated high-pressure experiments from 2–7 GPa and 1200–1600 °C. Consistent with previous results, C solubility is high (4–6 wt.%) in metal-rich sulfide melts and diminishes with increasing S content. Melts with near M/S = 1 (XS > 0.4) have

To improve the accuracy of X-ray absorption near-edge structure (XANES) calibrations for the Fe3 +/ΣFe ratio in basaltic glasses, we reevaluated the Fe3 +/ΣFe ratios of glasses used as standards by Cottrell et al. (2009), and available to the community (NMNH catalog #117393). Here we take into account the effect of recoilless fraction on the apparent Fe3 +/ΣFe ratio measured from room temperature Mossbauer spectra in that original study. Recoilless fractions were determined from Mossbauer spectra collected from 40 to 320 K for one basaltic glass, AII_25, and from spectra acquired at 10 K for the 13 basaltic glass standards from the study of Cottrell et al. (2009). The recoilless fractions, f, of Fe2 + and Fe3 + in glass AII_25 were calculated from variable-temperature Mossbauer spectra by a relative method (RM), based on the temperature dependence of the absorption area ratios of Fe3 + and Fe2 + paramagnetic doublets. The resulting correction factor applicable to room temperature determinations (C293, the ratio of recoilless fractions for Fe3 + and Fe2 +) is 1.125 ± 0.068 (2σ). Comparison of the spectra at 10 K for the 13 basaltic glasses with those from 293 K suggests C293 equal to 1.105 ± 0.015 (2σ). Although the 10 K estimate is more precise, the relative method determination is believed to be more accurate, as it does not depend on the assumption that recoilless fractions are equal at 10 K. Applying the effects of recoilless fraction to the relationship between Mossbauer-determined Fe3 +/ΣFe ratios and revised average XANES pre-edge centroids for the 13 standard glasses allows regression of a new calibration of the relationship between the Fe XANES pre-edge centroid energy and the Fe3 +/ΣFe ratio of silicate glass. We also update the calibration of Zhang et al. (2016) for andesites and present a more general calibration for mafic glasses including both basaltic and andesitic compositions. Recalculation of Fe3 +/ΣFe ratios for the mid-ocean ridge basalt (MORB) glasses analyzed previously by XANES by Cottrell and Kelley (2011) results in an average Fe3 +/ΣFe ratio for MORB of 0.143 ± 0.008 (1σ), taking into account only analytical precision, and 0.14 ± 0.01(1σ), taking into account uncertainty on the value of C293. This revised average is lower than the average of 0.16 ± 0.01 given by Cottrell and Kelley (2011). The revised average oxygen fugacity for MORB based on the database of Cottrell and Kelley (2011) is − 0.18 ± 0.16 log units less than the quartz-fayalite-magnetite buffer of Frost (1991) at 100 kPa (∆ QFM = − 0.18 ± 0.16).

Pyroxene is the principal host of Fe3+ in basalt source regions, hosting 79 and 81% of the Fe3+ in spinel and garnet lherzolite, respectively. In spinel peridotite, orthopyroxene (opx) and clinopyroxene (cpx) host 48% and 31%, respectively, of the total Fe3+. Yet the relationship between mantle mineralogy, pyroxene chemistry, and the oxygen fugacity (fO2) recorded by mantle-derived basalts remains unclear. To better understand partitioning of Fe3+ between pyroxene and melt we conducted experiments at 100 kPa with fO2 controlled by CO-CO2 gas mixes between ∆QFM −1.19 to +2.06 in a system containing andesitic melt saturated with opx or cpx only. To produce large (100–150 μm), homogeneous pyroxenes, we employed a dynamic cooling technique with a 5–10 °C/h cooling rate, and initial and final dwell temperatures 5–10 °C and 20–30 °C super and sub-liquidus, respectively. Resulting pyroxene crystals have absolute variation in Al2O3 and TiO2

The Earth's deep interior contains significant reservoirs of volatiles such as H, C, and N. Due to the incompatible nature of these volatile species, it has been difficult to reconcile their storage in the residual mantle immediately following crystallization of the terrestrial magma ocean (MO). As the magma ocean freezes, it is commonly assumed that very small amounts of melt are retained in the residual mantle, limiting the trapped volatile concentration in the primordial mantle. In this article, we show that inefficient melt drainage out of the freezing front can retain large amounts of volatiles hosted in the trapped melt in the residual mantle while creating a thick early atmosphere. Using a two-phase flow model, we demonstrate that compaction within the moving freezing front is inefficient over time scales characteristic of magma ocean solidification. We employ a scaling relation between the trapped melt fraction, the rate of compaction, and the rate of freezing in our magma ocean evolution model. For cosmochemically plausible fractions of volatiles delivered during the later stages of accretion, our calculations suggest that up to 77% of total H2O and 12% of CO2 could have been trapped in the mantle during magma ocean crystallization. The assumption of a constant trapped melt fraction underestimates the mass of volatiles in the residual mantle by more than an order of magnitude.

Although partial melt in the asthenosphere is important geodynamically, geophysical constraints on its abundance remain ambiguous. We use a database of seamounts detected using satellite altimetry to constrain the temporal history of erupted asthenospheric melt. We find that intraplate volcanism on young seafloor (

Experiments establishing the effect of pressure on the Fe3+/ΣFe ratio of andesitic silicate melts buffered by coexisting Ru and RuO2 were performed from 100 kPa to 7 GPa and 1400–1750 °C. Fe3+/ΣFe ratios were determined by room temperature Mossbauer spectroscopy, but corrected for the effects of recoilless fraction. Fe3+/ΣFe ratios in quenched glasses decrease with increasing pressure consistent with previous results between 100 kPa and 3 GPa (O’Neill et al., 2006), but show only small pressure effects above 5 GPa. Ratios also decrease with increasing temperature. Mossbauer hyperfine parameters indicate mean coordination of Fe3+ ions of ∼5 in glasses, with no dependence on the pressure from which the glasses were quenched, but show an increase with pressure in mean coordination of Fe2+ ions, from ∼5 to ∼6. XANES spectra on these glasses show variations in pre-edge intensities and centroid positions that are systematic with Fe3+/ΣFe, but are displaced from those established from otherwise identical andesitic glasses quenched at 100 kPa (Zhang et al., 2016). These systematics permit construction of a new XANES calibration curve relating pre-edge sub-peak intensities to Fe3+/ΣFe applicable to high pressure glasses. Consistent with interpretations of the Mossbauer hyperfine parameters, XANES pre-edge peak features in high pressure glasses are owing chiefly to the effects of pressure on the coordination of Fe2+ ions from ∼5.5 to ∼6, with negligible effects evident for Fe3+ ions. We use the new data to construct a thermodynamic model relating the effects of oxygen fugacity and pressure on Fe3+/ΣFe. We apply this model to calculate variations in oxygen fugacity in isochemical (constant Fe3+/ΣFe) columns of magma representative of magma oceans, in which fO2 is fixed at the base by equilibration with molten Fe. These calculations indicate that oxygen fugacities at the surface of shallow magma oceans are more reduced than at depth. For magma oceans in which the pressure at the base is near 5 GPa, as may be appropriate for Mercury and the Moon, conditions at the surface are ∼1.5 log unit more reduced at the surface than at their base. If the results calibrated up to pressures of 7 GPa can be extrapolated to higher pressures appropriate for magma oceans on larger terrestrial planets such as Mars or Earth, then conditions at the surface are ∼2 or 2.5 log units more reduced at the surface than at the base, respectively. Thus, atmospheres overlying shallow magma oceans should be highly reduced and rich in H2 and CO.

