Effects of changing solution chemistry on Fe/Fe 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 Fe²⁺-Fe³⁺-Cl-H₂O solution as the chlorinity was varied. We made use of the unique solubility of FeCl₄⁻ 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 Fe²⁺- and Fe³⁺-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(H₂O)₆²⁺, 5.98‰ to 5.34‰ for FeCl(H₂O)₅⁺, and 5.91‰ to 4.86‰ for FeCl₂(H₂O)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.     

First experimental determination of iron isotope fractionation between hematite and aqueous solution at hydrothermal conditions

Although iron isotopes provide a new powerful tool for tracing a variety of geochemical processes, the unambiguous interpretation of iron isotope ratios in natural systems and the development of predictive theoretical models require accurate data on equilibrium isotope fractionation between fluids and minerals. We investigated Fe isotope fractionation between hematite (Fe₂O₃) and aqueous acidic NaCl fluids via hematite dissolution and precipitation experiments at temperatures from 200 to 450 °C and pressures from saturated vapor pressure (P_sat) to 600 bar. Precipitation experiments at 200 °C and P_sat from aqueous solution, in which Fe aqueous speciation is dominated by ferric iron (Fe^III) chloride complexes, show no detectable Fe isotope fractionation between hematite and fluid, Δ⁵⁷Fe_{fluid-hematite} = δ⁵⁷Fe_fluid − δ⁵⁷Fe_hematite = 0.01 ± 0.08‰ (2 × standard error, 2SE). In contrast, experiments at 300 °C and P_sat, where ferrous iron chloride species (FeCl₂ and FeCl⁺) dominate in the fluid, yield significant fluid enrichment in the light isotope, with identical values of Δ⁵⁷Fe_{fluid-hematite} = −0.54 ± 0.15‰ (2SE) both for dissolution and precipitation runs. Hematite dissolution experiments at 450 °C and 600 bar, in which Fe speciation is also dominated by ferrous chloride species, yield Δ⁵⁷Fe_{fluid-hematite} values close to zero within errors, 0.15 ± 0.17‰ (2SE). In most experiments, chemical, redox, and isotopic equilibrium was attained, as shown by constancy over time of total dissolved Fe concentrations, aqueous Fe^II and Fe^III fractions, and Fe isotope ratios in solution, and identical Δ⁵⁷Fe values from dissolution and precipitation runs. Our measured equilibrium Δ⁵⁷Fe_{fluid-hematite} values at different temperatures, fluid compositions and iron redox state are within the range of fractionations in the system fluid-hematite estimated using reported theoretical β-factors for hematite and aqueous Fe species and the distribution of Fe aqueous complexes in solution. These theoretical predictions are however affected by large discrepancies among different studies, typically ±1‰ for the Δ⁵⁷Fe _{Fe(aq)-hematite} value at 200 °C. Our data may thus help to refine theoretical models for β-factors of aqueous iron species. This study provides the first experimental calibration of Fe isotope fractionation in the system hematite-saline aqueous fluid at elevated temperatures; it demonstrates the importance of redox control on Fe isotope fractionation at hydrothermal conditions.     

Iron and magnesium isotopic compositions of peridotite xenoliths from Eastern China

We present high-precision measurements of Mg and Fe isotopic compositions of olivine, orthopyroxene (opx), and clinopyroxene (cpx) for 18 lherzolite xenoliths from east central China and provide the first combined Fe and Mg isotopic study of the upper mantle. δ⁵⁶Fe in olivines varies from 0.18‰ to −0.22‰ with an average of −0.01 ± 0.18‰ (2SD, n = 18), opx from 0.24‰ to −0.22‰ with an average of 0.04 ± 0.20‰, and cpx from 0.24‰ to −0.16‰ with an average of 0.10 ± 0.19‰. δ²⁶Mg of olivines varies from −0.25‰ to −0.42‰ with an average of −0.34 ± 0.10‰ (2SD, n = 18), opx from −0.19‰ to −0.34‰ with an average of −0.25 ± 0.10‰, and cpx from −0.09‰ to −0.43‰ with an average of −0.24 ± 0.18‰. Although current precision (∼±0.06‰ for δ⁵⁶Fe; ±0.10‰ for δ²⁶Mg, 2SD) limits the ability to analytically distinguish inter-mineral isotopic fractionations, systematic behavior of inter-mineral fractionation for both Fe and Mg is statistically observed: Δ⁵⁶Fe_ol-cpx = −0.10 ± 0.12‰ (2SD, n = 18); Δ⁵⁶Fe_ol-opx = −0.05 ± 0.11‰; Δ²⁶Mg_ol-opx = −0.09 ± 0.12‰; Δ²⁶Mg_ol-cpx = −0.10 ± 0.15‰. Fe and Mg isotopic composition of bulk rocks were calculated based on the modes of olivine, opx, and cpx. The average δ⁵⁶Fe of peridotites in this study is 0.01 ± 0.17‰ (2SD, n = 18), similar to the values of chondrites but slightly lower than mid-ocean ridge basalts (MORB) and oceanic island basalts (OIB). The average δ²⁶Mg is −0.30 ± 0.09‰, indistinguishable from chondrites, MORB, and OIB. Our data support the conclusion that the bulk silicate Earth (BSE) has chondritic δ⁵⁶Fe and δ²⁶Mg.The origin of inter-mineral fractionations of Fe and Mg isotopic ratios remains debated. δ⁵⁶Fe between the main peridotite minerals shows positive linear correlations with slopes within error of unity, strongly suggesting intra-sample mineral-mineral Fe and Mg isotopic equilibrium. Because inter-mineral isotopic equilibrium should be reached earlier than major element equilibrium via chemical diffusion at mantle temperatures, Fe and Mg isotope ratios of coexisting minerals could be useful tools for justifying mineral thermometry and barometry on the basis of chemical equilibrium between minerals. Although most peridotites in this study exhibit a narrow range in δ⁵⁶Fe, the larger deviations from average δ⁵⁶Fe for three samples likely indicate changes due to metasomatic processes. Two samples show heavy δ⁵⁶Fe relative to the average and they also have high La/Yb and total Fe content, consistent with metasomatic reaction between peridotite and Fe-rich and isotopically heavy melt. The other sample has light δ⁵⁶Fe and slightly heavy δ²⁶Mg, which may reflect Fe-Mg inter-diffusion between peridotite and percolating melt.     

