Equilibrium Fe Isotope Fractionation Research Articles

Iron isotope fractionation between metal and silicate during core-mantle differentiation in rocky bodies

Fe isotope variations in rocky bodies reveal fundamental information about planetary evolution. However, experimental results have come to contradictory conclusions on the equilibrium Fe isotope fractionation between metal and silicate during core-mantle differentiation. Many different processes, including evaporation, core formation, partial melting and disproportion of mantle silicate, have been consequently proposed to explain the observed Fe isotope variations in rocky solar system bodies. Here we perform ab initio molecular dynamics simulations and find that the anharmonicity in iron strongly decreases the force constant of Fe at low pressures (<∼50 GPa), which even reverses the equilibrium Fe isotope fractionation between metal and silicate. We conclude that pyrolitic melt is always enriched in heavy Fe isotopes relative to liquid Fe-alloys, no matter what pressure. Therefore core-mantle differentiation will play a significant role in explaining the heavy Fe isotope compositions of the mantles of some rocky bodies (e.g., Earth, the ureilite parent body, and possibly the asteroid Vesta). As all previously proposed processes for Fe isotope fractionation can only enrich the mantle-derived rocks in heavy Fe isotopes, the near/sub-chondritic Fe isotope signatures of Mars and the aubrite parent body thus imply that iron sulfide enriched in light Fe isotopes may significantly contribute to the iron components of those meteoritic samples.

Fe isotope fractionation caused by phase transition of FeS and implications for Fe isotope signatures of the mantle and core

Iron is the most important component of the cores of terrestrial planets, and iron sulfide (FeS) is one of the preferred candidates present in these cores. FeS is also ubiquitous in Earth’s crust, peridotites, and extraterrestrial samples. Knowledge of the phase stability of FeS and Fe isotope fractionation between FeS phases and mantle silicates is of great importance for understanding the interior of the Earth and terrestrial planets. In this study, first-principles methods were used to study the pressure-dependent phase stability of FeS and equilibrium Fe isotope fractionation in FeS, hexagonal close-packed (hcp) Fe, and mantle silicates at the pressure of Earth’s interior. FeS underwent four phase transitions at 0 K. The first is the transition from FeS I to FeS II at 2.8 GPa, the second from FeS II to FeS III at 7.5 GPa, the third from FeS III to FeS VI at 74.2 GPa, and the fourth from FeS VI to FeS VII at 122.2 GPa. Apart from the fact that the transition from FeS I to FeS II causes negligible Fe isotope fractionation, other phase transitions can cause measurable Fe isotope fractionation at corresponding pressures along the geotherm. Fe isotopes exhibit measurable fractionation between FeS and mantle silicates under mantle pressure–temperature conditions. Each phase was more enriched in heavy Fe with increasing depth in the pressure range of 7.5–90 GPa. If the silicate mantle is enriched in heavy Fe relative to the core or Fe has negligible isotope fractionation between them under the core-mantle boundary (CMB) conditions of the Earth, the Fe2+/(Fe2++Mg) in (Fe2+, Mg)SiO3 post-perovskite is less than 50%. At the temperature–pressure conditions of Earth’s core, equilibrium Fe isotope fractionation between hcp Fe and FeS VII can be neglected. FeS III is more likely to exist in the Martian core relative to FeS VI.

First-principles investigation of equilibrium iron isotope fractionation in Fe1−xSx alloys at Earth's core formation conditions

Iron is one of the most abundant non-volatile elements in the solar system. As a major component of planetary metallic alloys, its immiscibility with silicates plays a major role in planetary formation and differentiation. Information about these processes can be gained by studying the equilibrium Fe isotope fractionation between metal alloys and molten silicates at conditions of core formation. In particular, recent attention has been paid to 56Fe/54Fe equilibrium isotope fractionation at conditions relevant to Earth's core formation and the influence that light elements (O, H, C, Ni, Si and S) have had in this process. Most of these experimental studies relied on the measurement of Fe isotope fractionation from quenched phases of silicate melts and molten iron alloys. The experimental works are extremely challenging, and may suffer different drawbacks. To overcome this, we use ab-initio computational methods to perform a systematic study of the 56Fe/54Fe equilibrium isotope fractionation in molten and solid Fe1−xSx alloys at conditions of the core formation (60 GPa, 3000 K). We show for the first time, that equilibrium isotope fractionation factors from solid systems can be used as proxies for molten systems with differences between these two methods less than 0.01‰ at the relevant P-T conditions. Additionally, we discuss the effect of sulphur concentration on the equilibrium Fe isotope fractionation and show that although there are some structural changes due to atom substitutions, the wide range of studied concentrations produces β-factors that are constant within ∼0.02‰. Finally, we discuss the implications of our results for the interpretation of recent experiments and the understanding of core crystallisation processes.