One of the most remarkable observations regarding volatile elements in the solar system is the depletion of N in the bulk silicate Earth (BSE) relative to chondrites, leading to a particularly high and non-chondritic C:N ratio. The N depletion may reflect large-scale differentiation events such as sequestration in Earth's core or massive blow off of Earth's early atmosphere, or alternatively the characteristics of a late-added volatile-rich veneer. As the behavior of N during early planetary differentiation processes is poorly constrained, we determined together the partitioning of N and C between Fe–N–C metal alloy and two different silicate melts (a terrestrial and a martian basalt). Conditions spanned a range of fO2 from ΔIW−0.4 to ΔIW−3.5 at 1.2 to 3 GPa, and 1400 °C or 1600 °C, where ΔIW is the logarithmic difference between experimental fO2 and that imposed by the coexistence of crystalline Fe and wustite. N partitioning ( D N metal / silicate ) depends chiefly on fO2, decreasing from 24 ± 3 to 0.3 ± 0.1 with decreasing fO2. D N metal / silicate also decreases with increasing temperature and pressure at similar fO2, though the effect is subordinate. In contrast, C partition coefficients ( D C metal / silicate ) show no evidence of a pressure dependence but diminish with temperature. At 1400 °C, D C metal / silicate partition coefficients increase linearly with decreasing fO2 from 300 ± 30 to 670 ± 50 . At 1600 °C, however, they increase from ΔIW−0.7 to ΔIW−2 ( 87 ± 3 to 240 ± 50 ) and decrease from ΔIW−2 to ΔIW−3.3 ( 99 ± 6 ) . Enhanced C in melts at high temperatures under reduced conditions may reflect stabilization of C–H species (most likely CH4). No significant compositional dependence for either N or C partitioning is evident, perhaps owing to the comparatively similar basalts investigated. At modestly reduced conditions (ΔIW−0.4 to −2.2), N is more compatible in core-forming metal than in molten silicate ( 1 ≤ D N metal / silicate ≤ 24 ), while at more reduced conditions (ΔIW−2.2 to ΔIW−3.5), N becomes more compatible in the magma ocean than in the metal phase. In contrast, C is highly siderophile at all conditions investigated ( 100 ≤ D C metal / silicate ≤ 700 ). Therefore, sequestration of volatiles in the core affects C more than N, and lowers the C:N ratio of the BSE. Consequently, the N depletion and the high C:N ratio of the BSE cannot be explained by core formation. Mass balance modeling suggests that core formation combined with atmosphere blow-off also cannot produce a non-metallic Earth with a C:N ratio similar to the BSE, but that the accretion of a C-rich late veneer can account for the observed high BSE C:N ratio.

A quantitative understanding of nitrogen incorporation in Earth materials is important for constraining volatile evolution in planetary bodies. We used a combination of chemical (SIMS, EPMA, and laser-extraction mass spectrometry) and spectroscopic (FTIR) observations to study nitrogen contents and speciation mechanisms in silicate glasses, metal alloys, and an N-bearing silicate mineral (hyalophane). One suite of Fe-free basaltic glasses was studied by all four methods. Concentrations of N in these glasses determined by EPMA are systematically higher than those measured by laser extraction but agree within mutual 2s uncertainties, demonstrating the general veracity of the EPMA method. SIMS working curves based on measurement of ^(14)N^+ and ^(14)N^(16)O^- as a function of N content determined by EPMA (or laser extraction) are best fit with exponential functions rather than the linear regressions that are most commonly applied to SIMS data. On the other hand, the relationship based on ^(12)C^(14)N- for C-poor, Fe-free glasses is exceptionally well fit to a linear regression (r^2 = 1, p < 0.001), in contrast to expectations from previous work on glasses with lower N contents. Matrix effects on the SIMS signals associated with Fe or H_2O content are not justified by the data, but volatile data (both N and H) for hyalophane, which contains 20 wt% BaO, reveal matrix effects possibly induced by its high average molar mass. A combination of FTIR and chemical data, together with a thorough review of the literature, was used to determine incorporation mechanisms for N in the Fe-free glasses. We infer that under reducing conditions at high pressure and temperature N is dissolved in basaltic melts chiefly as NH^−_2) and NH^(2–), with N_2 and/or nitride (X-N_3^–) complexes becoming increasingly important at low f_(O_2), increasing N content, and decreasing H content. Our results have implications for future studies seeking to accurately measure N by SIMS and for studies of N partitioning at high pressure relevant to planetary accretion and differentiation.

Andesitic glasses equilibrated at 1350 °C over a range of oxygen fugacities (log f O 2 from − 8.63 to − 0.68) were examined with Fe K-edge X-ray absorption near-edge structure (XANES) and Mossbauer spectra. XANES spectral features were then calibrated as a function of Mossbauer-derived Fe 3+ /∑Fe ratios. Additionally, both methods help characterize the local structure of iron ions in andesitic glasses. Fe 3+ /∑Fe ratios were determined from Mossbauer spectra collected at room temperature but corrected with recoilless fractions obtained from previously reported Mossbauer data collected on one of the glasses from 47 to 293 K. An empirical model was derived for the correlation between the pre-edge centroid energy and Fe 3+ /∑Fe ratio for andesitic glasses. This trend is intermediate between those previously determined for rhyolitic and basaltic glasses, but the distinction from basaltic compositions may be owing chiefly to differences in calibrations for Fe 3+ /∑Fe ratio, rather than to intrinsic differences in the spectra as a function of Fe 3+ /∑Fe ratio for mafic glasses. The ratios of intensities of pre-edge sub-peaks and Fe 3+ /∑Fe ratios for andesitic, basaltic, and rhyolitic glasses plot along a common trend, indicating that these measures provide a XANES calibration for Fe 3+ /∑Fe ratio that is less dependent on silicate composition. The coordination numbers of Fe 3+ and Fe 2+ ions in andesitic glass can be calculated from observations of pre-edge centroid energies and total intensities, combined with independent constraints on Fe 3+ /∑Fe ratio from Mossbauer spectra. The mean coordination of Fe 2+ ions calculated this way is close to 5.5 for reduced and oxidized compositions, and this is consistent with inferences from hyperfine features of the Mossbauer spectra. The mean coordination number of Fe 3+ inferred from XANES increases from ~ 4.5 to ~ 5 as andesitic glasses vary from reduced to oxidized; Mossbauer hyperfine parameters also suggest network-forming behavior of Fe 3+ , but with higher coordination for more reduced glasses.

