Equilibrium Iron Isotope Fractionation Research Articles

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.

An assessment of iron isotope fractionation during core formation

Iron is a ubiquitous element in terrestrial and extra-terrestrial settings and can provide clues as to the conditions during which planetary scale processes occurred. One example of this is determining the conditions accompanying metal core formation in rocky planets and planetesimals. For at least two decades there has been a growing database of experimental and natural data aimed at understanding whether iron isotopes fractionate during the separation of silicate and metal. While it has been argued that the data are not in agreement with one another, it is apparent that once criteria are established that prove equilibrated samples, there is good agreement amongst the different studies when looked at as a function of the metallic composition. Proving equilibrium is critical in these types of experiments. The three-isotope experimental technique for establishing equilibrium is found to be both mathematically and fundamentally sound. Further it is clear that the question of whether there is an equilibrium iron isotope fractionation between metal and silicate is not straightforward and that it can vary significantly as a function of temperature, pressure, metallic composition, oxygen fugacity, and silicate composition.

Combined iron and magnesium isotope geochemistry of pyroxenite xenoliths from Hannuoba, North China Craton: implications for mantle metasomatism

We present high-precision iron and magnesium isotopic data for diverse mantle pyroxenite xenoliths collected from Hannuoba, North China Craton and provide the first combined iron and magnesium isotopic study of such rocks. Compositionally, these xenoliths range from Cr-diopside pyroxenites and Al-augite pyroxenites to garnet-bearing pyroxenites and are taken as physical evidence for different episodes of melt injection. Our results show that both Cr-diopside pyroxenites and Al-augite pyroxenites of cumulate origin display narrow ranges in iron and magnesium isotopic compositions (δ57Fe = −0.01 to 0.09 with an average of 0.03 ± 0.08 (2SD, n = 6); δ26Mg = − 0.28 to −0.25 with an average of −0.26 ± 0.03 (2SD, n = 3), respectively). These values are identical to those in the normal upper mantle and show equilibrium inter-mineral iron and magnesium isotope fractionation between coexisting mantle minerals. In contrast, the garnet-bearing pyroxenites, which are products of reactions between peridotites and silicate melts from an ancient subducted oceanic slab, exhibit larger iron isotopic variations, with δ57Fe ranging from 0.12 to 0.30. The δ57Fe values of minerals in these garnet-bearing pyroxenites also vary widely (−0.25 to 0.08 in olivines, −0.04 to 0.25 in orthopyroxenes, −0.07 to 0.31 in clinopyroxenes, 0.07 to 0.48 in spinels and 0.31–0.42 in garnets). In addition, the garnet-bearing pyroxenite shows light δ26Mg (−0.43) relative to the mantle. The δ26Mg of minerals in the garnet-bearing pyroxenite range from −0.35 for olivine and orthopyroxene, to −0.34 for clinopyroxene, 0.04 for spinel and −0.68 for garnet. These measured values stand in marked contrast to calculated equilibrium iron and magnesium isotope fractionation between coexisting mantle minerals at mantle temperatures derived from theory, indicating disequilibrium isotope fractionation. Notably, one phlogopite clinopyroxenite with an apparent later metasomatic overprint has the heaviest δ57Fe (as high as 1.00) but the lightest δ26Mg (as low as −1.50) values of all investigated samples. Overall, there appears to be a negative co-variation between δ57Fe and δ26Mg in the Hannuoba garnet-bearing pyroxenite and in the phlogopite clinopyroxenite xenoliths and minerals therein. These features may reflect kinetic isotopic fractionation due to iron and magnesium inter-diffusion during melt–rock interaction. Such processes play an important role in producing inter-mineral iron and magnesium isotopic disequilibrium and local iron and magnesium isotopic heterogeneity in the subcontinental mantle.

Iron isotopic fractionation between silicate mantle and metallic core at high pressure

The +0.1‰ elevated 56Fe/54Fe ratio of terrestrial basalts relative to chondrites was proposed to be a fingerprint of core-mantle segregation. However, the extent of iron isotopic fractionation between molten metal and silicate under high pressure–temperature conditions is poorly known. Here we show that iron forms chemical bonds of similar strengths in basaltic glasses and iron-rich alloys, even at high pressure. From the measured mean force constants of iron bonds, we calculate an equilibrium iron isotope fractionation between silicate and iron under core formation conditions in Earth of ∼0–0.02‰, which is small relative to the +0.1‰ shift of terrestrial basalts. This result is unaffected by small amounts of nickel and candidate core-forming light elements, as the isotopic shifts associated with such alloying are small. This study suggests that the variability in iron isotopic composition in planetary objects cannot be due to core formation.