Iron, magnesium, and titanium isotopic fractionations between garnet, ilmenite, fayalite, biotite, and tourmaline: Results from NRIXS, ab initio, and study of mineral separates from the Moosilauke metapelite

Interpreting isotopic signatures documented in natural rocks requires knowledge of equilibrium isotopic fractionation factors. Here, we determine equilibrium Fe isotope fractionation factors between several common rock-forming minerals using a comparative approach involving three independent methods: (i) isotopic analyses of natural minerals from a metapelite from Mt. Moosilauke, New Hampshire, for which equilibration temperature and pressure are well constrained to be near the aluminosilicate triple point (T ≃ 500 °C, P ≃ 4 kbar), (ii) Nuclear Resonant Inelastic X-ray Scattering (NRIXS) measurements of Fe force constants of minerals, and (iii) Density Functional Theory (DFT) ab initio calculations of Fe force constants of minerals.The minerals studied for Fe isotopes include, in increasing order of their β-factors: garnet < ilmenite ≈ fayalite < biotite < tourmaline < muscovite ≈ plagioclase. Some of this ordering is affected by the presence of Fe3+ in the minerals, which tends to form stiffer bonds and be associated with heavy Fe isotope enrichments relative to Fe2+. We are, however, able to assess the magnitude of the effect of the ratio Fe3+/ΣFe on equilibrium fractionation factors, notably on the ilmenite-hematite solid solution. Equilibrium Fe isotopic fractionation factors between garnet, ilmenite, biotite, tourmaline and fayalite are determined. We also report Mg and Ti isotopic compositions of selected Moosilauke minerals that allow us to better constrain the equilibrium fractionation factors for garnet-biotite-tourmaline (Mg isotopes) and biotite-ilmenite (Ti isotopes).We show how the newly determined equilibrium fractionation factors can be used to address diverse problems in Earth and planetary sciences, notably (i) Fe and Mg isotopic fractionation during anatexis, (ii) Fe isotopic fractionation in lunar ilmenite, and (iii) Ti isotopic fractionation during fluvial transport of minerals.

Iron isotope exchange and fractionation between jarosite and aqueous Fe(II)

Jarosite is one of the critical minerals that regulates acidity and contaminants in acid-sulfate environments and its Fe isotope composition may shed light on its formation, transformation and recrystallization over time. Interpretation of its Fe isotope composition requires understanding the equilibrium Fe isotope fractionation factor between jarosite and other Fe-bearing minerals and aqueous species. Here we explore Fe isotope exchange and fractionation between jarosite and Fe(II)aq under acidic conditions using the three-isotope method (54Fe-56Fe-57Fe). A reversal-approach to equilibrium was applied by reacting synthetic jarosite and natural natrojarosite with two 57Fe-enriched Fe(II)aq solutions that had initial 56Fe/54Fe ratios above and below the predicted equilibrium value. No change in dissolved Fe(II) concentrations were observed with time but the 57Fe/56Fe ratio of Fe(II)aq decreased towards the system mass balance, suggesting a high degree of equilibration of the fluid with the solid phase despite no net Fe(II) sorption (within error). There is a negative relationship between pH and Fe isotope exchange, with Fe isotope exchange proceeding as pH decreases. This may be explained by dissolution of hydronium jarosite and reprecipitation of natrojarosite, coupled H3O+ - Na+ exchange, or jarosite decomposition, although no Fe-oxyhydroxide phases were identified from XRD. Calculation of the amount of Fe atoms in jarosite that exchanged with Fe(II)aq indicates that jarosite recrystallization was limited to a few percent. When the initial δ56Fe value of Fe(II)aq was greater than the presumed equilibrium value its isotopic value substantially decreased with time whereas the δ56Fe values of Fe(II)aq increased with time when it had an initial value below the suspected equilibrium composition. In each case, the isotopic composition of Fe(II)aq approached similar values, providing a high degree of confidence of an attainment of equilibrium. Calculation of the Fe(II)aq–jarosite and Fe(II)aq-natrojarosite equilibrium fractionation factors at 22 °C were −2.26‰ (±0.27‰, 2σ) and −2.19‰ (±0.18‰, 2σ), respectively. This indicates that during jarosite recrystallization in the presence of Fe(II), jarosite is expected to become isotopically heavier as lighter isotopes are fractionated into Fe(II). These values differ from the estimated fractionation factors derived from NRIXS spectroscopy and molecular modeling. The differences between experiments and theory may reflect surface exchange, which was likely in our study, versus predicted bulk thermodynamic properties of the mineral.

Calibrating equilibrium Fe isotope fractionation factors between magnetite, garnet, amphibole, and biotite

Equilibrium isotope fractionation factors are the basis for application of stable isotopes in geological studies. Experimental calibration and theoretical calculations have been employed to determine Fe isotope fractionation factors for a variety of minerals, however, these methods have their limitations. An alternative approach to calibrating inter-mineral Fe isotope fractionation factors is to use well-characterized geological samples; this approach has unique advantages over the other methods including attainment of equilibrium at relatively low temperatures. In this study, we investigated the Fe isotope composition of magnetite (Mt), garnet (Grt), amphibole (Amp) and biotite (Bt) from the metamorphosed Yingshan iron formation in South China. Two independent geothermometers, quartz–magnetite oxygen isotope and amphibole–garnet–biotite Fe–Mg exchange geothermometers, give a consistent metamorphic temperature of 538 ± 39 °C. The Fe isotope composition of the different Fe-bearing minerals is highly variable in different ironstone samples, with δ56Fe values (relative to IRMM-014) ranging from −0.23 to +0.37‰ in magnetite, −0.44 to +0.09‰ in amphibole, −0.78 to +0.02‰ in garnet, and −0.61 to +0.04‰ in biotite. Despite that, the offsets of δ56Fe values for mineral pairs are consistent, implying attainment of equilibrium isotope fraction between these minerals. Inter-mineral Fe isotopic fractionations (±2 standard deviation) measured from the multiple mineral pairs are Δ56FeMt–Grt = +0.55 ± 0.08‰, Δ56FeMt–Amp = +0.25 ± 0.06‰, Δ56FeMt–Bt = +0.42 ± 0.11‰, Δ56FeAmp–Grt = +0.34 ± 0.15‰, Δ56FeAmp–Bt = +0.26 ± 0.04‰, and Δ56FeBt–Grt = +0.13 ± 0.08‰. Based on the well-defined metamorphic temperature (538 ± 39 °C) and internally consistent inter-mineral fractionation factors, the temperature-dependent functions for equilibrium Fe isotope fractionation between the following mineral pairs are derived:103lnαMt–Grt = 0.36(±0.05) × 106/T2, 103lnαMt–Bt = 0.28(±0.07) × 106/T2,103lnαMt–Amp = 0.16(±0.04) × 106/T2, 103lnαAmp–Grt = 0.22(±0.10) × 106/T2,103lnαAmp–Bt = 0.17(±0.03) × 106/T2, and 103lnαBt–Grt = 0.09(±0.05) × 106/T2,where Fe3+/ΣFe ratio is 0.05 ± 0.02 in garnet, 0.29 ± 0.04 in biotite, and 0.24 ± 0.06 in amphibole for the above functions. The equilibrium Fe isotope fractionation factors derived in this study enable estimation of the metamorphic temperature of rocks that contain these mineral pairs, and identification of secondary processes that may have induced disequilibrium Fe isotope distribution in rocks, such as retrograde metamorphism and hydrothermal alteration.

