Iron isotope fractionation by Fe(II)-oxidizing photoautotrophic bacteria

Photoautotrophic bacteria that oxidize ferrous iron (Fe[II]) under anaerobic conditions are thought to be ancient in origin, and the ferric (hydr)oxide mineral products of their metabolism are likely to be preserved in ancient rocks. Here, two enrichment cultures of Fe(II)-oxidizing photoautotrophs and a culture of the genus Thiodictyon were studied with respect to their ability to fractionate Fe isotopes. Fe isotope fractionations produced by both the enrichment cultures and the Thiodictyon culture were relatively constant at early stages of the reaction progress, where the ⁵⁶Fe/⁵⁴Fe ratios of poorly crystalline hydrous ferric oxide (HFO) metabolic products were enriched in the heavier isotope relative to aqueous ferrous iron (Fe[II]_aq) by ∼1.5 ± 0.2‰. This fractionation appears to be independent of the rate of photoautotrophic Fe(II)-oxidation, and is comparable to that observed for Fe isotope fractionation by dissimilatory Fe(III)-reducing bacteria. Although there remain a number of uncertainties regarding how the overall measured isotopic fractionation is produced, the most likely mechanisms include (1) an equilibrium effect produced by biological ligands, or (2) a kinetic effect produced by precipitation of HFO overlaid upon equilibrium exchange between Fe(II) and Fe(III) species. The fractionation we observe is similar in direction to that measured for abiotic oxidation of Fe(II)_aq by molecular oxygen. This suggests that the use of Fe isotopes to identify phototrophic Fe(II)-oxidation in the rock record may only be possible during time periods in Earth’s history when independent evidence exists for low ambient oxygen contents.     

Evidence for equilibrium iron isotope fractionation by nitrate-reducing iron(II)-oxidizing bacteria

Iron isotope fractionations produced during chemical and biological Fe(II) oxidation are sensitive to the proportions and nature of dissolved and solid-phase Fe species present, as well as the extent of isotopic exchange between precipitates and aqueous Fe. Iron isotopes therefore potentially constrain the mechanisms and pathways of Fe redox transformations in modern and ancient environments. In the present study, we followed in batch experiments Fe isotope fractionations between Fe(II)_aq and Fe(III) oxide/hydroxide precipitates produced by the Fe(III) mineral encrusting, nitrate-reducing, Fe(II)-oxidizing Acidovorax sp. strain BoFeN1. Isotopic fractionation in ⁵⁶Fe/⁵⁴Fe approached that expected for equilibrium conditions, assuming an equilibrium Δ⁵⁶Fe_{Fe(OH)3-Fe(II)aq} fractionation factor of +3.0‰. Previous studies have shown that Fe(II) oxidation by this Acidovorax strain occurs in the periplasm, and we propose that Fe isotope equilibrium is maintained through redox cycling via coupled electron and atom exchange between Fe(II)_aq and Fe(III) precipitates in the contained environment of the periplasm. In addition to the apparent equilibrium isotopic fractionation, these experiments also record the kinetic effects of initial rapid oxidation, and possible phase transformations of the Fe(III) precipitates. Attainment of Fe isotope equilibrium between Fe(III) oxide/hydroxide precipitates and Fe(II)_aq by neutrophilic, Fe(II)-oxidizing bacteria or through abiologic Fe(II)_aq oxidation is generally not expected or observed, because the poor solubility of their metabolic product, i.e. Fe(III), usually leads to rapid precipitation of Fe(III) minerals, and hence expression of a kinetic fractionation upon precipitation; in the absence of redox cycling between Fe(II)_aq and precipitate, kinetic isotope fractionations are likely to be retained. These results highlight the distinct Fe isotope fractionations that are produced by different pathways of biological and abiological Fe(II) oxidation.     