Earth’s inventory of principle volatiles C, H, N, and S is a legacy of its early stages of accretion and differentiation. Elemental ratios (C/H, C/N, C/S) are powerful tools for understanding early processing of Earth’s volatiles, as they monitor relative fractionations through important processes even when absolute concentrations are less well defined. The C/H ratio of the bulk silicate Earth (BSE), defined from surface reservoirs and minimally degassed oceanic basalts is 1.3 ± 0.3, which is 5–15 times lower than the C/H ratio of carbonaceous and enstatite chondrites and 2–5 times lower than ordinary chondrites. The BSE C/N ratio is superchondritic (40 ± 8; Bergin et al. 2015) while the C/S ratio (0.49 ± 0.14) is nearly chondritic. Successful models of volatile acquisition and processing must account for the effects of accretion, core formation, and atmospheric loss on all three of these ratios. Simple models of equilibration between a magma ocean, the overlying atmosphere, and alloy destined for the core are used to explore the influence of core formation and atmospheric loss on major volatile concentrations and ratios. Among major volatile elements, C is most siderophile, and consequently core formation leaves behind a non-metallic Earth with low C/H, C/N, and C/S ratios compared to originally accreted materials and compared to the BSE. Compared to the predicted effect of early differentiation, the relatively high C/X ratios of the BSE argue in part that significant volatile replenishment occurred after core formation ceased, possibly in the form of a late veneer. However, a late veneer with chondritic composition is insufficient to explain the pattern of major volatile enrichments and depletions because BSE C/H and C/N ratios are non-chondritic. The C/H ratio is best explained if an appreciable fraction of H in the BSE predates delivery in the late veneer. Although atmospheric blow-off is an attractive explanation for the high C/N ratio, available data for C and N solubility and metal/silicate partitioning suggest that atmospheric blow-off cannot counter core formation to produce subchondritic C/N. Thus, unless virtually all core-forming metal segregated prior to volatile accretion (or relative C and N solubilities are appreciably different from those assumed here), the BSE C/N ratio suggests that accreting materials had elevated ratios compared to carbonaceous chondrites. One possibility is that a fraction of Earth’s volatiles accreted from differentiated C-rich planetesimals similar to the ureilite parent body. Reconciling C/H, C/N, and C/S ratios of the BSE simultaneously presents a major challenge that almost certainly involves a combination of parent body processing, core formation, catastrophic atmospheric loss, and partial replenishment by a late veneer. The chondritic C/S ratio of the BSE and relatively low S content of the BSE constrains the BSE C concentration, but a potential complicating factor in interpreting the BSE C/S ratio is the possible effect of segregation of an S-rich matte to the core during the later parts of core-mantle differentiation.

We present high-pressure experiments up to 8 GPa that constrain the solidus and liquidus of a composition, Fe 0.69 Ni 0.23 Cu 0.01 S 1.00 , typical of upper mantle sulfide. Solidus and liquidus brackets of this monosulfide are parameterized according to a relation similar to the Simon-Glatzel equation, yielding, respectively, T (°C) = 1015.1 [ P (GPa)/1.88 + 1] 0.206 and T (°C) = 1067.3 [ P (GPa)/1.19 + 1] 0.149 (1 ≤ P ≤ 8). The solidus fit is accurate within ±15 °C over the pressure intervals 1–3.5 GPa and within ±30 °C over the pressure intervals 3.5–8.0 GPa. The solidus of the material examined is cooler than the geotherm for convecting mantle, but hotter than typical continental geotherms, suggesting that sulfide is molten or partially molten through much of the convecting upper mantle, but potentially solid in the continental mantle. However, the material examined is one of the more refractory among the spectrum of natural mantle sulfide compositions. This, together with the solidus-lowering effects of O and C not constrained by the present experiments, indicates that the experimentally derived melting curves are upper bounds on sulfide melting in the Earth’s upper mantle and that the regions where sulfide is molten are likely extensive in both the convecting upper mantle and, potentially, the deeper parts of the oceanic and continental lithosphere, including common source regions of many diamonds.

The behavior of nickel in the Earth’s mantle is controlled by sulfide melt–olivine reaction. Prior to this study, experiments were carried out at low pressures with narrow range of Ni/Fe in sulfide melt. As the mantle becomes more reduced with depth, experiments at comparable conditions provide an assessment of the effect of pressure at low-oxygen fugacity conditions. In this study, we constrain the Fe–Ni composition of molten sulfide in the Earth’s upper mantle via sulfide melt–olivine reaction experiments at 2 GPa, 1200 and 1400 °C, with sulfide melt $$X_{{{\text{Ni}}}}^{{{\text{Sulfide}}}}=\frac{{{\text{Ni}}}}{{{\text{Ni}}+{\text{Fe}}}}$$ (atomic ratio) ranging from 0 to 0.94. To verify the approach to equilibrium and to explore the effect of $${f_{{{\text{O}}_{\text{2}}}}}$$ on Fe–Ni exchange between phases, four different suites of experiments were conducted, varying in their experimental geometry and initial composition. Effects of Ni secondary fluorescence on olivine analyses were corrected using the PENELOPE algorithm (Baro et al., Nucl Instrum Methods Phys Res B 100:31–46, 1995), “zero time” experiments, and measurements before and after dissolution of surrounding sulfides. Oxygen fugacities in the experiments, estimated from the measured O contents of sulfide melts and from the compositions of coexisting olivines, were 3.0 ± 1.0 log units more reduced than the fayalite–magnetite-quartz (FMQ) buffer (suite 1, 2 and 3), and FMQ − 1 or more oxidized (suite 4). For the reduced (suites 1–3) experiments, Fe–Ni distribution coefficients $$K_{{\text{D}}}^{{}}=\frac{{(X_{{{\text{Ni}}}}^{{{\text{sulfide}}}}/X_{{{\text{Fe}}}}^{{{\text{sulfide}}}})}}{{(X_{{{\text{Ni}}}}^{{{\text{olivine}}}}/X_{{{\text{Fe}}}}^{{{\text{olivine}}}})}}$$ are small, averaging 10.0 ± 5.7, with little variation as a function of total Ni content. More oxidized experiments (suite 4) give larger values of KD (21.1–25.2). Compared to previous determinations at 100 kPa, values of KD from this study are chiefly lower, in large part owing to the more reduced conditions of the experiments. The observed difference does not seem attributable to differences in temperature and pressure between experimental studies. It may be related in part to the effects of metal/sulfur ratio in sulfide melt. Application of these results to the composition of molten sulfide in peridotite indicates that compositions are intermediate in composition ( $$X_{{{\text{Ni}}}}^{{{\text{sulfide}}}}$$ ~ 0.4–0.6) in the shallow mantle at 50 km, becomes more Ni rich with depth as the O content of the melt diminishes, reaching a maximum (0.6–0.7) at depths near 80–120 km, and then becomes more Fe rich in the deeper mantle where conditions are more reduced, approaching ( $$X_{{{\text{Ni}}}}^{{{\text{sulfide}}}}$$ ~ 0.28) > 140 km depth. Because Ni-rich sulfide in the shallow upper mantle melts at lower temperature than more Fe-rich compositions, mantle sulfide is likely molten in much of the deep continental lithosphere, including regions of diamond formation.