The effect of core composition on iron isotope fractionation between planetary cores and mantles

We have conducted high-pressure, high-temperature isotope exchange experiments between molten silicate and molten Fe–Si–C-alloys to constrain the effect of Si on equilibrium Fe isotope fractionation during planetary core formation. The values of Δ57FeMetal-Silicate at 1850°C and 1 GPa determined by high-resolution MC-ICP-MS in this study range from −0.013±0.054‰ (2SE) to 0.072 ± 0.085‰ with 1.34–8.14 atom % Si in the alloy, respectively. These results, although not definitive on their own, are consistent with previous experimental results from our group and a model in which elements that substitute for Fe atoms in the alloy structure (i.e., Ni, S, and Si) induce a fractionation of Fe isotopes between molten silicate and molten Fe-alloys during planetary differentiation. Using in situ synchrotron X-ray diffraction data for molten Fe-rich alloys from the literature, we propose a model to explain this fractionation behavior in which impurity elements in Fe-alloys cause the nearest neighbor atomic distances to shorten, thereby stiffening metallic bonds and increasing the preference of the alloy for heavy Fe isotopes relative to the silicate melt. This fractionation results in the bulk silicate mantles of the smaller terrestrial planets and asteroids becoming isotopically light relative to chondrites, with an enrichment of heavy Fe isotopes in their cores, consistent with magmatic iron meteorite compositions. Our model predicts a bulk silicate mantle δ57Fe ranging from −0.01‰ to −0.12‰ for the Moon, −0.06‰ to −0.33‰ for Mars, and −0.08‰ to −0.33‰ for Vesta. Independent estimates of the δ57Fe of primitive mantle source regions that account for Fe isotope fractionation during partial melting agree well with these ranges for all three planetary bodies and suggest that Mars and Vesta have cores with impurity (i.e., Ni, S, Si) abundances near the low end of published ranges. Therefore, we favor a model in which core formation results in isotopically light bulk silicate mantles for the Moon, Mars, and Vesta. The processes of magma ocean crystallization, mantle partial melting, and fractional crystallization of mantle-derived melts are all likely to result in heavy Fe isotope enrichment in the melt phase, which can explain why basaltic samples from these planetary bodies have variable δ57Fe values consistently heavier than our bulk mantle estimates. Additionally, we find no clear evidence that Fe isotopes were fractionated to a detectable level by volatile depletion processes during or after planetary accretion, although it cannot be ruled out.

Iron isotope exchange and fractionation between hematite (α-Fe2O3) and aqueous Fe(II): A combined three-isotope and reversal-approach to equilibrium study

Hematite is the most thermodynamically stable Fe oxide at the Earth’s surface and its isotopic composition may record past biogeochemical Fe cycling and paleo-environmental conditions. Proper interpretation of its isotopic values requires an understanding of the equilibrium Fe isotope fractionation factor between hematite and other Fe-bearing minerals and aqueous species. Here, we use the three-isotope method (54Fe-56Fe-57Fe, with 56Fe/54Fe and 57Fe/56Fe ratios expressed in delta notation relative to the IRMM-014 isotope standard) to simultaneously determine Fe isotope exchange, via tracking an enriched tracer-isotope (i.e. 57Fe), and measure mass-dependent isotope fractionation between aqueous Fe(II) (Fe(II)aq) with structural-Fe(III) in hematite, as determined from variations in 56Fe/54Fe ratios. Specifically, we used a reversal-approach to equilibrium by reacting three hematite samples of varying particle size and reactivity with two Fe(II)aq solutions that had initial 56Fe/54Fe ratios above and below the presumed equilibrium value. We confirm that Fe(II)aq readily exchanges with hematite at low temperature, and that the extent of exchange depends on hematite particle size. In three-isotope plots, δ56Fe values of Fe(II)aq exhibit unique trajectories depending on the initial hematite particle size. Rapid exchange occurring between Fe(II)aq and small hematite particles are affected by an apparent mixing reaction, which results in the δ56Fe value of Fe(II)aq to approach the isotopic composition of hematite. These trajectories exhibit inflections to more negative δ56Fe values during continued exchange. The location of these inflections depend on particle size and are interpreted to represent a change in the recrystallization mechanism from rapid “heterogeneous” recrystallization, which is limited to surface exchange, to “homogeneous” recrystallization involving the bulk hematite. Slow isotopic exchange occurring at longer reaction times appears to approximate equilibrium for all particle sizes with measured Fe(II)aq-hematite Fe isotope fractionation factors at 22 °C ranging from −2.61 ‰ (±0.26‰, 2σ) to −3.14 ‰ (±0.34‰, 2σ) in δ56Fe, and an average value of −2.8‰. This value is consistent with estimated fractionation factors measured during microbial Fe reduction experiments and with calculated and spectroscopically derived reduced partition function ratios. The three-isotope method provides a robust measure of equilibrium fractionation if proper constrains are utilized to eliminate extrapolation from partial exchange. Furthermore, it can provide insight into the mechanism of mineral-fluid exchange which is necessary for accurate quantification of the extent of mineral recrystallization when using isotopic tracers.