Iron isotopic systematics of UHP eclogites respond to oxidizing fluid during exhumation

AbstractThe iron stable isotope compositions (δ56Fe) and iron valence states of ultrahigh‐pressure eclogites from Bixiling in the Dabie orogen belt, China, were measured to trace the changes of geochemical conditions during vertical transportation of earth materials, for example, oxygen fugacity. The bulk Fe3+/ΣFe ratios of retrograde eclogites, determined by Mössbauer spectroscopy, are consistently higher than those of fresh eclogites, suggesting oxidation during retrograde metamorphism and fluid infiltration. The studied eclogites (five samples) display limited mid‐ocean ridge basalts (MORB)‐like (~0.10‰) δ56Fe values, which are indistinguishable from their protoliths, that is, gabbro cumulates formed through differentiation of mantle‐derived basaltic magma. This suggests that Fe isotope fractionation during continental subduction is limited. Garnet separates display limited δ56Fe variation ranging from −0.08 ± 0.07 ‰ to 0.02 ± 0.07‰, whereas coexisting omphacite displays a large variation of δ56Fe values from 0.15 ± 0.07‰ to 0.47 ± 0.07‰. Omphacite also has highly variable Fe3+/ΣFe ratios from 0.367 ± 0.025 to 0.598 ± 0.024, indicating modification after peak metamorphism. Omphacite from retrograde eclogites has elevated Fe3+/ΣFe ratios (0.54–0.60) compared to that from fresh eclogites (~0.37), whereas garnet displays a narrow range of ferric iron content with Fe3+/ΣFe ratios from 0.039 ± 0.013 to 0.065 ± 0.022. The homogenous δ56Fe values and Fe3+/ΣFe ratios of garnet suggest that it survived the retrograde metamorphism and preserved its Fe‐isotopic features and ferric contents of peak metamorphism. Because of similar diffusion rates of Fe and Mg in garnet and omphacite, and constant Δ26Mgomphacite‐garnet values (1.14 ± 0.04‰), equilibrium iron isotope fractionation between garnet and omphacite was probably achieved during peak metamorphism. Elevated Fe3+/ΣFe ratios of omphacite from retrograde eclogites and variant Δ56Feomphacite‐garnet values of the studied eclogites (0.13 ± 0.10‰ to 0.48 ± 0.10‰) indicate that oxidized geofluid infiltration resulted in the elevation of δ56Fe values of omphacite during retrograde metamorphism.

Crystal chemical constraints on inter-mineral Fe isotope fractionation and implications for Fe isotope disequilibrium in San Carlos mantle xenoliths

The origin of variations in iron isotope compositions of mantle minerals is uncertain, and predictions of equilibrium inter-mineral iron isotope fractionation conflict. This hinders interpretation of the petrologic and geochemical implications of Fe isotope data from mantle lithologies. To address this, we present a revised ionic model for predicting equilibrium iron isotope fractionation between mantle minerals and use it to interpret measured inter-mineral iron isotopic fractionation from five distinct mantle xenolith lithologies from San Carlos, Arizona. The samples represent a broad range of modal abundances and include lherzolite, harzburgite, dunite, clinopyroxenite, and websterite. The xenoliths exhibit Fe-isotopic variation between minerals in a single sample, and between samples. In all cases where spinel and olivine coexist, the 57Fe/54Fe of spinel is greater than that of the corresponding olivine, agreeing with expectations for equilibrium fractionation from theory (ionic model), but disagreeing with predictions based on Mössbauer data. The 57Fe/54Fe values of clinopyroxenes from the xenoliths show no clear systematic differences. We interpret this to be a result of varying degrees of metasomatism, perhaps involving interaction with a melt. The spinel peridotite samples (lherzolite, harzburgite, and dunite) are partially melted residual mantle that exhibit a decrease in whole-rock 57Fe/54Fe with increasing olivine abundance. This is consistent with progressive extraction of a 57Fe-rich partial melt. The clinopyroxenite has the highest whole-rock 57Fe/54Fe, consistent with its origin as a cumulate from an unrelated magma possessing elevated 57Fe/54Fe. The websterite sample is transitional to Group II type xenoliths, has the lowest whole-rock 57Fe/54Fe of the investigated samples, and likely experienced a more complex metasomatic history. This study demonstrates that the Fe isotope compositions of San Carlos xenoliths and their component minerals record the complex petrologic history and local heterogeneity of the subcontinental mantle lithosphere.