Iron isotope fractionation during leaching of granite and basalt by hydrochloric and oxalic acids

Transport of iron (Fe) within hydrothermal and soil environments involves the transferral into aqueous solutions by leaching of complex, polyminerallic rocks. Understanding the isotope fractionation mechanisms during this process is key for any application of the Fe-isotope system to biogeochemical studies. Here, we reacted biotite granite and tholeiite-basalt with 0.5 M hydrochloric acid and 5 mM oxalic acid solutions at ambient temperature. Solution aliquots were recovered over a seven-day period and analysed for major and trace element concentrations and Fe isotopic compositions. In all experiments, Fe initially released into solution was isotopically lighter, with Δ⁵⁶Fe_{solution-rock} as low as −1.80‰ in the granite-hydrochloric acid system. The oxalic acid experiments showed similar patterns but smaller fractionation. In all experiments, the Δ⁵⁶Fe_{solution-rock} reduced over time, which would be in line with the formation of a leached layer as proposed before [Brantley S. L., Liermann L. J., Guynn R. L., Anbar A., Icopini G. A., and Barling J. (2004) Fe isotopic fractionation during mineral dissolution with and without bacteria. Geochim. Cosmochim. Acta68(15), 3189-3204]. Granite and basalts reacting with hydrochloric acid reached apparent steady-state values of −0.60 ± 0.15‰ and −0.40 ± 0.20‰, respectively, whilst experimental values with oxalic acid were −1.0 ± 0.15‰ and −0.50 ± 0.15‰. During the granite experiments, alteration of biotite to chlorite, followed by dissolution of chlorite, were likely the dominant processes, whilst in the basalt experiments, dissolution of pigeonite was likely the principal source of Fe. Variations in pH during the hydrochloric acid experiments were minimal, remaining below 0.5 at all times. In oxalic acid solutions, the pH increased to over 4, leading likely to precipitation of secondary minerals and adsorption/co-precipitation of Fe onto mineral surfaces. These processes could contribute to the greater fractionation observed in the final stages of the oxalic acid experiments. Our results highlight the importance of mineralogy and fluid composition on the Fe-isotope systematics during weathering. The fractionation processes identified for granites and basalts are in line with those inferred from field observations in soils, sediments, groundwater and hydrothermal deposits and from laboratory studies of single-mineral leaching.     

Kinetic and equilibrium Fe isotope fractionation between aqueous Fe(III) and hematite

Application of the Fe isotope system to studies of natural rocks and fluids requires precise knowledge of equilibrium Fe isotope fractionation factors among various aqueous Fe species and minerals. These are difficult to obtain at the low temperatures at which Fe isotope fractionation is expected to be largest and requires careful distinction between kinetic and equilibrium isotope effects. A detailed investigation of Fe isotope fractionation between [Fe^III(H₂O)₆]³⁺ and hematite at 98°C allows the equilibrium ⁵⁶Fe/⁵⁴Fe fractionation to be inferred, which we estimate at 10³lnα_{Fe(III)-hematite} = −0.10 ± 0.20‰. We also infer that the slope of Fe(III)-hematite fractionation is modest relative to 10⁶/T², which would imply that this fractionation remains close to zero at lower temperatures. These results indicate that Fe isotope compositions of hematite may closely approximate those of the fluids from which they precipitated if equilibrium isotopic fractionation is assumed, allowing inference of δ⁵⁶Fe values of ancient fluids from the rock record. The equilibrium Fe(III)-hematite fractionation factor determined in this study is significantly smaller than that obtained from the reduced partition function ratios calculated for [Fe^III(H₂O)₆]³⁺ and hematite based on vibrational frequencies and Mössbauer shifts by [Polyakov 1997] and [Polyakov and Mineev 2000], and Schauble et al. (2001), highlighting the importance of experimental calibration of Fe isotope fractionation factors. In contrast to the long-term (up to 203 d) experiments, short-term experiments indicate that kinetic isotope effects dominate during rapid precipitation of ferric oxides. Precipitation of hematite over ∼12 h produces a kinetic isotope fractionation where 10³lnα_{Fe(III)-hematite} = +1.32 ± 0.12‰. Precipitation under nonequilibrium conditions, however, can be recognized through stepwise dissolution in concentrated acids. As expected, our results demonstrate that dissolution by itself does not measurably fractionate Fe isotopes.     