Iron isotope fractionation during crustal anatexis: Constraints from migmatites from the Dabie orogen, Central China

Iron (Fe) isotope fractionation could occur during crustal melting, which may contribute to the isotopic variations of granites. Our current knowledge on Fe isotope fractionation during crustal melting remains rare. Thirteen leucosomes, 9 melanosomes and 1 amphibolitic schollen in 11 migmatites from the Dabie orogen, Central China were measured to investigate Fe isotope fractionation during crustal anatexis. The melanosomes and amphibolitic schollen yield δ56Fe values from 0.018±0.031‰ to 0.152±0.027‰ with an average of 0.106‰. The leucosomes have δ56Fe values from 0.107±0.035‰ to 0.512±0.028‰, variably higher than their coexisting melanosomes by 0.00~0.36‰. The δ56Fe of leucosomes and melanosomes, as well as apparent isotope fractionation (Δ56FeL-M) between them, do not correlate with the abundance of crystalline carbonate, loss on ignition, and Th/U, indicating the insignificant effect of fluid components. High δ56Fe values of three leucosomes (≥0.34‰), giving large Δ56FeL-M from 0.22 to 0.36‰, were likely produced by feldspar accumulation, evidenced by their high Eu*, low FeOt/Al2O3 and high plagioclase abundance up to 50vol%. Capture of peritectic amphiboles in another leucosome may explain its low Δ56FeL-M ~0‰. After screening out these samples, 7 leucosome – melanosome pairs yield identical Δ56FeL-M averaging 0.093±0.056‰ (2SD, N=7) within analytical uncertainties. No correlation between Δ56FeL-M and Mg#, FeOt/Al2O3 or TiO2/FeOt rules out the possibility that the fractionation observed here is produced by fractional crystallization and/or sub-solidus isotope re-equilibrium between leucosomes and melanosomes. Therefore, we suggest these consistent Δ56FeL-M should record equilibrium Fe isotope fractionation during crustal anatexis producing these migmatites. The leucosomes and melanosomes have comparable Fe3+/ΣFe~0.40, and Δ56FeL-M approximates 0.07‰ at Fe3+/ΣFeleucosomeFe3+/ΣFemenalosome=1. Fe isotope fractionation during crustal anatexis revealed here thus is not dominantly controlled by distribution of Fe3+ and Fe2+, but by the difference in coordination number of iron between granitic melts and residual mafic minerals. The observation here argues that Fe isotope fractionation during crustal partial melting should also contribute to the previously revealed δ56Fe variation in high silica granitic rocks.

Fe isotope fractionation between chalcopyrite and dissolved Fe during hydrothermal recrystallization: An experimental study at 350 °C and 500 bars

Equilibrium Fe isotope fractionation between chalcopyrite and dissolved Fe was determined in acidic chloride-bearing fluid at 350°C and 500bars. The study utilized deformable gold-cell technology, which allowed time-series sampling of solution during chalcopyrite recrystallization and isotope exchange. A key element of the experimental design involved the addition of anomalous dissolved 57Fe to an on-going experiment as a means of determining the degree and rate of isotope exchange. Taking explicit account of imposed chemical and isotopic mass balance constraints of Fe in fluid and mineral (chalcopyrite) reservoirs, these data indicate that no more than 1000h is required for the isotopically anomalous dissolved Fe reservoir to exchange completely with the coexisting chalcopyrite. The experimental calibration of the rate of Fe isotope exchange for the δ57Fe-spiked experiment provides critical insight for the time necessary to achieve Fe isotope exchange in two non-spiked, but otherwise identical experiments. The Fe isotope data indicate that the equilibrium fractionation between chalcopyrite and dissolved Fe, Δ56FeCpy-Fe (aq), at 350°C is small, 0.09±0.17‰ (2σ), and is in good agreement with recent theoretical equilibrium predictions. Owing to the apparent rate of Fe isotope exchange at 350°C, it is likely that chalcopyrite formed at high temperature deep-sea vents (black smoker systems) achieves isotopic equilibrium, and effectively records the Fe isotopic composition of the coexisting end-member hydrothermal fluid. Comparison of the experimental mineral–fluid equilibrium fractionation factors with conjugate chalcopyrite and dissolved Fe pairs sampled from high temperature hydrothermal vent systems at Axial Caldera and Main Endeavour Field (Juan de Fuca Ridge) are in agreement with this inference. The experimental data were further used to determine the mineral–mineral equilibrium Fe isotope fractionation between pyrite-chalcopyrite, Δ56FePyr-Cpy, at 350°C by combining previously determined pyrite-Fe2+(aq) equilibrium fractionation data with chalcopyrite-Fe2+(aq) from this study. The empirically determined Δ56FePyr-Cpy value, 0.90±0.34‰ (2σ), is consistent with theoretical predictions, and when coupled with mineral–fluid Fe isotope fractionation systematics and experimentally determined exchange rates, helps to delineate processes of sulfide mineralization in hydrothermal systems.