Magma redox and structural controls on iron isotope variations in Earth's mantle and crust

The heavy iron isotopic composition of Earth's crust relative to chondrites has been explained by vaporization during the Moon-forming impact, equilibrium partitioning between metal and silicate at core–mantle-boundary conditions, or partial melting and magma differentiation. The latter view is supported by the observed difference in the iron isotopic compositions of MORBS and peridotites. However, the precise controls on iron isotope variations in igneous rocks remain unknown. Here, we show that equilibrium iron isotope fractionation is mainly controlled by redox (Fe3+/Fetot ratio) and structural (e.g., polymerization) conditions in magmas. We measured, for the first time, the mean force constants of iron bonds in silicate glasses by synchrotron Nuclear Resonant Inelastic X-ray Scattering (NRIXS, also known as Nuclear Resonance Vibrational Spectroscopy – NRVS, or Nuclear Inelastic Scattering – NIS). The same samples were studied by conventional Mössbauer and X-ray Absorption Near Edge Structure (XANES) spectroscopy. The NRIXS results reveal a +0.2 to +0.4‰ equilibrium fractionation on 56Fe/54Fe ratio between Fe2+ and Fe3+ end-members in basalt, andesite, and dacite glasses at magmatic temperatures. These first measurements can already explain ∼1/3 of the iron isotopic shift measured in MORBs relative to their source. Further work will be required to investigate how pressure, temperature, and structural differences between melts and glasses affect equilibrium fractionation factors. In addition, large fractionation is also found between rhyolitic glass and commonly occurring oxide and silicate minerals. This fractionation reflects mainly changes in the coordination environment of Fe2+ in rhyolites relative to less silicic magmas and mantle minerals, as also seen by XANES. We provide a new calibration of XANES features vs. Fe3+/Fetot ratio determinations by Mössbauer to estimate Fe3+/Fetot ratio in situ in glasses of basaltic, andesitic, dacitic, and rhyolitic compositions. Modeling of magma differentiation using rhyolite-MELTS shows that iron structural changes in silicic magmas can explain the heavy iron isotopic compositions of granitoids and rhyolites. This study demonstrates that iron stable isotopes can help reveal planetary redox conditions and igneous processes. Other heterovalent elements such as Ti, V, Eu, Cr, Ce, or U may show similar isotopic variations in bulk rocks and individual minerals, which could be used to establish past and present redox condition in the mantles of Earth and other planets.

Determination of the Fe(II)aq–magnetite equilibrium iron isotope fractionation factor using the three-isotope method and a multi-direction approach to equilibrium