Inter-mineral Fe isotope variations in mantle-derived rocks and implications for the Fe geochemical cycle

Iron isotope and major- and minor-element compositions of coexisting olivine, clinopyroxene, and orthopyroxene from eight spinel peridotite mantle xenoliths; olivine, magnetite, amphibole, and biotite from four andesitic volcanic rocks; and garnet and clinopyroxene from seven garnet peridotite and eclogites have been measured to evaluate if inter-mineral Fe isotope fractionation occurs in high-temperature igneous and metamorphic minerals and if isotopic fractionation is related to equilibrium Fe isotope partitioning or a result of open-system behavior. There is no measurable fractionation between silicate minerals and magnetite in andesitic volcanic rocks, nor between olivine and orthopyroxene in spinel peridotite mantle xenoliths. There are some inter-mineral differences (up to 0.2‰ in ⁵⁶Fe/⁵⁴Fe) in the Fe isotope composition of coexisting olivine and clinopyroxene in spinel peridotites. The Fe isotope fractionation observed between clinopyroxene and olivine appears to be a result of open-system behavior based on a positive correlation between the Δ⁵⁶Fe_{clinopyroxene-olivine} fractionation and the δ⁵⁶Fe value of clinopyroxene and olivine. There is also a significant difference in the isotopic compositions of garnet and clinopyroxene in garnet peridotites and eclogites, where the average Δ⁵⁶Fe_{clinopyroxene-garnet} fractionation is +0.32 ± 0.07‰ for six of the seven samples. The one sample that has a lower Δ⁵⁶Fe_{clinopyroxene-garnet} fractionation of 0.08‰ has a low Ca content in garnet, which may reflect some crystal chemical control on Fe isotope fractionation. The Fe isotope variability in mantle-derived minerals is interpreted to reflect subduction of isotopically variable oceanic crust, followed by transport through metasomatic fluids. Isotopic variability in the mantle might also occur during crystal fractionation of basaltic magmas within the mantle if garnet is a liquidus phase. The isotopic variations in the mantle are apparently homogenized during melting processes, producing homogenous Fe isotope compositions during crust formation.     

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: , , , FeCl₃(H₂O)₃, FeCl₃(H₂O)₂, , FeCl₅H₂O²⁻, , ) 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 , all models show a nearly linear decrease in ⁵⁶Fe/⁵⁴Fe as the number of Cl⁻ ions per Fe³⁺ 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 , 8.04-9.12‰ for , 7.61-8.73‰ for , 7.14-8.25‰ for , and 3.09-4.41‰ for . The fractionation between and ranges from 1.5‰ to 2.6‰, depending on the model; this is comparable in magnitude to fractionation effects due to Fe³⁺/Fe²⁺ 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 , ,, 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 FeCl₃(H₂O)₂ vs. FeCl₃(H₂O)₃ and 5.25‰ vs. 4.94‰ for vs. FeCl₅H₂O²⁻ with the B3LYP/6-31G(d) model). As more Fe-Cl bonds substitute for Fe-OH₂ bonds (with a resulting decrease in β), the lengths of the Fe-Cl bonds and the Fe-O bonds increase.Preliminary modeling of shows an Fe³⁺/Fe²⁺ 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 H₂O molecules to models of improves agreement with measured vibrational frequencies and bond lengths; 1000 ln β increases by 0.8-1.0‰. An additional hydration sphere around 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.     

Copper stable isotopes as tracers of metal-sulphide segregation and fractional crystallisation processes on iron meteorite parent bodies

We report high precision Cu isotope data coupled with Cu concentration measurements for metal, troilite and silicate fractions separated from magmatic and non-magmatic iron meteorites, analysed for Fe isotopes (δ⁵⁷Fe; permil deviation in ⁵⁷Fe/⁵⁴Fe relative to the pure iron standard IRMM-014) in an earlier study (Williams et al., 2006). The Cu isotope compositions (δ⁶⁵Cu; permil deviation in ⁶⁵Cu/⁶³Cu relative to the pure copper standard NIST 976) of both metals (δ⁶⁵Cu_M) and sulphides (δ⁶⁵Cu_FeS) span much wider ranges (−9.30 to 0.99‰ and −8.90 to 0.63‰, respectively) than reported previously. Metal-troilite fractionation factors (Δ⁶⁵Cu_M-FeS = δ⁶⁵Cu_M − δ⁶⁵Cu_FeS) are variable, ranging from −0.07 to 5.28‰, and cannot be explained by equilibrium stable isotope fractionation coupled with either mixing or reservoir effects, i.e. differences in the relative proportions of metal and sulphide in the meteorites. Strong negative correlations exist between troilite Cu and Fe (δ⁵⁷Fe_FeS) isotope compositions and between metal-troilite Cu and Fe (Δ⁵⁷Fe_M-FeS) isotope fractionation factors, for both magmatic and non-magmatic irons, which suggests that similar processes control isotopic variations in both systems. Clear linear arrays between δ⁶⁵Cu_FeS and δ⁵⁷Fe_FeS and calculated Cu metal-sulphide partition coefficients (D_Cu = [Cu]_metal/[Cu]_FeS) are also present. A strong negative correlation exists between Δ⁵⁷Fe_M-FeS and D_Cu; a more diffuse positive array is defined by Δ⁶⁵Cu_M-FeS and D_Cu. The value of D_Cu can be used to approximate the degree of Cu concentration equilibrium as experimental studies constrain the range of D_Cu between Fe metal and FeS at equilibrium to be in the range of 0.05-0.2; D_Cu values for the magmatic and non-magmatic irons studied here range from 0.34 to 1.11 and from 0.04 to 0.87, respectively. The irons with low D_Cu values (closer to Cu concentration equilibrium) display the largest Δ⁵⁷Fe_M-FeS and the lowest Δ⁶⁵Cu_M-FeS values, whereas the converse is observed in the irons with large values D_Cu that deviate most from Cu concentration equilibrium. The magnitudes of Cu and Fe isotope fractionation between metal and FeS in the most equilibrated samples are similar: 0.25 and 0.32‰/amu, respectively. As proposed in an earlier study (Williams et al., 2006) the range in Δ⁵⁷Fe_M-FeS values can be explained by incomplete Fe isotope equilibrium between metal and sulphide during cooling, where the most rapidly-cooled samples are furthest from isotopic equilibrium and display the smallest Δ⁵⁷Fe_M-FeS and largest D_Cu values. The range in Δ⁶⁵Cu_M-FeS, however, reflects the combined effects of partial isotopic equilibrium overprinting an initial kinetic signature produced by the diffusion of Cu from metal into exsolving sulphides and the faster diffusion of the lighter isotope. In this scenario, newly-exsolved sulphides initially have low Cu contents (i.e. high D_Cu) and extremely light δ⁶⁵Cu_FeS values; with progressive equilibrium and fractional crystallisation the Cu contents of the sulphides increase as their isotopic composition becomes less extreme and closer to the metal value. The correlation between Δ⁶⁵Cu_M-FeS and Δ⁵⁷Fe_M-FeS is therefore a product of the superimposed effects of kinetic fractionation of Cu and incomplete equilibrium between metal and sulphide for both isotope systems during cooling. The correlations between Δ⁶⁵Cu_M-FeS and Δ⁵⁷Fe_M-FeS are defined by both magmatic and non-magmatic irons record fractional crystallisation and cooling of metallic melts on their respective parent bodies as sulphur and chalcophile elements become excluded from crystallised solid iron and concentrated in the residual melt. Fractional crystallisation processes at shallow levels have been implicated in the two main classes of models for the origin of the non-magmatic iron meteorites; at (i) shallow levels in impact melt models and (ii) at much deeper levels in models where the non-magmatic irons represent metallic melts that crystallised within the interior of a disrupted and re-aggregated parent body. The presence of non-magmatic irons with a range of Fe and Cu isotope compositions, some of which record near-complete isotopic equilibrium implies crystallisation at a range of cooling rates and depths, which is most consistent with cooling within the interior of a meteorite parent body. Our data therefore lend support to models where the non-magmatic irons are metallic melts that crystallised in the interior of re-aggregated, partially differentiated parent bodies.     