Iron isotope fractionation between aqueous Fe(II) and goethite revisited: New insights based on a multi-direction approach to equilibrium and isotopic exchange rate modification

The Fe isotope compositions of naturally occurring Fe oxide minerals provide insights into biogeochemical processes that occur in modern and ancient environments. Key to understanding isotopic variations in such minerals is knowledge of the equilibrium Fe isotope fractionation factors between common minerals and aqueous Fe species. Because experimental measurements of isotopic fractionation may reflect a combination of kinetic and equilibrium fractionations during rapid dissolution and precipitation, even in experiments that employ the three-isotope method, assessment of the attainment of equilibrium is often difficult. Here, we re-examine Fe isotope exchange, via a 57Fe tracer, and natural mass-dependent fractionation, through changes in initial 56Fe/54Fe ratios, between aqueous Fe(II) (Fe(II)aq) and goethite. This approach uses the three-isotope method, but is distinct in its evaluation of kinetic isotope fractionation and the attainment of equilibrium by: (i) employing a multi-direction approach to equilibrium at 22°C via reaction of three Fe(II)aq solutions that had different initial 56Fe/54Fe ratios, (ii) conducting isotopic exchange experiments at elevated temperature (50°C), and (iii) modifying the rate of isotopic exchange through a combination of trace-element substitutions and particle coarsening to evaluate corresponding temporal changes in fractionation trajectories that may reflect changing instantaneous fractionation factors. We find that rapid isotopic exchange produces kinetic isotope effects between Fe(II)aq and goethite, which shifts the 56Fe/54Fe ratios of Fe(II)aq early in reactions toward that of goethite, indicating that the instantaneous Fe(II)aq–goethite fractionation factor under kinetic conditions is small. Importantly, however, this kinetic fractionation is “erased” with continued reaction, and this is evident by the congruence for multiple-exchange trajectories of distinct initial Fe(II)aq solutions toward the same final value over long reaction times. Experiments at higher temperature result in a smaller fractionation between Fe(II)aq and goethite, consistent with a 1/T2 temperature dependence. Coarsened and trace-element substituted goethites that had low surface areas produced much slower rates of isotopic exchange than the chemically pure forms or goethite of smaller crystal size, resulting in only partial isotopic mixing (10–40%). Fractionation-exchange trajectories produced during slow isotopic exchange are linear, however, and extrapolate to the same (within error) Fe(II)aq–goethite fractionations at 100% isotopic mixing as that for reactions of pure goethite. We conclude that the equilibrium 56Fe/54Fe fractionation for Fe(II)aq–goethite at 22°C ranges between −1.04±0.08‰ and −1.22±0.11‰ (2σ), depending on particle size, where the more negative fractionation is influenced by surface Fe that has distinct isotopic properties; these results are consistent with earlier measurements by Beard et al. (2010). This work highlights the utility of using multiple exchange-fractionation trajectories to verify the attainment of equilibrium and resolving kinetic isotope effects, and the importance of isotopic exchange rate on disequilibrium mixing between components. We recommend that these techniques are essential for unambiguously demonstrating that measured fractionations during isotopic exchange experiments are, in fact, equilibrium fractionations.

Experimental determination of equilibrium Fe isotopic fractionation between pyrite and dissolved Fe under hydrothermal conditions

Fe isotope fractionation between pyrite (FeS2) and dissolved Fe in NaCl- and sulfur-bearing aqueous fluids was determined under hydrothermal conditions (300–350°C, 500bars). The data were collected using two different, but complementary, experimental approaches, one involving classical Fe isotope exchange between Fe in pyrite and dissolved Fe in coexisting fluid, while the other involved homogenous precipitation of pyrite in a redox and pH buffered chemical system. Results from these experiments indicate equilibrium Fe isotopic fractionation between pyrite and fluid, Δ56FePyr–Fe(aq), of 0.99±0.29‰ (2σ), in 56Fe/54Fe. The experimentally determined equilibrium pyrite–fluid Fe isotopic fractionation agrees with theoretical and spectrally-based predictions. The second series of experiments were conducted to better constrain the effect of precipitation rate on the temporal evolution of the Fe isotopic composition of pyrite and Fe bearing fluids in dynamic mixing environments, such as hydrothermal vent sites at mid-ocean ridges. Rapid homogenous precipitation of pyrite at 300 and 350°C indicates that δ56Fe of dissolved Fe is significantly greater than pyrite that formed during the earliest stage of the experiment, possibly facilitated by either equilibrium or kinetic isotope effects involving FeS as a reactant during pyrite formation. Subsequent recrystallization of pyrite results in a Fe isotopic fractionation with dissolved Fe that moves towards the experimentally determined equilibrium Fe isotopic fractionation with reaction progress. The experimental data reported here may help to decipher the complex kinetic and thermodynamic processes involved in pyrite formation at deep sea vents, while also providing constraints for the rapidly developing theoretical models used to estimate equilibrium Fe isotope fractionation between pyrite and fluid at elevated temperatures and pressures.