Magnetite is ubiquitous in the Earth's crust and its presence in modern marine sediments has been taken as an indicator of biogeochemical Fe cycling. Magnetite is also the most abundant Fe oxide in banded iron formations (BIFs) that have not been subjected to ore-forming alteration. Magnetite is therefore an important target of stable Fe isotope studies, and yet interpretations are currently difficult because of large uncertainties in the equilibrium stable Fe isotope fractionation factors for magnetite relative to fluids and other minerals. In this study, we utilized the three-isotope method (57Fe–56Fe–54Fe) to explore isotopic exchange via an enriched-57Fe tracer, and natural mass-dependent fractionation using 56Fe/54Fe variations, during reaction of aqueous Fe(II) (Fe(II)aq) with magnetite. Importantly, we employed a multi-direction approach to equilibrium by reacting four 57Fe-enriched Fe(II) solutions that had distinct 56Fe/54Fe ratios, which identifies changes in the instantaneous Fe isotope fractionation factor and hence identifies kinetic isotope effects. We find that isotopic exchange can be described by two 56Fe/54Fe fractionations, where an initial rapid exchange (∼66% isotopic mixing within 1 day) involved a relatively small Fe(II)aq–magnetite 56Fe/54Fe fractionation, followed by slower exchange (∼25% isotopic mixing over 50 days) that was associated with a larger Fe(II)aq–magnetite 56Fe/54Fe fractionation; this later fractionation is interpreted to approach isotopic equilibrium between Fe(II)aq and the total magnetite. All four Fe(II) solutions extrapolate to the same final equilibrium 56Fe/54Fe fractionation for Fe(II)aq–magnetite of −1.56±0.20‰(2σ) at 22 °C. Additional experiments that synthesized magnetite via conversion of ferrihydrite by reaction with aqueous Fe(II) yield final 56Fe/54Fe fractionations that are identical to those of the exchange experiments. Our experimental results agree well with calculated fractionation factors using the reduced partition function ratios for Fe(H2O)62+ from Rustad et al. (2010) and stoichiometric magnetite from Mineev et al. (2007), and these relations may be combined with the experimental constraints to determine the temperature dependence of the Fe(II)aq–magnetite fractionation factor:103lnαFe(II)aq–magnetite=−0.145(±0.002)×106/T2+0.10(±0.02) where T is in K.Part of the reason for large discrepancies in calculated Fe isotope fractionation factors for magnetite likely lies in the stoichiometry of the mineral in specific studies, given the significant effect of octahedral versus tetrahedral Fe isotope fractionation that has been calculated. When our results are applied to BIF genesis, our experimentally determined Fe(II)aq–magnetite fractionation factor indicates that magnetite–siderite mineral pairs in ∼2.5 Ga BIFs did not form in Fe isotope equilibrium with each other, or with ancient seawater. Iron oxides in such BIFs are therefore more likely to have formed through processes that were isolated from equilibrium with the oceans, indicating that such BIF minerals may not be suitable proxies for ancient paleoenvironments.

Comment on “New data on equilibrium iron isotope fractionation among sulfides: Constraints on mechanisms of sulfide formation in hydrothermal and igneous systems” by V.B. Polyakov and D.M. Soultanov

Comment on “New data on equilibrium iron isotope fractionation among sulfides: Constraints on mechanisms of sulfide formation in hydrothermal and igneous systems” by V.B. Polyakov and D.M. Soultanov

Response to the comment by M. Blanchard, F. Poitrasson, M. Méheut, M. Lazzeri, F. Mauri, E. Balan on “New data on equilibrium iron isotope fractionation among sulfides: Constraints on mechanisms of sulfide formation in hydrothermal and igneous systems” published in Geochim. Cosmochim. Acta 75 (2011) 1957–1974

Response to the comment by M. Blanchard, F. Poitrasson, M. Méheut, M. Lazzeri, F. Mauri, E. Balan on “New data on equilibrium iron isotope fractionation among sulfides: Constraints on mechanisms of sulfide formation in hydrothermal and igneous systems” published in Geochim. Cosmochim. Acta 75 (2011) 1957–1974

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.

Iron isotope fractionation between pyrite (FeS 2), hematite (Fe 2O 3) and siderite (FeCO 3): A first-principles density functional theory study