Grain-scale iron isotopic distribution of pyrite from Precambrian shallow marine carbonate revealed by a femtosecond laser ablation multicollector ICP-MS technique: Possible proxy for the redox state of ancient seawater

The redox state of Precambrian shallow seas has been linked with material cycle and evolution of the photosynthesis-based ecosystem. Iron is a redox-sensitive element and exists as a soluble Fe(II) species or insoluble Fe(III) species on Earth’s surface. Previous studies have shown that the iron isotopic ratio of marine sedimentary minerals is useful for understanding the ocean redox state, although the redox state of the Archean shallow sea is poorly known. This is partly because the conventional bulk isotope analytical technique has often been used, wherein the iron isotopic record may be dampened by the presence of isotopically different iron-bearing minerals within the same sample. Here we report a microscale iron isotopic ratio of individual pyrite grains in shallow marine stromatolitic carbonates over geological time using a newly developed, near-infrared femtosecond laser ablation multicollector ICP-MS technique (NIR-fs-LA-MC-ICP-MS).We have determined that the grain-scale iron isotopic distribution of pyrite from coeval samples shows a bimodal (2.7 and 2.3 Ga) or unimodal pattern (2.9, 2.6, and 0.7 Ga). In particular, pyrite from the 2.7 Ga Fortescue Group shows a unique bimodal distribution with highly positive (+1.0‰ defined as Type 1) and negative δ⁵⁶Fe values (−1.8‰ defined as Type 2). Type 1 and 2 pyrites occasionally occur within different siliceous layers in the same rock specimen. Layer-scale iron isotopic heterogeneity indicates that the iron isotopic ratios of the two types of pyrite are not homogenized by diagenesis after deposition. Some cubic pyrites have a core with a positive δ⁵⁶Fe value (1‰) and a rim with a crustal δ⁵⁶Fe value (0‰). The observed isotopic zoning suggests that the positive δ⁵⁶Fe value is a primary signature at the time of stromatolite formation, while secondary pyrite precipitated during diagenesis.The positive δ⁵⁶Fe value of Type 1 and the large iron isotopic difference between Type 1 and 2 (2.8‰.) suggest partial Fe(II) oxidation in the 2.7-Ga shallow sea, i.e., pyritization of ⁵⁶Fe-enriched ferric oxyhydroxide (Type 1) and ⁵⁶Fe depleted Fe²⁺aq in seawater (Type 2). Type 2 pyrite was probably not produced by microbial iron redox cycling during diagenesis because this scenario requires a higher abundance of pyrite with δ⁵⁶Fe of 0‰ than of −1.8‰. Consequently, the degree of Fe(II) oxidation in the 2.7-Ga shallow sea can be estimated by a Fe²⁺aq steady-state model. The model calculation shows that half the Fe²⁺aq influx was oxidized in the seawater. This implies that O₂ produced by photosynthesis would have been completely consumed by oxidation of the Fe²⁺aq influx. Grain-scale iron isotopic distribution of pyrite could be a useful index for reconstructing the redox state of the Archean shallow sea.     