Experimental evidence for the absence of iron isotope fractionation between metal and silicate liquids at 1 GPa and 1250–1300 °C and its cosmochemical consequences

Iron isotope fractionation during metal–silicate differentiation has been proposed as a cause for differences in iron isotope compositions of chondrites, iron meteorites and the bulk silicate Earth. Stable isotope fractionation, however, rapidly decreases with increasing temperature. We have thus performed liquid metal–liquid silicate equilibration experiments at 1250–1300°C and 1GPa to address whether Fe isotope fractionation is resolvable at the lowest possible temperatures for magmatic metal–silicate differentiation. A centrifuging piston cylinder apparatus enabled quantitative metal–silicate segregation. Elemental tin or sulphur was used in the synthetic metal-oxide mixtures to lower the melting temperature of the metal. The analyses demonstrate that eight of the 10 experimental systems equilibrated in a closed isotopic system, as was assessed by varying run durations and starting Fe isotope compositions. Statistically significant iron isotope fractionation between quenched metals and silicates was absent in nine of the 10 experiments and all 10 experiments yield an average metal–silicate fractionation factor of 0.01±0.04‰, independent of whether graphite or silica glass capsules were used. This implies that Fe isotopes do not fractionate during low pressure metal–silicate segregation under magmatic conditions. In large bodies such as the Earth, fractionation between metal and high pressure (>20GPa) silicate phases may still be a possible process for equilibrium fractionation during metal–silicate differentiation. However, the 0.07±0.02‰ heavier composition of bulk magmatic iron meteorites relative to the average of bulk ordinary/carbonaceous chondrites cannot result from equilibrium Fe isotope fractionation during core segregation. The up to 0.5‰ lighter sulphide than metal fraction in iron meteorites and in one ordinary chondrite can only be explained by fractionation during subsolidus processes.

Experimental determination of iron isotope fractionations among [formula omitted]–FeSaq–Mackinawite at low temperatures: Implications for the rock record

The Fe isotope fractionation factors among aqueous ferrous iron (Feaq2+), aqueous FeS clusters (FeSaq), and nanoparticulate mackinawite under neutral and mildly acidic and alkaline pH conditions have been determined using the three-isotope method. Combined voltammetric analysis and geochemical modeling were used to determine the Fe speciation in the experimental systems. The equilibrium 56Fe/54Fe fractionation factor at 20°C and pH 7 has been determined to be −0.32±0.29 (2σ)‰ between Feaq2+ (minor FeSaq also present in the experiment) and mackinawite. This fractionation factor was essentially constant when pH was changed to 6 or 8. When equal molarity of HS− and Feaq2+ were added to the system, however, the isotopic fractionation at pH 7 changed to −0.64±0.36 (2σ)‰, correlating with a significant increase in the proportion of FeHS+ and FeSaq. These results highlight a more important role of aqueous Fe–S speciation in the equilibrium Fe isotope fractionation factor than recognized in previous studies. The isotopic fractionation remained constant when temperature was increased from 20°C to 35°C for fractionation factors between Feaq2+, and mackinawite and between dominantly FeHS+ and mackinawite. Synthesis experiments similar to those of Butler et al. (2005) and Guilbaud et al. (2010) at pH 4 show consistent results: over time, the aqueous Fe–mackinawite fractionation decreases but even after 38days of aging the fractionation factor is far from the equilibrium value inferred using the three-isotope method. In contrast, at near-neutral pH the fractionation factor for the synthesis experiment reached the equilibrium value in 38days. These differences are best explained by noting that at low pH the FeS mackinawite particles coarsen more rapidly via particle aggregation, which limits isotopic exchange, whereas at higher pH mackinawite aggregation is limited, and Fe isotope exchange occurs more rapidly, converging on the equilibrium value. These results suggest that mackinawite formed in natural environments at near-neutral or alkaline pH are unlikely to retain kinetic isotope fractionations, but are more likely to reflect equilibrium isotope compositions. This in turn has important implications for interpreting iron isotope compositions of Fe sulfides in natural systems.

Coupled lithium- and iron isotope fractionation during magmatic differentiation

In this study, we investigated Fe and Li isotope fractionation between mineral separates of olivine pheno- and xenocrysts (including one clinopyroxyene phenocryst) and their basaltic hosts. Samples were collected from the Canary Islands (Teneriffa, La Palma) and some German volcanic regions (Vogelsberg, Westerwald and Hegau). All investigated bulk samples fall in a tight range of Li and Fe isotope compositions (δ56Fewr=0.06–0.17‰ and δ7Lima=2.5–5.2‰, assuming δ7Li of the olivine-free matrix is virtually identical to that of the bulk sample for mass balance reasons). In contrast, olivine phenocrysts display highly variable, but generally light Fe and mostly light Li isotope compositions compared to their respective olivine-free basaltic matrix, which was considered to represent the melt (with δ56Feol=−0.24 to 0.14‰ and δ7Liol=−10.5 to +6.5‰, respectively). Single olivine crystals from one sample display even a larger range of δ56Feol between −0.7 and +0.1‰. One single clinopyroxene phenocryst displays the lightest Li isotope composition (δ7Licpx=−17.7‰), but no Fe isotope fractionation relative to melt. The olivine phenocrysts show variable Mg# and Ni (correlated in most cases) that range between 0.89 and 0.74 and between 300 and 3000μg/g, respectively. These olivines likely grew by fractional crystallization in an evolving magma. One sample from the Vogelsberg volcano contained olivine xenocrysts (Mg#>0.89 and Ni>3000μg/g), in addition to olivine phenocrysts. This sample displays the highest Li- and the second highest Fe-isotope fractionation between olivine and melt (Δ7Liol-melt=−13; Δ56Feol-melt=−0.29).Our data, i.e. the variable olivine- at constant whole rock and matrix isotope compositions, strongly indicate disequilibrium, i.e. kinetic Fe and Li isotope fractionation between olivine and melt (for Li also between cpx and melt) during fractional crystallization. Δ7Liol-melt is correlated with the Li partitioning between olivine and melt (i.e. with Liol/Limelt), indicating Li isotope fractionation due to preferential (faster) diffusion of 6Li into olivine during fractional crystallization. Olivine with low Δ7Liol-melt, also have low Δ56Feol-melt, indicating that Fe isotope fractionation is also driven by diffusion of isotopically light Fe into olivine, potentially, as Fe–Mg inter-diffusion. The lowest Δ56Feol-melt (−0.40) was observed in a sample from Westerwald (Germany) with abundant magnetite, indicating relatively oxidizing conditions during magma differentiation. This may have enhanced equilibrium Fe isotope fractionation between olivine and melt or fine dispersed magnetite in the basalt matrix may have shifted its Fe isotope composition towards higher δ56Fe. The decoupling of Li- and Fe isotope fractionation in cpx is likely due to faster diffusion of Li relative to Fe in cpx, implying that the large investigated cpx phenocryst resided in the magma for only a short period of time which was sufficient for Li- but not for Fe diffusion. The absence of any equilibrium Fe isotope fractionation between the investigated cpx phenocryst and its basaltic host may be related to the similar Fe3+/Fe2+ of cpx and melt. In contrast to cpx, the generally light Fe isotope composition of all investigated olivine separates implies the existence of equilibrium- (in addition to diffusion-driven) isotope fractionation between olivine and melt, on the order of 0.1‰.