Using vibrational spectroscopy and SIMS, we determined the solubility and speciation of C–O–H–N dissolved volatiles in mafic glasses quenched from high pressure under reduced conditions, with f O 2 from −3.65 to +1.46 relative to the iron–wustite buffer (IW). Experiments were performed on martian and terrestrial basalts at 1.2 GPa and 1400 °C in graphite containers with variable availability of H 2 O, and in the presence of FePt alloys or Fe–C liquids. The dominant C–O–H–N species varies systematically with f O 2 and H 2 O content: the carbonate ion prevails above IW + 1, but for dry conditions between IW−2 and IW + 1, C O species are most important. Below IW, reduced NH-bearing species are present. At the most reducing and hydrous (∼0.5 wt% H 2 O) conditions, small amounts of CH 4 are present. Concentrations of C diminish as conditions become more reduced, amounting to 10 s to ∼100 ppm in the interval ∼IW−2 to IW + 1 where C O species dominate, and as little as 1–3 ppm at more reduced conditions. Concentrations of non-carbonate carbon, dominated by C O species, correlate with CO fugacities along a trend implying that the species stoichiometry has just one C O group and suggesting that carbonyl complexes (transition metals with multiple carbon monoxide ligands) are not important species under these conditions. C partition coefficients between Fe–C liquid and silicate melt increase with decreasing f O 2 , becoming as great as 10 4 for the most reducing conditions investigated. The low solubility of C in silicate liquids under reducing conditions means that most C during the magma ocean stage of planetary differentiation is either segregated to the core or in the overlying atmosphere. Precipitation of C-rich phases in a carbon-saturated magma ocean is also possible, and is one mechanism by which some C can be retained in the mantle of a planet. The predominant magmatic carbonaceous species for both martian and lunar volcanism is likely C O.

To evaluate the accuracy of Fe3+ and Fe2+ ratios in silicate glasses determined by Mossbauer spectroscopy, we examine in detail the temperature (47–293 K) of Mossbauer spectra for two andesitic glasses, one quenched at 1 atm, 1400 °C (VF3) and the other at 3.5 GPa, 1600 °C (M544). Variable-temperature Mossbauer spectra of these two glasses are used to characterize the recoilless fraction, f , by two different methods—a relative method (RM) based on the temperature dependence of the ratios of Fe3+ and Fe2+ Mossbauer doublets and the second based on the temperature dependence of the center shift (CS) of the doublets. The ratio of the recoilless fractions for Fe3+ and Fe2+, C T, can then be used to adjust the observed area of the Mossbauer doublets into the Fe3+/∑Fe ratio in the sample. We also evaluated the contributions of non-paramagnetic components to the Fe in the glasses by determining the influence of applied magnetic field on sample magnetization. Finally, for the VF3 glass, we determined the Fe3+/∑Fe independently by wet chemical determination of the FeO content combined with careful electron microprobe analyses of total Fe. Recoilless fractions determined with the CS method (CSM) are significantly smaller than those determined with the relative method and suggest larger corrections to room-temperature Fe3+/∑Fe ratios. However, the RM determinations are believed to be more accurate because they depend less on the assumption of the Debye harmonic model and because they produce more nearly temperature-independent estimates of Fe3+/∑Fe ratios. Non-linear responses of sample magnetizations to applied magnetic fields indicate that the glasses contain a small (0.4–1.1% for VF3) superparamagnetic component that is most likely to be nanophase precipitates of (Fe,Mg)Fe2O4 oxide, but corrections for this component have negligible influence on the total Fe3+/∑Fe determined for the glass. For the VF3 glass, the Fe3+/∑Fe produced by uncorrected room-temperature Mossbauer spectroscopy [0.685 ± 0.014 in two standard deviation (2σ)] agrees within 3% of that determined by wet chemistry (0.666 ± 0.030 in 2σ). The Fe3+/∑Fe corrected for recoilless fraction contributions is 0.634 ± 0.078(2σ), which is 7.5% lower than the uncorrected room-temperature ratio, but also agrees within 5% of wet chemical ratio. At least for this andesitic glass, the room-temperature determination of Fe3+/∑Fe is accurate within analytical uncertainty, but room-temperature Mossbauer determinations of Fe3+/∑Fe are always systematically higher compared to recoilless-fraction corrected ratios.

To determine partitioning of C between upper mantle silicate minerals and basaltic melts, we executed 26 experiments between 0.8 and 3 GPa and 1250–1500 °C which yielded 37 mineral/glass pairs suitable for C analysis by secondary ion mass spectrometry (SIMS). To enhance detection limits, experiments were conducted with 13C-enriched bulk compositions. Independent measurements of 13C and 12C in coexisting phases produced two C partition coefficients for each mineral pair and allowed assessment of the approach to equilibrium during each experiment. Concentrations of C in olivine (ol), orthopyroxene (opx), clinopyroxene (cpx) and garnet (gt) range from 0.2 to 3.5 ppm, and resulting C partition coefficients for ol/melt, opx/melt, cpx/melt and gt/melt are, respectively, 0.0007 ± 0.0004 ( n = 2 ), 0.0003 ± 0.0002 ( n = 45 ), 0.0005 ± 0.0004 ( n = 17 ) and 0.0001 ± 0.00007 ( n = 5 ). The effective partition coefficient of C during partial melting of peridotite is 0.00055 ± 0.00025 , and therefore C is significantly more incompatible than Nb, slightly more compatible than Ba, and, among refractory trace elements, most similar in behavior to U or Th. Experiments also yielded partition coefficients for F and H between minerals and melts. Combining new and previous values of D F mineral / melt yields bulk D F peridotite / melt = 0.011 ± 0.002 , which suggests that F behaves similarly to La during partial melting of peridotite. Values of D H pyx / melt correlate with tetrahedral Al along a trend consistent with previously published determinations. Small-degree partial melting of the mantle results in considerable CO2/Nb fractionation, which is likely the cause of high CO2/Nb evident in some Nb-rich oceanic basalts. CO2/Ba is much less easily fractionated, with incompatible-element-enriched partial melts having lower CO2/Ba than less enriched basalts. Comparison of calculated behavior of CO2, Nb, and Ba to systematics of oceanic basalts suggests that depleted (DMM-like) sources have 75 ± 25 ppm CO2 (CO2/Nb = 505 ± 168, CO2/Ba = 133 ± 44), whereas enriched sources of intraplate basalts similar in concentrations to primitive mantle have 600 ± 200 ppm CO2. If all mantle reservoirs are expressed in the current inventory of oceanic basalts for which nearly undegassed CO2 concentrations are available, then we estimate the likely range of mantle C concentrations to be 1.4–4.8 × 1023 grams of C, or 1.5–5.2 times the mass of the current C surface reservoir. Depending on the assumed Ba and Nb contents of average oceanic crust, resulting ridge fluxes of C range from 7.2 × 1013 to 2.9 × 1014 g/yr.