In addition to equilibrium isotopic fractionation factors experimentally derived, theoretical predictions are needed for interpreting isotopic compositions measured on natural samples because they allow exploring more easily a broader range of temperature and composition. For iron isotopes, only aqueous species were studied by first-principles methods and the combination of these data with those obtained by different methods for minerals leads to discrepancies between theoretical and experimental isotopic fractionation factors. In this paper, equilibrium iron isotope fractionation factors for the common minerals pyrite, hematite, and siderite were determined as a function of temperature, using first-principles methods based on the density functional theory (DFT). In these minerals belonging to the sulfide, oxide and carbonate class, iron is present under two different oxidation states and is involved in contrasted types of interatomic bonds. Equilibrium fractionation factors calculated between hematite and siderite compare well with the one estimated from experimental data (ln α 57Fe/ 54Fe = 4.59 ± 0.30‰ and 5.46 ± 0.63‰ at 20 °C for theoretical and experimental data, respectively) while those for Fe(III) aq-hematite and Fe(II) aq-siderite are significantly higher that experimental values. This suggests that the absolute values of the reduced partition functions ( β-factors) of aqueous species are not accurate enough to be combined with those calculated for minerals. When compared to previous predictions derived from Mössbauer or INRXS data [Polyakov V. B., Clayton R. N., Horita J. and Mineev S. D. (2007) Equilibrium iron isotope fractionation factors of minerals: reevaluation from the data of nuclear inelastic resonant X-ray scattering and Mössbauer spectroscopy. Geochim. Cosmochim. Acta 71, 3833–3846], our iron β-factors are in good agreement for siderite and hematite while a discrepancy is observed for pyrite. However, the detailed investigation of the structural, electronic and vibrational properties of pyrite as well as the study of sulfur isotope fractionation between pyrite and two other sulfides (sphalerite and galena) indicate that DFT-derived β-factors of pyrite are as accurate as for hematite and siderite. We thus suggest that experimental vibrational density of states of pyrite should be re-examined.

Equilibrium Iron Isotope Fractionation at Core-Mantle Boundary Conditions

The equilibrium iron isotope fractionation between lower mantle minerals and metallic iron at core-mantle boundary conditions can be evaluated from the high-pressure 57Fe partial vibrational density of states determined by synchrotron inelastic nuclear resonant x-ray scattering spectroscopy using a diamond anvil. Ferropericlase [(Mg,Fe)O] and (Fe,Mg)SiO3- post-perovskite are enriched in heavy iron isotopes relative to metallic iron at ultrahigh pressures, as opposed to the equilibrium iron isotope fractionation between these compounds at low pressure. The enrichment of Earth and Moon basalts in heavy iron isotopes relative to those from Mars and asteroid Vesta can be explained by the equilibrium iron isotope fractionation during the segregation of Earth's core and the assumption that Earth was already differentiated before the Moon-forming "giant impact."

Experimental studies of equilibrium iron isotope fractionation in ferric aquo–chloro complexes

Here we compare new experimental studies with theoretical predictions of equilibrium iron isotopic fractionation among aqueous ferric chloride complexes (Fe(H 2O) 6 3+, FeCl(H 2O) 5 2+, FeCl 2(H 2O) 4 +, FeCl 3 (H 2O) 3, and FeCl 4 –), using the Fe–Cl–H 2O system as a simple, easily-modeled example of the larger variety of iron–ligand compounds, such as chlorides, sulfides, simple organic acids, and siderophores. Isotopic fractionation ( 56Fe/ 54Fe) among naturally occuring iron-bearing species at Earth surface temperatures (up to ∼3‰) is usually attributed to redox effects in the environment. However, theoretical modeling of reduced isotopic partition functions among iron-bearing species in solution also predicts fractionations of similar magnitude due to non-redox changes in speciation (i.e., ligand bond strength and coordination number). In the present study, fractionations are measured in a series of low pH ([H +] = 5 M) solutions of ferric chloride (total Fe = 0.0749 mol/L) at chlorinities ranging from 0.5 to 5.0 mol/L. Advantage is taken of the unique solubility of FeCl 4 – in immiscible diethyl ether to create a separate spectator phase, used to monitor changing fractionation in the aqueous solution. Δ 56Fe aq-eth = δ 56Fe (total Fe remaining in aqueous phase)−δ 56Fe (FeCl 4 – in ether phase) is determined for each solution via MC-ICPMS analysis. Both experiments and theoretical calculations of Δ 56Fe aq-eth show a downward trend with increasing chlorinity: Δ 56Fe aq-eth is greatest at low chlorinity, where FeCl 2(H 2O) 4 + is the dominant species, and smallest at high chlorinity where FeCl 3(H 2O) 3 is dominant. The experimental Δ 56Fe aq-eth ranges from 0.8‰ at [Cl –] = 0.5 M to 0.0‰ at [Cl –] = 5.0 M, a decrease in aqueous–ether fractionation of 0.8‰. This is very close to the theoretically predicted decreases in Δ 56Fe aq-eth, which range from 1.0 to 0.7‰, depending on the ab initio model. The rate of isotopic exchange and attainment of equilibrium are shown using spiked reversal experiments in conjunction with the two-phase aqueous–ether system. Equilibrium under the experimental conditions is established within 30 min. The general agreement between theoretical predictions and experimental results points to substantial equilibrium isotopic fractionation among aqueous ferric chloride complexes and a decrease in 56Fe/ 54Fe as the Cl –/Fe 3+ ion ratio increases. The effects on isotopic fractionation shown by the modeling of this simple iron–ligand system imply that ligands present in an aqueous environment are potentially important drivers of fractionation, are indicative of possible fractionation effects due to other speciation effects (such as iron–sulfide systems or iron bonding with organic ligands), and must be considered when interpreting iron isotope fractionation in the geological record.