Experimental determination of the equilibrium Fe isotope fractionation between [formula omitted] and FeS m (mackinawite) at 25 and 2 °C

We report the first experimentally-determined metal isotope equilibrium fractionation factors for a metal sulphide at ambient temperatures and pressures. Mackinawite, referred here as FeS m (where the subscript m indicates mackinawite), can be a reactive component in diagenetic pyrite formation and the extent of equilibration between FeS m and dissolved Fe(II) has direct implications the δ 56Fe signatures recorded in diagenetic pyrite. The measured equilibrium Fe isotope fractionation between Fe(II) aq and FeS m is Δ 56Fe Fe(II)–FeS = −0.52 ± 0.16‰ at 2 °C and Δ 56Fe Fe(II)–FeS = −0.33 ± 0.12‰ at 25 °C and pH 4. At the experimental pH the equilibrium fractionation factor between all dissolved Fe(II) species and FeS m (Δ 56Fe Fe(II)–FeS) equates to the fractionation factor between Fe aq 2 + and FeS m ( Δ 56 Fe Fe 2 + – FeS ) . The measured fractionations are of the same order as other non-redox fractionations measured in low-temperature Fe–C–O systems. We show that at low temperature, the Fe(II) aq–FeS m system is slowly asymptotic to isotopic equilibrium and consequently, FeS m is likely to partially conserve kinetically derived isotopic signatures generated on precipitation. Combined with the range of published kinetic fractionations measured on FeS m precipitation, our data suggest that, subject to the degree of isotope exchange during equilibration, FeS m can display δ 56Fe compositions encompassing a range of ∼1.4‰.

Stable Iron Isotope Fractionation Between Aqueous Fe(II) and Hydrous Ferric Oxide

Despite the ubiquity of poorly crystalline ferric hydrous oxides (HFO, or ferrihydrite) in natural environments, stable Fe isotopic fractionation between HFO and other Fe phases remains unclear. In particular, it has been difficult to determine equilibrium Fe isotope fractionation between aqueous Fe(II) and HFO due to fast transformation of the latter to more stable minerals. Here we used HFO stabilized by the presence of dissolved silica (2.14 mM), or a Si-HFO coprecipitate, to determine an equilibrium Fe(II)-HFO fractionation factor using a three-isotope method. Iron isotope exchange between Fe(II) and HFO was rapid and near complete with the Si-HFO coprecipitate, and rapid but incomplete for HFO in the presence of dissolved silica, the latter case likely reflecting blockage of oxide surface sites by sorbed silica. Equilibrium Fe(II)-HFO (56)Fe/(54)Fe fractionation factors of -3.17 ± 0.08 (2σ)‰ and -2.58 ± 0.14 (2σ)‰ were obtained for HFO plus silica and the Si-HFO coprecipitate, respectively. Structural similarity between ferrihydrite and hematite, as suggested by spectroscopic studies, combined with the minor isotopic effect of dissolved silica, imply that the true equilibrium Fe(II)-HFO (56)Fe/(54)Fe fractionation factor in the absence of silica may be ∼-3.2‰. These results provide a critical interpretive context for inferring the stable isotope effects of Fe redox cycling in nature.

New data on equilibrium iron isotope fractionation among sulfides: Constraints on mechanisms of sulfide formation in hydrothermal and igneous systems