Journal of Petrology, 45 (12), ISSN:0022-3530, ISSN:1460-2415

To determine the speciation and concentrations of dissolved COH volatiles in graphite-saturated martian primitive magmas, we conducted piston-cylinder experiments on graphite-encapsulated synthetic melt of Adirondack-class Humphrey basaltic composition. Experiments were performed over three orders of magnitude in oxygen fugacity (IW+2.3 to IW−0.8), and at pressures (1–3.2 GPa) and temperatures (1340–1617 °C) similar to those of possible martian source regions. Oxygen fugacities were determined from compositions of coexisting silicate melt + FePt alloy, olivine + pyroxene + FePt alloy, or melt + FeC liquid. Infrared spectra of quenched glasses all show carbonate absorptions at 1430 and 1520 cm−1, with CO2 concentrations diminishing under more reduced conditions, from 0.50 wt% down to 26 ppm. Carbon contents of silicate glasses and FeC liquids were measured using secondary ion mass spectrometry (SIMS) yielding 36–716 ppm and 6.71–7.03 wt%, respectively. Fourier transform infrared (FTIR) and SIMS analysis produced similar H2O contents of 0.26–0.85 and 0.29–0.40 wt%, respectively. Raman spectra of glasses reveal evidence for OH− ions, but no indication of methane-related species. FTIR-measured concentrations of dissolved carbonate diminish linearly with oxygen fugacity, but more reduced conditions yield greater dissolved carbonate concentrations than would be expected based on oxidized conditions in previous work. C contents of silicate glasses determined by SIMS are consistently higher than C as carbonate determined by FTIR. Their difference, termed non-carbonate C, correlates well with additional IR absorptions found in reduced glasses (fO2 < IW+0.4) at 1615, 2205, and 3370 cm−1. These absorption bands are not seen in more oxidized glasses, except B441 (IW+1.7), presumably because they represent reduced C-bearing complexes. The 2205 cm−1 peak is attributed to a CO complex, possibly an Fe-carbonyl ion. The 1615 cm−1 peak does not correlate with that at 2205 cm−1, but does correlate with non-carbonate C and is in a region commonly associated with CO bonding. The origin of the peak at 3370 cm−1 is poorly understood and could potentially be owing to a variety of COH species or to NH bonding. The intensities of the 1615 and 3370 cm−1 peaks correlate with each other leading us to provisionally attribute both to an unspecified complex with both CO and NH bonds. These results suggest that dissolved species such as carbonyl or other CO-bearing species could be a significant source of C fluxes to the martian atmosphere, with minor additions of CO2 and negligible methane contributions. By assuming that degassed, reduced C ultimately becomes atmospheric CO2, reduced C outgassing may be incorporated in models of martian atmospheric evolution. At Humphrey source region conditions (1350 ± 50 °C, 1.2 ± 0.1 GPa) the total C contents are equivalent to 1200 ppm CO2 at IW+1 and 475 ppm CO2 at IW, which are 2 and 4 times higher than the CO2 derived from CO32− alone. For reasonable magmatic fluxes over the last 4.5 Ga of martian history, such graphite-saturated magmas would produce 0.25 and 0.60 bars from sources at IW and IW+1, significantly more than expected solely from consideration of dissolved CO2. The carbon contents of FeC liquids in this study are consistent with graphite-saturated carbide liquids becoming more C-rich with increasing temperature. Experiments with melt and FeC liquid have values of DCall/sil between 1.3 × 103 and 2.2 × 103, potentially allowing planetary mantles to retain significant C following core formation.

The Moon-forming giant impact extensively melts and partially vaporizes the silicate Earth and delivers a substantial mass of metal to Earth's core. Subsequent evolution of the magma ocean and overlying atmosphere has been described by theoretical models but observable constraints on this epoch have proved elusive. Here, we report calculations of the primordial atmosphere during the magma ocean and water ocean epochs and forge new links with observations to gain insight into the behavior of volatiles on the early Earth. As Earth's magma ocean crystallizes, it outgasses the bulk of the volatiles into the primordial atmosphere. The redox state of the magma ocean controls both the chemical composition of the outgassed volatiles and the hydrogen isotopic composition of water oceans that remain after hydrogen loss from the primordial atmosphere. Whereas water condenses and is retained, molecular hydrogen does not condense and can escape, allowing large quantities (~10^2 bars) of hydrogen - if present - to be lost from Earth in this epoch. Because the escaping inventory of H can be comparable to the hydrogen inventory in the early oceans, the corresponding deuterium enrichment can be large with a magnitude that depends on the initial H2 inventory. By contrast, the common view that terrestrial water has a carbonaceous chondrite source requires the oceans to preserve the isotopic composition of that source, undergoing minimal D-enrichment via H2 loss. Such minimal enrichment places upper limits on the amount of primordial H2 in contact with early water oceans (pH2, 39 pages, 6 figures; including supplementary materials

Enrichment in K2O in oceanic island basalts (OIB) is correlated with high SiO2, low CaO/Al2O3, and radiogenic isotopic signatures indicative of enriched mantle sources (EM1 and EM2). These are also chemical characteristics of the petit-spot lavas, which are highly enriched in K2O (3–4 wt%) compared to other primitive oceanic basalts. We present experimentally derived liquids with varying concentrations of K2O in equilibrium with a garnet lherzolite residue at 3 GPa to test the hypothesis that the major element characteristics of EM-type basalts are related to their enrichment in K2O. SiO2 is known to increase with K2O at pressures less than 3 GPa, but it was previously unknown if this effect was significant at the high pressures associated with partial melting at the base of the lithosphere. We find that at 3 GPa for each 1 wt% increase in the K2O content of a garnet lherzolite saturated melt, SiO2 increases by ~0.5 wt% and CaO decreases by ~0.5 wt%. MgO and $$K_{D}^{{{\text{Fe}} - {\text{Mg}}}}$$ each decrease slightly with K2O concentration, as do Na2O and Cr2O3. The effect of K2O alone is not strong enough to account for the SiO2 and CaO signatures associated with high-K2O OIB. The SiO2, CaO, and K2O concentrations of experimentally derived partial melts presented here resemble those of petit-spot lavas, but the Al2O3 concentrations from the experimental melts are greater. Partitioning of K2O between peridotite and melt suggests that petit spots, previously considered to sample ambient asthenosphere, require a source more enriched in K2O than the MORB source.

Ratios of first-row transition elements (FRTE), such as Fe/Mn and Zn/Fe, may be fractionated differently by partial melting of peridotite than by partial melting of recycled lithologies like eclogite, and therefore may be useful as indicators of the source lithologies of mantle-derived basalts. Interpretation of basalt source lithologies from FRTE ratios requires accurate assessment of FRTE partitioning behavior between peridotitic minerals and coexisting melts. We present experimental determinations of partition coefficients for several of the FRTE (Sc, Ti, V, Cr, Mn, Fe, Co, Zn) and Ga and Ge between basaltic melt and olivine, garnet, pyroxenes, and spinel at 3 GPa. Because mineral/melt partitioning is sensitive to phase compositions, a key feature of these experiments is that the melts and minerals are known from previous experiments to be in equilibrium at the solidus of garnet peridotite at 3 GPa. Therefore, these partition coefficients are directly applicable to near-solidus partial melting of the mantle at 3 GPa. We use these partition coefficients to calculate compositions of model partial melts of peridotite and compare these to natural OIB. Model partial melts of peridotite have lower Fe/Mn ( 7 * 10−4) than many primitive OIB, which implies that some other source lithology participates in the formation of many OIB. Alternatively, these ratios may potentially be produced by garnet peridotite if the source contains ∼0.3% Fe2O3, consistent with observations from continental xenoliths. Zn/Fe is a less sensitive indicator of non-peridotite source lithology than either Fe/Mn or Co/Fe, as Zn/Fe in partial melts of peridotite overlaps with >75% of primitive OIB. Ga and Sc are fractionated significantly by residual garnet, and high Ga/Sc may indicate the presence of garnet in basalt source regions. When taking into account several FRTE ratios simultaneously, few OIB appear to be consistent with derivation solely from a reduced peridotitic source. The source either must have a modest non-peridotitic component, be Fe-enriched, or be slightly oxidized.