Modeling the effects of bond environment on equilibrium iron isotope fractionation in ferric aquo-chloro complexes

Four or five sets of ab initio models, including Unrestricted Hartree Fock (UHF) and hybrid Density Functional Theory (DFT) are calculated for each species in a series of aqueous ferric aquo-chloro complexes: Fe ( H 2 O ) 6 3 + , FeCl ( H 2 O ) 5 2 + , FeCl 2 ( H 2 O ) 4 + , FeCl 3(H 2O) 3, FeCl 3(H 2O) 2, FeCl 4 - , FeCl 5H 2O 2−, FeCl 5 2 - , FeCl 6 3 - ) in order to determine the relative isotopic fractionation among the complexes, to compare the results of different models for the same complexes, to examine factors that influence the magnitude of the isotopic fractionation, and to compare bond-partner-driven fractionation with redox-driven fractionation. Relative to Fe ( H 2 O ) 6 3 + , all models show a nearly linear decrease in 56Fe/ 54Fe as the number of Cl − ions per Fe 3+ ion increases, with slopes of −0.8‰ to −1.0‰ per Cl − at 20 °C. At 20 °C, 1000 ln β ( β = 56Fe/ 54Fe reduced partition function ratio relative to a dissociated Fe atom) values range from 8.93‰ to 9.73‰ for Fe ( H 2 O ) 6 3 + , 8.04–9.12‰ for FeCl ( H 2 O ) 5 2 + , 7.61–8.73‰ for FeCl 2 ( H 2 O ) 4 + , 7.14–8.25‰ for FeCl 4 - , and 3.09–4.41‰ for FeCl 6 3 - . The fractionation between Fe ( H 2 O ) 6 3 + and FeCl 4 - ranges from 1.5‰ to 2.6‰, depending on the model; this is comparable in magnitude to fractionation effects due to Fe 3+/Fe 2+ redox reactions. β values from the UHF models are consistently higher than those from the hybrid DFT models. Isotopic fractionation is shown to be sensitive to differences in ligand bond stiffness (above), coordination number, bond length, and the frequency of the asymmetric Fe–X stretching vibrational mode, as predicted by previous theoretical studies. Complexes with smaller coordination numbers have higher 1000 ln β (7.46‰, 5.25‰, and 3.48‰ for FeCl 4 - , FeCl 5 2 - , FeCl 6 3 - , respectively, from the B3LYP/6-31G(d) model). Species with the same number of chlorides but fewer waters also show the effect of coordination number on 1000 ln β: (7.46‰ vs. 7.05‰ for FeCl 3(H 2O) 2 vs. FeCl 3(H 2O) 3 and 5.25‰ vs. 4.94‰ for FeCl 5 2 - vs. FeCl 5H 2O 2− with the B3LYP/6-31G(d) model). As more Fe–Cl bonds substitute for Fe–OH 2 bonds (with a resulting decrease in β), the lengths of the Fe–Cl bonds and the Fe–O bonds increase. Preliminary modeling of Fe ( H 2 O ) 6 2 + shows an Fe 3+/Fe 2+ fractionation of 3.2‰ for the B3LYP/6-31G(d) model, in agreement with previous studies. The addition of an explicit outer hydration sphere of 12 H 2O molecules to models of Fe ( H 2 O ) 6 3 + improves agreement with measured vibrational frequencies and bond lengths; 1000 ln β increases by 0.8–1.0‰. An additional hydration sphere around FeCl 4 - increases 1000 ln β by only 0.1‰. Isotopic fractionations predicted for this simple system imply that ligands present in an aqueous iron environment are potentially important drivers of fractionation, and suggest that significant fractionation effects are likely in other aqueous systems containing sulfides or organic ligands. Fractionation effects due to both speciation and redox must be considered when interpreting iron isotope fractionations in the geological record.