Fe, S, and Cu reduced partition function ratios (β-factors) allow calculation of equilibrium isotope fractionation factors. β-Factors for chalcopyrite are calculated from experimental and theoretical partial phonon densities of state states ( Kobayashi et al., 2007). The Fe β-factors for mackinawite are calculated from Mössbauer spectroscopy data ( Bertaut et al., 1965). Excellent agreement exists between Fe β-factors for chalcopyrite calculated from the experimental and theoretical 57Fe phonon densities of states, supporting the reliability of the Fe β-factors for chalcopyrite. The 34S β-factor for chalcopyrite is consistent with experimental data on equilibrium sulfur isotope fractionation factors among sulfides and theoretical 34S β-factors, except those recently calculated by a DFT approach. Up-to-date experimental isotope-exchange data on equilibrium Fe isotope fractionation factors between minerals and aqueous Fe were critically reevaluated in conjunction with Fe β-factors for minerals, and the following expressions for β-factors for aqueous Fe 2+ and Fe 3+ were obtained: 10 3 ln β Fe aq 3 + = 1.0063 10 6 T 2 ; T > 273 K 10 3 ln β Fe aq 2 + = 0.6537 10 6 T 2 ; T > 273 K. .Application of these new β-factors to Fe and S isotope fractionation at seafloor hydrothermal conditions shows that chalcopyrite may precipitate in equilibrium with Fe in hydrothermal fluids. In contrast, Fe and S isotope compositions of pyrite are may not be controlled by isotope equilibration with Fe in the hydrothermal fluid. Previously reported data on Fe isotope fractionation between pyrite and chalcopyrite in intrusions in the Grasberg Igneous Complex seem to be reasonable, and are in general agreement with previous “non-isotopic” estimates. Fe isotope fractionation between pyrite and chalcopyrite appear to reflect equilibrium in the intrusions of the Grasberg Igneous Complex, and Fe isotopes may therefore be considered as a potential geothermometer. The pyrite–chalcopyrite Fe isotope geothermometer, however, cannot be applied to the skarn facies of the Grasberg complex, where the initial Fe isotope compositions may have been disturbed by subsequent processes related to skarn formation.

Effects of changing solution chemistry on Fe 3+/Fe 2+ isotope fractionation in aqueous Fe–Cl solutions

The range in 56Fe/ 54Fe isotopic compositions measured in naturally occurring iron-bearing species is greater than 5‰. Both theoretical modeling and experimental studies of equilibrium isotopic fractionation among iron-bearing species have shown that significant fractionations can be caused by differences in oxidation state (i.e., redox effects in the environment) as well as by bond partner and coordination number (i.e., nonredox effects due to speciation). To test the relative effects of redox vs. nonredox attributes on total Fe equilibrium isotopic fractionation, we measured changes, both experimentally and theoretically, in the isotopic composition of an Fe 2+–Fe 3+–Cl–H 2O solution as the chlorinity was varied. We made use of the unique solubility of FeCl 4 − in immiscible diethyl ether to create a separate spectator phase against which changes in the aqueous phase could be quantified. Our experiments showed a reduction in the redox isotopic fractionation between Fe 2+- and Fe 3+-bearing species from 3.4‰ at [Cl −] = 1.5 M to 2.4‰ at [Cl −] = 5.0 M, due to changes in speciation in the Fe–Cl solution. This experimental design was also used to demonstrate the attainment of isotopic equilibrium between the two phases, using a 54Fe spike. To better understand speciation effects on redox fractionation, we created four new sets of ab initio models of the ferrous chloride complexes used in the experiments. These were combined with corresponding ab initio models for the ferric chloride complexes from previous work. At 20 °C, 1000 ln β ( β = 56Fe/ 54Fe reduced partition function ratio relative to a dissociated Fe atom) values range from 6.39‰ to 5.42‰ for Fe(H 2O) 6 2+, 5.98‰ to 5.34‰ for FeCl(H 2O) 5 +, and 5.91‰ to 4.86‰ for FeCl 2(H 2O) 4, depending on the model. The theoretical models predict ferric–ferrous fractionation about half as large (depending on model) as the experimental results. Our results show (1) oxidation state is likely to be the dominant factor controlling equilibrium Fe isotope fractionation in solution and (2) nonredox attributes (such as ligands present in the aqueous solution, speciation and relative abundances, and ionic strength of the solution) can also have significant effects. Changes in the isotopic composition of an Fe-bearing solution will influence the resultant Fe isotopic signature of any precipitates.

Microbial production of isotopically light iron(II) in a modern chemically precipitated sediment and implications for isotopic variations in ancient rocks

The inventories and Fe isotope composition of aqueous Fe(II) and solid-phase Fe compounds were quantified in neutral-pH, chemically precipitated sediments downstream of the Iron Mountain acid mine drainage site in northern California, USA. The sediments contain high concentrations of amorphous Fe(III) oxyhydroxides [Fe(III)(am)] that allow dissimilatory iron reduction (DIR) to predominate over Fe-S interactions in Fe redox transformation, as indicated by the very low abundance of Cr(II)-extractable reduced inorganic sulfur compared with dilute HCl-extractable Fe. delta(56)Fe values for bulk HCl- and HF-extractable Fe were approximately 0. These near-zero bulk delta(56)Fe values, together with the very low abundance of dissolved Fe in the overlying water column, suggest that the pyrite Fe source had near-zero delta(56)Fe values, and that complete oxidation of Fe(II) took place prior to deposition of the Fe(III) oxide-rich sediment. Sediment core analyses and incubation experiments demonstrated the production of millimolar quantities of isotopically light (delta(56)Fe approximately -1.5 to -0.5 per thousand) aqueous Fe(II) coupled to partial reduction of Fe(III)(am) by DIR. Trends in the Fe isotope composition of solid-associated Fe(II) and residual Fe(III)(am) are consistent with experiments with synthetic Fe(III) oxides, and collectively suggest an equilibrium Fe isotope fractionation between aqueous Fe(II) and Fe(III)(am) of approximately -2 per thousand. These Fe(III) oxide-rich sediments provide a model for early diagenetic processes that are likely to have taken place in Archean and Paleoproterozoic marine sediments that served as precursors for banded iron formations. Our results suggest pathways whereby DIR could have led to the formation of large quantities of low-delta(56)Fe minerals during BIF genesis.