The onset of melting in the Earth's upper mantle influences the thermal evolution of the planet, fluxes of key volatiles to the exosphere, and geochemical and geophysical properties of the mantle. Although carbonatitic melt could be stable 250 km or less beneath mid-oceanic ridges, owing to the small fraction (∼0.03 wt%) its effects on the mantle properties are unclear. Geophysical measurements, however, suggest that melts of greater volume may be present at ∼200 km (refs 3-5) but large melt fractions are thought to be restricted to shallower depths. Here we present experiments on carbonated peridotites over 2-5 GPa that constrain the location and the slope of the onset of silicate melting in the mantle. We find that the pressure-temperature slope of carbonated silicate melting is steeper than the solidus of volatile-free peridotite and that silicate melting of dry peridotite + CO(2) beneath ridges commences at ∼180 km. Accounting for the effect of 50-200 p.p.m. H(2)O on freezing point depression, the onset of silicate melting for a sub-ridge mantle with ∼100 p.p.m. CO(2) becomes as deep as ∼220-300 km. We suggest that, on a global scale, carbonated silicate melt generation at a redox front ∼250-200 km deep, with destabilization of metal and majorite in the upwelling mantle, explains the oceanic low-velocity zone and the electrical conductivity structure of the mantle. In locally oxidized domains, deeper carbonated silicate melt may contribute to the seismic X-discontinuity. Furthermore, our results, along with the electrical conductivity of molten carbonated peridotite and that of the oceanic upper mantle, suggest that mantle at depth is CO(2)-rich but H(2)O-poor. Finally, carbonated silicate melts restrict the stability of carbonatite in the Earth's deep upper mantle, and the inventory of carbon, H(2)O and other highly incompatible elements at ridges becomes controlled by the flux of the former.

Deep-Earth volatile cycles couple the mantle with near-surface reservoirs. Volatiles are emitted by volcanism and, in particular, from mid-ocean ridges, which are the most prolific source of basaltic volcanism. Estimates of volatile extraction from the asthenosphere beneath ridges typically rely on measurements of undegassed lavas combined with simple petrogenetic models of the mean degree of melting. Estimated volatile fluxes have large uncertainties; this is partly due to a poor understanding of how volatiles are transported by magma in the asthenosphere. Here, we assess the fate of mantle volatiles through numerical simulations of melting and melt transport at mid-ocean ridges. Our simulations are based on two-phase, magma/mantle dynamics theory coupled to idealised thermodynamic model of mantle melting in the presence of water and carbon dioxide. We combine simulation results with catalogued observations of all ridge segments to estimate a range of likely volatile output from the global mid-ocean ridge system. We thus predict global MOR crust production of 66-73 Gt/yr (22-24 km3/yr) and global volatile output of 52-110 Mt/yr, corresponding to mantle volatile contents of 100--200~ppm. We find that volatile extraction is limited: up to half of deep, volatile-rich melt is not focused to the axis but is rather deposited along the LAB. As these distal melts crystallise and fractionate, they metasomatise the base of the lithosphere, creating rheological heterogeneity that could contribute to the seismic signature of the LAB., 42 pages; 8 figures; 2 appendices (incl 1 table); 7 suppl. figures; 1 suppl. table

To determine the influence of basalt composition on the CO2 solubility in martian lavas, we investigated experimentally a synthetic melt based on the martian meteorite Yamato 980459 (Y 980459), an olivine-phyric shergottite and a picritic rock (19 wt% MgO) thought to be a near-primary liquid derived from high-temperature (>1540 °C) partial melting of the martian mantle. Experiments were performed in a piston-cylinder apparatus at 1–2 GPa and 1600–1650 °C. CO2 contents in quenched glasses were determined using Fourier transform infrared spectroscopy (FTIR) and range from 0.45–1.26 wt%. Despite large differences in FeO* and MgO contents, the CO2 solubilities in Y 980459 are similar to that in a less primitive synthetic martian basalt based on the Humphrey rock and to a Hawaiian tholeiite. The lack of enhanced solubility in Fe2+- and Mg2+-rich melts is likely owing to the complex structural role of these cations in silicate melts, acting partly as network formers, rather than network modifiers. The small sensitivity of CO2 solubility to compositional variations among martian and tholeiitic basalts means that the experimentally determined solubilities may be applicable to a wide spectrum of martian magmatic products. Using experimentally determined CO2 solubilities of Y 980459 and Humphrey allows the calibration of the thermodynamic parameters governing dissolution of CO2 vapor as carbonate in martian basalts. This relation facilitates calculation of the CO2 dissolved in magmas derived from graphite-saturated martian basalt source regions as a function of P , T , and f O2. The hot conditions in the source of Y 980459, 1540 ± 10 °C, and 1.2 ± 0.5 GPa, are plausible for plume-related magmas forming the giant Tharsis volcanic complex, which accounts for 50% of martian igneous activity since stabilization of the primordial crust. If oxygen fugacity in the sources of hot Tharsis magmatism were equivalent to that at the iron-wustite buffer (IW) or 1 log unit above (IW+1), respectively, then the entire Tharsis event would outgas 30–300 mbars of CO2 to the martian atmosphere, which is far from the 2 bars required to stabilize an equable climate in the late Noachian and early Hesperian epochs. This mismatch could be reconciled if significant martian igneous activity derived from comparatively oxidized mantle sources (i.e., IW+2) similar to those responsible for the nakhlite meteorites.

The H2O storage capacities of peridotitic minerals place crucial constraints on the onset of hydrous partial melting in the mantle. The storage capacities of minerals in equilibrium with a peridotite mineral assemblage (“peridotite-saturated” minerals) are lower than when the minerals coexist only with fluid because hydrous partial melt is stabilized at a lower activity of H2O. Here, we determine peridotite-saturated olivine H2O storage capacities from 5 to 8 GPa and 1400–1500 °C in layered experiments designed to grow large (∼100–150 μm) olivine crystals in equilibrium with the full hydrous peridotite assemblage (melt+ol+opx+gar+cpx). The peridotite-saturated H2O storage capacity of olivine at 1450 °C rises from 57±26 ppm (by wt.) at 5 GPa to 254±60 ppm at 8 GPa. Combining these with results of a parallel study at 10–13 GPa ( Tenner et al., 2011 , CMP) yields a linear relation applicable from 5 to 13 GPa for peridotite-saturated H2O storage capacity of olivine at 1450 °C, C H 2 O olivine ( ppm ) = 57.6 ( ± 16 ) × P ( GPa ) − 169 ( ± 18 ) . Storage capacity diminishes with increasing temperature, but is unaffected by variable total H2O concentration between 0.47 and 1.0 wt%. Both of these are as predicted for the condition in which the water activity in the melt is governed principally by the cryoscopic requirement of melt stability for a given temperature below the dry solidus. Measured olivine storage capacities are in agreement or slightly greater than those predicted by a model that combines data from experimental freezing point depression and olivine/melt partition coefficients of H2O ( Hirschmann et al., 2009 ). Considering the temperature along the mantle geotherm, as well as available constraints on garnet/olivine and pyroxene/olivine partitioning of H2O ( D H 2 O gar / ol , D H 2 O px / ol ), we estimate the peridotite H2O storage capacity in the low velocity zone. The C H 2 O required to initiate melting between 150 and 250 km depth is between 270 and 855 ppm. We conclude that hydrous partial melting does not occur at these depths for H2O concentrations (50–200 ppm) typical of the convecting upper mantle sampled by mid-ocean ridge basalts.