Equilibrium iron isotope fractionation factors of minerals: Reevaluation from the data of nuclear inelastic resonant X-ray scattering and Mössbauer spectroscopy

We have critically reevaluated equilibrium iron isotope fractionation factors for oxide and sulfide minerals using recently acquired data obtained by Mössbauer spectroscopy and inelastic nuclear resonant X-ray scattering (INRXS) synchrotron radiation. Good agreement was observed in the iron β-factors of metallic iron (α-Fe) and hematite calculated using both Mössbauer- and INRXS-derived data, which supports the validity and reliability of the calculations. Based on this excellent agreement, we suggest the use of the present data on the iron β-factors of hematite as a reference. The previous Mössbauer-derived iron β-factor for magnetite has been modified significantly based on the Fe-sublattice density of states obtained from the INRXS experiments. This resolves the disagreement between naturally observed iron isotope fractionation factors for mineral pairs involving magnetite and those obtained from the calculated β-factors. The correctness of iron β-factor for pyrite has been corroborated by the good agreement with experimental data of sulfur isotope geothermometers of pyrite–galena and pyrite–sphalerite. A good correlation between the potential energy of the cation site, the oxidation state of iron and the iron β-factor value has been established. Specifically, ferric compounds, which have a higher potential energy of iron than ferrous compounds, have higher β-factors. A similar dependence of β-factors on the oxidation state and potential energy could be extended to other transition metals. Extremely low values of INRXS-derived iron β-factors for troilite and Fe 3S significantly widen the range of iron β-factors for covalently bonded compounds.

The experimental calibration of the iron isotope fractionation factor between pyrrhotite and peralkaline rhyolitic melt

A first experimental study was conducted to determine the equilibrium iron isotope fractionation between pyrrhotite and silicate melt at magmatic conditions. Experiments were performed in an internally heated gas pressure vessel at 500 MPa and temperatures between 840 and 1000 °C for 120–168 h. Three different types of experiments were conducted and after phase separation the iron isotope composition of the run products was measured by MC-ICP-MS. (i) Kinetic experiments using 57Fe-enriched glass and natural pyrrhotite revealed that a close approach to equilibrium is attained already after 48 h. (ii) Isotope exchange experiments—using mixtures of hydrous peralkaline rhyolitic glass powder (∼4 wt% H 2O) and natural pyrrhotites (Fe 1 − x S) as starting materials— and (iii) crystallisation experiments, in which pyrrhotite was formed by reaction between elemental sulphur and rhyolitic melt, consistently showed that pyrrhotite preferentially incorporates light iron. No temperature dependence of the fractionation factor was found between 840 and 1000 °C, within experimental and analytical precision. An average fractionation factor of Δ 56Fe/ 54Fe pyrrhotite-melt = −0. 35 ± 0.04‰ (2SE, n = 13) was determined for this temperature range. Predictions of Fe isotope fractionation between FeS and ferric iron-dominated silicate minerals are consistent with our experimental results, indicating that the marked contrast in both ligand and redox state of iron control the isotope fractionation between pyrrhotite and silicate melt. Consequently, the fractionation factor determined in this study is representative for the specific Fe 2+/ΣFe ratio of our peralkaline rhyolitic melt of 0.38 ± 0.02. At higher Fe 2+/ΣFe ratios a smaller fractionation factor is expected. Further investigation on Fe isotope fractionation between other mineral phases and silicate melts is needed, but the presented experimental results already suggest that even at high temperatures resolvable variations in the Fe isotope composition can be generated by equilibrium isotope fractionation in natural magmatic systems.