Experimental constraints on Fe isotope fractionation during magnetite and Fe carbonate formation coupled to dissimilatory hydrous ferric oxide reduction

Iron isotope fractionation between aqueous Fe(II) and biogenic magnetite and Fe carbonates produced during reduction of hydrous ferric oxide (HFO) by Shewanella putrefaciens, Shewanella algae, and Geobacter sulfurreducens in laboratory experiments is a function of Fe(III) reduction rates and pathways by which biogenic minerals are formed. High Fe(III) reduction rates produced ⁵⁶Fe/⁵⁴Fe ratios for Fe(II)_aq that are 2-3‰ lower than the HFO substrate, reflecting a kinetic isotope fractionation that was associated with rapid sorption of Fe(II) to HFO. In long-term experiments at low Fe(III) reduction rates, the Fe(II)_aq-magnetite fractionation is −1.3‰, and this is interpreted to be the equilibrium fractionation factor at 22°C in the biologic reduction systems studied here. In experiments where Fe carbonate was the major ferrous product of HFO reduction, the estimated equilibrium Fe(II)_aq-Fe carbonate fractionations were ca. 0.0‰ for siderite (FeCO₃) and ca. +0.9‰ for Ca-substituted siderite (Ca_0.15Fe_0.85CO₃) at 22°C. Formation of precursor phases such as amorphous nonmagnetic, noncarbonate Fe(II) solids are important in the pathways to formation of biogenic magnetite or siderite, particularly at high Fe(III) reduction rates, and these solids may have ⁵⁶Fe/⁵⁴Fe ratios that are up to 1‰ lower than Fe(II)_aq. Under low Fe(III) reduction rates, where equilibrium is likely to be attained, it appears that both sorbed Fe(II) and amorphous Fe(II)(s) components have isotopic compositions that are similar to those of Fe(II)_aq.The relative order of δ⁵⁶Fe values for these biogenic minerals and aqueous Fe(II) is: magnetite > siderite ≈ Fe(II)_aq > Ca-bearing Fe carbonate, and this is similar to that observed for minerals from natural samples such as Banded Iron Formations (BIFs). Where magnetite from BIFs has δ⁵⁶Fe >0‰, the calculated δ⁵⁶Fe value for aqueous Fe(II) suggests a source from midocean ridge (MOR) hydrothermal fluids. In contrast, magnetite from BIFs that has δ⁵⁶Fe ≤0‰ apparently requires formation from aqueous Fe(II) that had very low δ⁵⁶Fe values. Based on this experimental study, formation of low-δ⁵⁶Fe Fe(II)_aq in nonsulfidic systems seems most likely to have been produced by dissimilatory reduction of ferric oxides by Fe(III)-reducing bacteria.     

Kinetic fractionation of Fe isotopes during transport through a porous quartz-sand column

Sorption and desorption processes are an important part of biological and geochemical metallic isotope cycles. Here, we address the dynamic aspects of metallic isotopic fractionation in a theoretical and experimental study of Fe sorption and desorption during the transport of aqueous Fe(III) through a quartz-sand matrix. Transport equations describing the behavior of sorbing isotopic species in a water saturated homogeneous porous medium are presented; isotopic fractionation of the system (Δ_{sorbedmetal-soln}) being defined in terms of two parameters: (i) an equilibrium fractionation factor, α_e; and (ii) a kinetic sorption factor, α₁. These equations are applied in a numerical model that simulates the sorption-desorption of Fe isotopes during injection of a Fe(III) solution pulse into a quartz matrix at pH 0-2 and explores the effects of the kinetic and equilibrium parameters on the Fe-isotope evolution of porewater. The kinetic transport theory is applied to a series of experiments in which pulses of Na and Fe(III) chloride solutions were injected into a porous sand grain column. Fractionation factors of α_e = 1.0003 ± 0.0001 and α₁ = 0.9997 ± 0.0004 yielded the best fit between the transport model and the Fe concentration and δ⁵⁶Fe data. The equilibrium fractionation (Δ⁵⁶Fe_{sorbedFe-soln}) of 0.3‰ is comparable with values deduced for adsorption of metallic cations on iron and manganese oxide surfaces and suggests that sandstone aquifers will fractionate metallic isotopes during sorption-desorption reactions. The ability of the equilibrium fractionation factor to describe a natural system, however, depends on the proximity to equilibrium, which is determined by the relative time scales of mass transfer and chemical reaction; low fluid transport rates should produce a system that is less dependent on kinetic effects. The results of this study are applicable to Fe-isotope fractionation in clastic sediments formed in highly acidic conditions; such conditions may have existed on Mars where acidic oxidizing ground and surface waters may have been responsible for clastic sedimentation and metallic element transport.     

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% H₂O) and natural pyrrhotites (Fe_1 − xS) 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 Fe²⁺/ΣFe ratio of our peralkaline rhyolitic melt of 0.38 ± 0.02. At higher Fe²⁺/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.     

Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions

Oxygen and iron isotope analyses of low-Ti and high-Ti mare basalts are presented to constrain their petrogenesis and to assess stable isotope variations within lunar mantle sources. An internally-consistent dataset of oxygen isotope compositions of mare basalts encompasses five types of low-Ti basalts from the Apollo 12 and 15 missions and eight types of high-Ti basalts from the Apollo 11 and 17 missions. High-precision whole-rock δ¹⁸O values (referenced to VSMOW) of low-Ti and high-Ti basalts correlate with major-element compositions (Mg#, TiO₂, Al₂O₃). The observed oxygen isotope variations within low-Ti and high-Ti basalts are consistent with crystal fractionation and match the results of mass-balance models assuming equilibrium crystallization. Whole-rock δ⁵⁶Fe values (referenced to IRMM-014) of high-Ti and low-Ti basalts range from 0.134‰ to 0.217‰ and 0.038‰ to 0.104‰, respectively. Iron isotope compositions of both low-Ti and high-Ti basalts do not correlate with indices of crystal fractionation, possibly owing to small mineral-melt iron fractionation factors anticipated under lunar reducing conditions.The δ¹⁸O and δ⁵⁶Fe values of low-Ti and the least differentiated high-Ti mare basalts are negatively correlated, which reflects their different mantle source characteristics (e.g., the presence or absence of ilmenite). The average δ⁵⁶Fe values of low-Ti basalts (0.073 ± 0.018‰, n = 8) and high-Ti basalts (0.191 ± 0.020‰, n = 7) may directly record that of their parent mantle sources. Oxygen isotope compositions of mantle sources of low-Ti and high-Ti basalts are calculated using existing models of lunar magma ocean crystallization and mixing, the estimated equilibrium mantle olivine δ¹⁸O value, and equilibrium oxygen-fractionation between olivine and other mineral phases. The differences between the calculated whole-rock δ¹⁸O values for source regions, 5.57‰ for low-Ti and 5.30‰ for high-Ti mare basalt mantle source regions, are solely a function of the assumed source mineralogy. The oxygen and iron isotope compositions of lunar upper mantle can be approximated using these mantle source values. The δ¹⁸O and δ⁵⁶Fe values of the lunar upper mantle are estimated to be 5.5 ± 0.2‰ (2σ) and 0.085 ± 0.040‰ (2σ), respectively. The oxygen isotope composition of lunar upper mantle is identical to the current estimate of Earth’s upper mantle (5.5 ± 0.2‰), and the iron isotope composition of the lunar upper mantle overlaps within uncertainty of estimates for the terrestrial upper mantle (0.044 ± 0.030‰).     

Calculation of equilibrium stable isotope partition function ratios for aqueous zinc complexes and metallic zinc

The goal of this study is to determine reduced partition function ratios for a variety of species of zinc, both as a metal and in aqueous solutions in order to calculate equilibrium stable isotope partitioning. We present calculations of the magnitude of Zn stable-isotope fractionation (^66,67,68Zn/⁶⁴Zn) between aqueous species and metallic zinc using measured vibrational spectra (fit from neutron scattering studies of metallic zinc) and a variety of electronic structure models. The results show that the reduced metal, Zn(0), will be light in equilibrium with oxidized Zn(II) aqueous species, with the best estimates for the Zn(II)-Zn(0) fractionation between hexaquo species and metallic zinc being Δ^66/64Zn_aq-metal ∼ 1.6‰ at 25 °C, and Δ^66/64Zn_aq-metal ∼ 0.8‰ between the tetrachloro zinc complex and metallic zinc at 25 °C using B3LYP/aug-cc-pVDZ level of theory and basis set. To examine the behavior of zinc in various aqueous solution chemistries, models for Zn(II) complex speciation were used to determine which species are thermodynamically favorable and abundant under a variety of different conditions relevant to natural waters, experimental and industrial solutions. The optimal molecular geometries for [Zn(H₂O)₆]²⁺, [Zn(H₂O)₆]·SO₄, [ZnCl₄]²⁻ and [Zn(H₂O)₃(C₃H₅O(COO)₃)]⁻ complexes in various states of solvation, protonation and coordination were calculated at various levels of electronic structure theory and basis set size. Isotopic reduced partition function ratios were calculated from frequency analyses of these optimized structures. Increasing the basis set size typically led to a decrease in the calculated reduced partition function ratios of ∼0.5‰ with values approaching a plateau using the aug-cc-pVDZ basis set or larger. The widest range of species were studied at the B3LYP/LAN2DZ/6-31G^∗ level of theory and basis-set size for comparison. Aqueous zinc complexes where oxygen is bound to the metal center tended to have the largest reduced partition function ratios, with estimated fractionations ranging from 2.2 to 2.9‰ (⁶⁶Zn/⁶⁴Zn) at 25 °C relative to metallic zinc. The tetrahedrally coordinated tetrachloro zinc complex, where zinc is bound exclusively to chloride, had the lowest reduced partition function ratio for a Zn(II) species (Δ^66/64Zn_aq-metal ∼ 1-1.3‰ at 25 °C). Increasing the number of waters in the second shell of solvation of the above complexes led to variable results, most commonly leading to a decrease of ∼0.2 to 0.3‰ in calculated Δ^66/64Zn_aq-metal at 25 °C.These estimates are useful in the interpretation of observed fractionations during the electrochemical deposition of zinc, where aqueous-metal fractionations of up to 5.5‰ are observed. The models show these are not caused by an equilibrium fractionation process. These results suggest that the redox cycle of zinc during industrial processing may be responsible for isotopically distinct reservoirs of zinc observed in polluted environments. The leaching of metallic zinc or zinc tailings from industrial sites could lead to the observed heavy signature in river systems, the magnitude of which will be reliant on the source material and the aqueous species that form.     

O and S isotopic composition of dissolved and attached oxidation products of pyrite by Acidithiobacillus ferrooxidans: Comparison with abiotic oxidations