We present experiments from 0.7 to 3 GPa that quantify solubility of H 2 in silicate melts under controlled hydrogen fugacities ( f H2 ). Two experimental series, one on synthetic basalt+COH and other with a synthetic andesite+OH, were conducted using a double capsule technique to impose a range of f H2 , on the samples. Quenched glasses were analyzed by FTIR and SIMS. Both series follow simple solubility laws in which molecular H 2 concentrations are proportional to f H2 and with a partial molar volume of molecular H 2 of 11 cm 3 /mole. Solubilities in andesitic melt are systematically greater than in basaltic liquid in a relationship consistent with control by the ionic porosity (IP) of the melts. Extrapolation based on IP allows estimation of the solubility of H 2 in peridotitic melts applicable to magma oceans. The H 2 /(H 2 +H 2 O) ratio in silicate melts (where H 2 O includes molecular H 2 O and OH − ) increases as conditions become more reduced, with increasing pressure, and with increasing total H. Under some conditions prevailing in the early Earth and terrestrial planets as well as in the deep Earth today, H 2 can be a significant fraction of the dissolved H and at high pressure it may exceed “water” (H 2 O and OH − ). Therefore, magmatic H 2 may influence the initial distribution of volatiles and the redox evolution of terrestrial planets, as well as the ongoing formation and fate of hydrous melts in the deep Earth today. Hydrous species in melts in equilibrium with Fe-rich alloy at high pressure, for example during core formation from a magma ocean, could be chiefly H 2 , rather than H 2 O. Hence, delivery of H 2 to the core by removal of Fe hydride need not be coupled to oxidation of the residual mantle. Although lunar basalts are much reduced, the fraction of H dissolved as molecular H 2 is small owing to low total H concentrations. Extrapolation to conditions of potential hydrous partial melting in the deep Earth suggests that the chief magmatic volatile may be H 2 rather than H 2 O. The very small partial specific density of magmatic H 2 (0.18 g/cm 3 at low pressure) may provide significant positive buoyancy to deep partial melts.

Redox conditions in magma oceans (MOs) have a key influence on the mass and composition of Earth's early atmosphere. If the shallow part of the MO is oxidized, it may be overlain by an H 2 O–CO 2 atmosphere, but if the near-surface magma is close to equilibrium with Fe-rich alloy, then the atmosphere will consist chiefly of H 2 , H 2 O, and CO, and on cooling will be rich in CH 4 . Although MOs are intimately associated with core-forming metal, the redox conditions in their shallow parts are not necessarily reducing. The magmatic Fe 3+ /Fe T ratio is set by equilibrium with metal at depth and homogenized through the magma column by convection. Indirect evidence suggests that the Fe 3+ /Fe T ratio of magmas in equilibrium with alloy at high pressure is greater than at low pressure, such that the shallow part of the MO may be comparatively oxidized and coexist with an atmosphere consisting chiefly of H 2 O and CO 2 . The mass of the atmosphere is dictated by the concentrations of volatile-species dissolved in the magma, which in turn are determined by partitioning between magma and alloy. Very strong partitioning of C into alloy may capture most of the carbon delivered to the growing planet, leaving behind a C-poor bulk silicate Earth (BSE) and a C-poor atmosphere. However, modest solubility of CH 4 in the magma may allow the BSE to retain significant C. Alternatively, if partitioning of C into alloy is extreme but the fraction of metal equilibrated with the MO is small, the alloy may become saturated with diamond. Floatation of diamond in the MO may retain a substantial inventory of C in the early mantle. BSE C may also have been replenished in a late veneer. Following segregation of metal to the core, crystallization of the MO may have prompted precipitation of C-rich phases (graphite, diamond, carbide), limiting the C in the early atmosphere and creating a substantial interior C inventory that may account for the large fraction of BSE carbon in the mantle today. Such precipitation could have occurred owing to a combination of the redox evolution of the crystallizing MO and cooling.

In the form of ionic impurities in rocks and minerals, water lubricates tectonic plates, influences rock viscosities and melting processes, and slows down seismic waves.

To understand possible volcanogenic fluxes of CO 2 to the Martian atmosphere, we investigated experimentally carbonate solubility in a synthetic melt based on the Adirondack-class Humphrey basalt at 1–2.5 GPa and 1400–1625 °C. Starting materials included both oxidized and reduced compositions, allowing a test of the effect of iron oxidation state on CO 2 solubility. CO 2 contents in experimental glasses were determined using Fourier transform infrared spectroscopy (FTIR) and Fe 3+ /Fe T was measured by Mossbauer spectroscopy. The CO 2 contents of glasses show no dependence on Fe 3+ /Fe T and range from 0.34 to 2.12 wt.%. For Humphrey basalt, analysis of glasses with gravimetrically-determined CO 2 contents allowed calibration of an integrated molar absorptivity of 81,500 ± 1500 L mol −1 cm −2 for the integrated area under the carbonate doublet at 1430 and 1520 cm −1 . The experimentally determined CO 2 solubilities allow calibration of the thermodynamic parameters governing dissolution of CO 2 vapor as carbonate in silicate melt, K II , ( Stolper and Holloway, 1988 ) as follows: ln K II 0 = - 15.42 ± 0.20 , Δ V 0 = 20.85 ± 0.91 cm 3 mol −1 , and Δ H 0 = −17.96 ± 10.2 kJ mol −1 . This relation, combined with the known thermodynamics of graphite oxidation, facilitates calculation of the CO 2 dissolved in magmas derived from graphite-saturated Martian basalt source regions as a function of P , T , and f O 2 . For the source region for Humphrey, constrained by phase equilibria to be near 1350 °C and 1.2 GPa, the resulting CO 2 contents are 51 ppm at the iron–wustite buffer (IW), and 510 ppm at one order of magnitude above IW (IW + 1). However, solubilities are expected to be greater for depolymerized partial melts similar to primitive shergottite Yamato 980459 (Y 980459). This, combined with hotter source temperatures (1540 °C and 1.2 GPa) could allow hot plume-like magmas similar to Y 980459 to dissolve 240 ppm CO 2 at IW and 0.24 wt.% of CO 2 at IW + 1. For expected magmatic fluxes over the last 4.5 Ga of Martian history, magmas similar to Humphrey would only produce 0.03 and 0.26 bars from sources at IW and IW + 1, respectively. On the other hand, more primitive magmas like Y 980459 could plausibly produce 0.12 and 1.2 bars at IW and IW + 1, respectively. Thus, if typical Martian volcanic activity was reduced and the melting conditions cool, then degassing of CO 2 to the atmosphere may not be sufficient to create greenhouse conditions required by observations of liquid surface water. However, if a significant fraction of Martian magmas derive from hot and primitive sources, as may have been true during the formation of Tharsis in the late Noachian, that are also slightly oxidized (IW + 1.2), then significant contribution of volcanogenic CO 2 to an early Martian greenhouse is plausible.