Theoretical investigation of iron isotope fractionation between Fe(H 2O) 63+ and Fe(H 2O) 62+: Implications for iron stable isotope geochemistry

The magnitude of equilibrium iron isotope fractionation between Fe(H 2O) 6 3+ and Fe(H 2O) 6 2+ is calculated using density functional theory (DFT) and compared to prior theoretical and experimental results. DFT is a quantum chemical approach that permits a priori estimation of all vibrational modes and frequencies of these complexes and the effects of isotopic substitution. This information is used to calculate reduced partition function ratios of the complexes (10 3 · ln(β)), and hence, the equilibrium isotope fractionation factor (10 3 · ln(α)). Solvent effects are considered using the polarization continuum model (PCM). DFT calculations predict fractionations of several per mil in 56Fe/ 54Fe favoring partitioning of heavy isotopes in the ferric complex. Quantitatively, 10 3 · ln(α) predicted at 22°C, ∼ 3 ‰, agrees with experimental determinations but is roughly half the size predicted by prior theoretical results using the Modified Urey-Bradley Force Field (MUBFF) model. Similar comparisons are seen at other temperatures. MUBFF makes a number of simplifying assumptions about molecular geometry and requires as input IR spectroscopic data. The difference between DFT and MUBFF results is primarily due to the difference between the DFT-predicted frequency for the ν 4 mode (O-Fe-O deformation) of Fe(H 2O) 6 3+ and spectroscopic determinations of this frequency used as input for MUBFF models (185–190 cm −1 vs. 304 cm −1, respectively). Hence, DFT-PCM estimates of 10 3 · ln(β) for this complex are ∼ 20% smaller than MUBFF estimates. The DFT derived values can be used to refine predictions of equilibrium fractionation between ferric minerals and dissolved ferric iron, important for the interpretation of Fe isotope variations in ancient sediments. Our findings increase confidence in experimental determinations of the Fe(H 2O) 6 3+ − Fe(H 2O) 6 2+ fractionation factor and demonstrate the utility of DFT for applications in “heavy” stable isotope geochemistry.

Comment on “Isotopic fractionation between Fe(III) and Fe(II) in aqueous solutions” by Clark Johnson et al., [Earth Planet. Sci. Lett. 195 (2002) 141–153

In a recent contribution [1], Johnson et al. reported the equilibrium isotope fractionation factor between dissolved Fe(II) and Fe(III) in aqueous solutions at pH=2.5 and 5.5. They suggest that because the iron isotope fractionation observed in their experiments spans virtually the entire range observed in sedimentary rocks, Fe(II)^ Fe(III) aqueous speciation may play a major role in determining iron isotope variations in nature where Fe(II) and Fe(III) can become physically separated. They discounted earlier conclusions by us and others [2,3] that signi¢cant equilibrium fractionation between speci¢c coexisting Fe(II)or Fe(III)-aqueous complexes (e.g., between aqueous Fe(II)(OH)xðaqÞ and Fe(II)ðaqÞ ion) is capable of producing iron isotope contrasts that can be preserved in nature. This is an important contribution not only because the authors recognize the importance of abiotic equilibrium iron isotope fractionation in nature in contrast to previous assertions [4], but also because it will help to focus discussion on the development and evaluation of experimental approaches that can reveal abiotic fractionation mechanisms. However, in this Comment we propose that the experiments presented in this paper cannot be interpreted as straightforwardly as Johnson et al. contend. In particular, we show that in one of their critical experiments attainment of either isotope mass balance or equilibrium was not demonstrated, and thus the results of that experiment cannot be used to calculate an Fe(II)^Fe(III) equilibrium fractionation factor. Johnson et al. determined isotopic fractionation between aqueous Fe(II) and Fe(III) based on two experiments in which isotopically ‘normal’ Fe(II)and Fe(III)-chloride solutions were mixed and presumed to attain isotopic equilibrium at pH=2.5 and 5.5. The time allowed for attainment of isotopic equilibrium was constrained by a third experiment in which isotopically ‘enriched’ Fe(III)-chloride solution was mixed with isotopically ‘normal’ Fe(II)-chloride solution at pH=2.5, and a time series of Fe(II) and Fe(III) isotopic compositions approaching equilibrium were ¢t with a second order rate function. This latter experiment demonstrated the approach to isotopic equilibrium through electron transfer between coexisting Fe(II)and Fe(III)-aqueous complexes in the absence of oxidation through the ferrous hydroxyl pathway that is common under higher pH, aerobic conditions. We point out that the ¢tted rate function was for conditions where aqueous