The acidophilic iron-oxidizing bacterium, Acidithiobacillus ferrooxidans, plays a part in the pyrite oxidation process and has been widely studied in order to determine the kinetics of the reactions and the isotopic composition of dissolved product sulphates, but the details of the oxidation processes at the surface of pyrite are still poorly known. In this study, oxygen and sulphur isotopic compositions (δ¹⁸O and δ³⁴S) were analyzed for dissolved sulphates and water from experimental aerobic acidic (pH < 2) pyrite oxidation by A. ferrooxidans. The oxidation products attached to the pyrite surfaces were studied for their morphology (SEM), their chemistry (Raman spectroscopy) and for their δ¹⁸O (ion microprobe). They were compared to abiotically (Fe³⁺, H₂O₂, O₂) oxidized pyrite surface compounds in order to constrain the oxidation pathways and to look for the existence of potential biosignatures for this system.The pyrite dissolution evolved from non-stoichiometric (during the first days) to stoichiometric (with increasing time) resulting in dissolved sulphates having distinct δ¹⁸O (e.g. +11.0‰ and −2.0‰, respectively) and δ³⁴S (+4.5‰ and +2.8‰, respectively) values. The “oxidation layer” at the surface of pyrite is complex and made of iron oxides, sulphate, polysulphide, elemental sulphur and polythionates. Bio- and Fe³⁺-oxidation favour the development of monophased micrometric bumps made of hematite or sulphate while other abiotic oxidation processes result in more variable oxidation products. The δ¹⁸O of these oxidation products at the surface of oxidized pyrites are strongly variable (from ≈−40‰ to ≈+30‰) for all experiments.Isotopic fractionation between sulphates and pyrite, Δ³⁴S_SO4-pyrite, is equal to −1.3‰ and +0.4‰ for sulphates formed by stoichiometric and non-stoichiometric processes, respectively. These two values likely reflect either a S-S or a Fe-S bond breaking process. The Δ¹⁸O_SO4-H2O and Δ¹⁸O_SO4-O2 are estimated to be ≈+16‰ and ≈−25‰, respectively. These values are higher than previously published data and may reflect biological effects. The large δ¹⁸O heterogeneity measured at the surfaces of oxidized pyrites, whatever the oxidant, may be related (i) to the existence of local surface environments isolated from the solution in which the oxidation processes are different and (ii) to the stabilization at the pyrite surface of reaction intermediates that are not in isotopic equilibrium with the solution. Though the oxygen isotopic composition of surface oxidation products cannot be taken as a direct biosignature, the combined morphological, chemical and isotopic characterization of the surfaces of oxidized pyrites may furnish clues about a biological activity on a mineral surface.     

Isotopic fractionation of the major elements of molten basalt by chemical and thermal diffusion

Samples produced in piston cylinder experiments were used to document the thermal isotopic fractionation of all the major elements of basalt except for aluminum and the fractionation of iron isotopes by chemical diffusion between a natural basalt and rhyolite. The thermal isotopic fractionations are summarized in terms of a parameter Ω_i defined as the fractionation in per mil per 100 °C per atomic mass units difference between the isotopes. For molten basalt we report Ω_Ca = 1.6, Ω_Fe = 1.1, Ω_Si = 0.6, Ω_O = 1.5. In an earlier paper we reported Ω_Mg = 3.6. These fractionations represent a steady state balance between thermal diffusion and chemical diffusion with the mass dependence of the thermal diffusion coefficient being significantly larger than the mass dependence of the chemical diffusion coefficients for isotopes of the same element. The iron isotopic measurements of the basalt-rhyolite diffusion couple showed significant fractionation that are parameterized in terms of a parameter β_Fe = 0.03 when the ratio of the diffusion coefficients D₅₄ and D₅₆ of ⁵⁴Fe and ⁵⁶Fe is expressed in terms of the atomic mass as D₅₄/D₅₆ = (56/54)^β_Fe. This value of β_Fe is smaller than what we had measured earlier for lithium, magnesium and calcium (i.e., β_Li = 0.215, β_Ca = 0.05, β_Mg = 0.05) but still significant when one takes into account the high precision with which iron isotopic compositions can be measured (i.e., ±0.03‰) and that iron isotope fractionations at magmatic temperatures from other causes are extremely small. In a closing section we discuss technological and geological applications of isotopic fractionations driven by either or both chemical and thermal gradients.     

Oxygen isotope systematics of chondrules in the Allende CV3 chondrite: High precision ion microprobe studies

The oxygen three-isotope systematics of 36 chondrules from the Allende CV3 chondrite are reported using high precision secondary ion mass spectrometer (CAMECA IMS-1280). Twenty-six chondrules have shown internally homogenous Δ¹⁷O values among olivine, pyroxene, and spinel within a single chondrule. The average Δ¹⁷O values of 19 FeO-poor chondrules (13 porphyritic chondrules, 2 barred olivine chondrules, and 4 chondrule fragments) show a peak at −5.3 ± 0.6‰ (2SD). Another 5 porphyritic chondrules including both FeO-poor and FeO-rich ones show average Δ¹⁷O values between −3‰ and −2‰, and 2 other FeO-poor barred olivine chondrules show average Δ¹⁷O values of −3.6‰ and 0‰. These results are similar to those for Acfer 094 chondrules, showing bimodal Δ¹⁷O values at −5‰ and −2‰. Nine porphyritic chondrules contain olivine grains with heterogeneous Δ¹⁷O values as low as −18‰, indicating that they are relict olivine grains and some of them were derived from precursors related to refractory inclusions. However, most relict olivine grains show oxygen isotope ratios that overlap with those in homogeneous chondrules. The Δ¹⁷O values of four barred olivine chondrules range from −5‰ to 0‰, indicating that not all BO chondrules plot near the terrestrial fractionation line as suggested by previous bulk chondrule analyses. Based on these data, we suggest the presence of multiple oxygen isotope reservoirs in local dust-rich protoplanetary disk, from which the CV3 parent asteroid formed.A compilation of 225 olivine and low-Ca pyroxene isotopic data from 36 chondrules analyzed in the present study lie between carbonaceous chondrite anhydrous mineral (CCAM) and Young and Russell lines. These data define a correlation line of δ¹⁷O = (0.982 ± 0.019) × δ¹⁸O − (2.91 ± 0.10), which is similar to those defined by chondrules in CV3 chondrites and Acfer 094 in previous studies. Plagioclase analyses in two chondrules plot slightly below the CCAM line with Δ¹⁷O values of −2.6‰, which might be the result of oxygen isotope exchange between chondrule mesostasis and aqueous fluid in the CV parent body.     