We investigate charge-balanced hydrous magnesium and silicon defects [(2H) X Mg, (4H) X Si] by first principles. Two new lowest-energy hydrogen configurations are proposed for (4H) X Si. With these new configurations, the distribution of O-H stretching phonon frequencies in Group I (>3450 cm –1 ) are better reproduced. Substitution of silicon with four hydrogen atoms gives rise to significant elongation of distances between O atoms at the tetrahedron of the silicon vacancy. Our calculations indicate that the correlation between O-O distances and O-H stretching phonon frequencies, which has been well established for hydrous minerals, does not apply directly to nominally anhydrous minerals and should not be used to determine the identity of hydrous defects responsible for infrared absorption peaks.

The onset of hydrous partial melting in the mantle above the transition zone is dictated by the H2O storage capacity of peridotite, which is defined as the maximum concentration that the solid assemblage can store at P and T without stabilizing a hydrous fluid or melt. H2O storage capacities of minerals in simple systems do not adequately constrain the peridotite water storage capacity because simpler systems do not account for enhanced hydrous melt stability and reduced H2O activity facilitated by the additional components of multiply saturated peridotite. In this study, we determine peridotite-saturated olivine and pyroxene water storage capacities at 10–13 GPa and 1,350–1,450°C by employing layered experiments, in which the bottom ~2/3 of the capsule consists of hydrated KLB-1 oxide analog peridotite and the top ~1/3 of the capsule is a nearly monomineralic layer of hydrated Mg# 89.6 olivine. This method facilitates the growth of ~200-μm olivine crystals, as well as accessory low-Ca pyroxenes up to ~50 μm in diameter. The presence of small amounts of hydrous melt ensures that crystalline phases have maximal H2O contents possible, while in equilibrium with the full peridotite assemblage (melt + ol + pyx + gt). At 12 GPa, olivine and pyroxene water storage capacities decrease from ~1,000 to 650 ppm, and ~1,400 to 1,100 ppm, respectively, as temperature increases from 1,350 to 1,450°C. Combining our results with those from a companion study at 5–8 GPa (Ardia et al., in prep.) at 1,450°C, the olivine water storage capacity increases linearly with increasing pressure and is defined by the relation $$ C_{{{\text{H}}_{2} {\text{O}}}}^{\text{olivine}} \left( {\text{ppm}} \right) = 57.6\left( { \pm 16} \right) \times P\left( {\text{GPa}} \right) - 169\left( { \pm 18} \right). $$ Adjustment of this trend for small increases in temperature along the mantle geotherm, combined with experimental determinations of $$ D_{{{\text{H}}_{2} {\text{O}}}}^{\text{pyx/olivine}} $$ from this study and estimates of $$ D_{{{\text{H}}_{2} {\text{O}}}}^{{{\text{gt}}/{\text{olivine}}}} $$ , allows for estimation of peridotite H2O storage capacity, which is 440 ± 200 ppm at 400 km. This suggests that MORB source upper mantle, which contains 50–200 ppm bulk H2O, is not wet enough to incite a global melt layer above the 410-km discontinuity. However, OIB source mantle and residues of subducted slabs, which contain 300–1,000 ppm bulk H2O, can exceed the peridotite H2O storage capacity and incite localized hydrous partial melting in the deep upper mantle. Experimentally determined values of $$ D_{{{\text{H}}_{2} {\text{O}}}}^{{{\text{pyx}}/{\text{olivine}}}} $$ at 10–13 GPa have a narrow range of 1.35 ± 0.13, meaning that olivine is probably the most important host of H2O in the deep upper mantle. The increase in hydration of olivine with depth in the upper mantle may have significant influence on viscosity and other transport properties.

To investigate the influence of chemical composition on the behavior of H 2 O in Fe-rich nominally anhydrous minerals, and to determine the difference between H 2 O behavior in the martian and terrestrial mantles, we conducted high-pressure H 2 O storage capacity experiments employing a wide range of olivine compositions. Experiments were conducted with bulk compositions in the system FeO-MgO-SiO 2 -H 2 O with Mg no. [Mg no. = 100 × molar Mg/(Mg+Fe)] ranging between 50 and 100 at 3 GPa in a piston-cylinder and at 6 GPa in a multi-anvil apparatus. Experiments at 3 GPa were conducted at 1200 °C, with f O 2 buffered by the coexistence of Fe and FeO, and at 1300–1500 °C in unbuffered assemblies. Experiments at 6 GPa were conducted at 1200 °C without buffers. Experiments at 1200 °C produced olivine+orthopyroxene+hydrous liquid (liq), and higher T experiments produced olivine+liq. Additionally, we synthesized a suite of 7 olivine standards (Mg no. = 90) for low blank secondary ion mass spectrometry (SIMS) analysis of H in multi-anvil experiments at 3–10 GPa and 1250 °C, resulting in large (200–400 μm) homogeneous crystals with 0.037 to 0.30 wt% H 2 O. Polarized Fourier transform infrared (FTIR) measurements on randomly oriented grains from the synthesis experiments were used to determine principal axis spectra through least-squares regression, and H contents were calculated from the total absorbance in the OH stretching region. Using these olivines as calibrants for SIMS analyses, the H contents of olivines and pyroxenes from the variable Mg no. experiments were measured by counting 16 OH ions. Ignoring any matrix effects owing to variation in Mg no., H contents of olivine and pyroxene increase linearly with decreasing Mg no. At 6 GPa and 1200 °C, olivine H contents increase from 0.05 to 0.13 wt% H 2 O (8360 to 23 900 H/10 6 Si) as olivine Mg no. decreases from 100 to 68, and at 3 GPa and 1200 °C olivine H contents increase from 0.017 to 0.054 wt% (278 to 10 000 H/10 6 Si) as Mg no. decreases from 100 to 55. The partition coefficient for H between pyroxene and olivine, D H opx/ol , decreases from 1.05 at 3 GPa and 1200 °C to 0.61 at 6 GPa and 1200 °C. The storage capacity of Fe-rich olivines with compositions expected in the martian mantle is ~1.5 times greater than those in the terrestrial mantle, suggesting that the geochemical behavior of H 2 O in the mantles of the two planets are quite similar. If 50% of the K 2 O on Mars remains in its mantle (Taylor et al. 2006), then a similar or greater proportion of the H 2 O is also in the mantle. Given accretionary models of the total martian H 2 O budget (Lunine et al. 2003), this suggests concentrations of 100–500 ppm H 2 O in the martian mantle and 0.1–1.9 wt% H 2 O in primary martian basalts.