High precision SIMS oxygen three isotope study of chondrules in LL3 chondrites: Role of ambient gas during chondrule formation

We report high precision SIMS oxygen three isotope analyses of 36 chondrules from some of the least equilibrated LL3 chondrites, and find systematic variations in oxygen isotope ratios with chondrule types. FeO-poor (type I) chondrules generally plot along a mass dependent fractionation line (Δ¹⁷O ∼ 0.7‰), with δ¹⁸O values lower in olivine-rich (IA) than pyroxene-rich (IB) chondrules. Data from FeO-rich (type II) chondrules show a limited range of δ¹⁸O and δ¹⁷O values at δ¹⁸O = 4.5‰, δ¹⁷O = 2.9‰, and Δ¹⁷O = 0.5‰, which is slightly ¹⁶O-enriched relative to bulk LL chondrites (Δ¹⁷O ∼ 1.3‰). Data from four chondrules show ¹⁶O-rich oxygen isotope ratios that plot near the CCAM (Carbonaceous Chondrite Anhydrous Mineral) line. Glass analyses in selected chondrules are systematically higher than co-existing minerals in both δ¹⁸O and Δ¹⁷O values, whereas high-Ca pyroxene data in the same chondrule are similar to those in olivine and pyroxene phenocrysts.Our results suggest that the LL chondrite chondrule-forming region contained two kinds of solid precursors, (1) ¹⁶O-poor precursors with Δ¹⁷O > 1.6‰ and (2) ¹⁶O-rich solid precursors derived from the same oxygen isotope reservoir as carbonaceous chondrites. Oxygen isotopes exhibited open system behavior during chondrule formation, and the interaction between the solid and ambient gas might occur as described in the following model. Significant evaporation and recondensation of solid precursors caused a large mass-dependent fractionation due to either kinetic or equilibrium isotope exchange between gas and solid to form type IA chondrules with higher bulk Mg/Si ratios. Type II chondrules formed under elevated dust/gas ratios and with water ice in the precursors, in which the ambient H₂O gas homogenized chondrule melts by isotope exchange. Low temperature oxygen isotope exchange may have occurred between chondrule glasses and aqueous fluids with high Δ¹⁷O (∼5‰) in LL the parent body. According to our model, oxygen isotope ratios of chondrules were strongly influenced by the local solid precursors in the proto-planetary disk and the ambient gas during chondrule melting events.     

Calibration of the calcite-water oxygen-isotope geothermometer at Devils Hole, Nevada, a natural laboratory

The δ¹⁸O of ground water (−13.54 ± 0.05 ‰) and inorganically precipitated Holocene vein calcite (+14.56 ± 0.03 ‰) from Devils Hole cave #2 in southcentral Nevada yield an oxygen isotopic fractionation factor between calcite and water at 33.7 °C of 1.02849 ± 0.00013 (1000 ln α_{calcite-water} = 28.09 ± 0.13). Using the commonly accepted value of ∂(α_{calcite-water})/∂T of −0.00020 K⁻¹, this corresponds to a 1000 ln α_{calcite-water} value at 25 °C of 29.80, which differs substantially from the current accepted value of 28.3. Use of previously published oxygen isotopic fractionation factors would yield a calcite precipitation temperature in Devils Hole that is 8 °C lower than the measured ground water temperature. Alternatively, previously published fractionation factors would yield a δ¹⁸O of water, from which the calcite precipitated, that is too negative by 1.5 ‰ using a temperature of 33.7 °C. Several lines of evidence indicate that the geochemical environment of Devils Hole has been remarkably constant for at least 10 ka. Accordingly, a re-evaluation of calcite-water oxygen isotopic fractionation factor may be in order.Assuming the Devils Hole oxygen isotopic value of α_{calcite-water} represents thermodynamic equilibrium, many marine carbonates are precipitated with a δ¹⁸O value that is too low, apparently due to a kinetic isotopic fractionation that preferentially enriches ¹⁶O in the solid carbonate over ¹⁸O, feigning oxygen isotopic equilibrium.     

Density functional theory predictions of equilibrium isotope fractionation of iron due to redox changes and organic complexation

We present molecular orbital/density functional theory (MO/DFT) calculations that predict a greater isotopic fractionation in redox reactions than in reactions involving ligand exchange. The predicted fractionation factors, reported as 1000·ln(^56-54α), associated with equilibrium between Fe-organic and Fe-H₂O species were <1.6‰ in vacuo and <1.2‰ in solution when the oxidation state of the system was held constant. These fractionation factors were significantly smaller than those predicted for equilibrium between different oxidation states of Fe, for which 1000·ln(^56-54α) was >2.7‰ in vacuo and >2.2‰ in solution when the bound ligands were unchanged. The predicted ⁵⁶Fe/⁵⁴Fe ratio was greater in complexes containing Fe³⁺ and in complexes with shorter Fe-O bond lengths; both of these trends follow previous theoretical results. Our predictions also agree with previous experimental measurements that suggest that the largest biological fractionations will be associated with processes that change the oxidation state of Fe, and that identification of biologically controlled Fe isotope fractionation may be difficult when abiotic redox fractionations are present in the system. The models studied here also have important implications for future theoretical isotope calculations, because we have discovered the necessity of using vibrational frequencies instead of reduced masses when predicting reduced partition functions in aqueous-phase species.