Stable Fe isotope fractionations produced by aqueous Fe(II)-hematite surface interactions

Stable Fe isotope fractionations were investigated during exposure of hematite to aqueous Fe(II) under conditions of variable Fe(II)/hematite ratios, the presence/absence of dissolved Si, and neutral versus alkaline pH. When Fe(II) undergoes electron transfer to hematite, Fe(II) is initially oxidized to Fe(III), and structural Fe(III) on the hematite surface is reduced to Fe(II). During this redox reaction, the newly formed reactive Fe(III) layer becomes enriched in heavy Fe isotopes and light Fe isotopes partition into aqueous and sorbed Fe(II). Our results indicate that in most cases the reactive Fe(III) that undergoes isotopic exchange accounts for less than one octahedral layer on the hematite surface. With higher Fe(II)/hematite molar ratios, and the presence of dissolved Si at alkaline pH, stable Fe isotope fractionations move away from those expected for equilibrium between aqueous Fe(II) and hematite, towards those expected for aqueous Fe(II) and goethite. These results point to formation of new phases on the hematite surface as a result of distortion of Fe-O bonds and Si polymerization at high pH. Our findings demonstrate how stable Fe isotope fractionations can be used to investigate changes in surface Fe phases during exposure of Fe(III) oxides to aqueous Fe(II) under different environmental conditions. These results confirm the coupled electron and atom exchange mechanism proposed to explain Fe isotope fractionation during dissimilatory iron reduction (DIR). Although abiologic Fe(II)_aq - oxide interaction will produce low δ⁵⁶Fe values for Fe(II)_aq, similar to that produced by Fe(II) oxidation, only small quantities of low-δ⁵⁶Fe Fe(II)_aq are formed by these processes. In contrast, DIR, which continually exposes new surface Fe(III) atoms during reduction, as well as production of Fe(II), remains the most efficient mechanism for generating large quantities of low-δ⁵⁶Fe aqueous Fe(II) in many natural systems.     

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 and equilibrium Fe isotope fractionation between aqueous Fe(II) and Fe(III)

Equilibrium and kinetic Fe isotope fractionation between aqueous ferrous and ferric species measured over a range of chloride concentrations (0, 11, 110 mM Cl⁻) and at two temperatures (0 and 22°C) indicate that Fe isotope fractionation is a function of temperature, but independent of chloride contents over the range studied. Using ⁵⁷Fe-enriched tracer experiments the kinetics of isotopic exchange can be fit by a second-order rate equation, or a first-order equation with respect to both ferrous and ferric iron. The exchange is rapid at 22°C, ∼60-80% complete within 5 seconds, whereas at 0°C, exchange rates are about an order of magnitude slower. Isotopic exchange rates vary with chloride contents, where ferrous-ferric isotope exchange rates were ∼25 to 40% slower in the 11 mM HCl solution compared to the 0 mM Cl⁻ (∼10 mM HNO₃) solutions; isotope exchange rates are comparable in the 0 and 110 mM Cl⁻ solutions.The average measured equilibrium isotope fractionations, Δ_{Fe(III)-Fe(II)}, in 0, 11, and 111 mM Cl⁻ solutions at 22°C are identical within experimental error at +2.76±0.09, +2.87±0.22, and +2.76±0.06 ‰, respectively. This is very similar to the value measured by Johnson et al. (2002a) in dilute HCl solutions. At 0°C, the average measured Δ_{Fe(III)-Fe(II)} fractionations are +3.25±0.38, +3.51±0.14 and +3.56±0.16 ‰ for 0, 11, and 111 mM Cl⁻ solutions. Assessment of the effects of partial re-equilibration on isotope fractionation during species separation suggests that the measured isotope fractionations are on average too low by ∼0.20 ‰ and ∼0.13 ‰ for the 22°C and 0°C experiments, respectively. Using corrected fractionation factors, we can define the temperature dependence of the isotope fractionation from 0°C to 22°C as: where the isotopic fractionation is independent of Cl⁻ contents over the range used in these experiments. These results confirm that the Fe(III)-Fe(II) fractionation is approximately half that predicted from spectroscopic data, and suggests that, at least in moderate Cl⁻ contents, the isotopic fractionation is relatively insensitive to Fe-Cl speciation.     

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.     

Ab initio calculations of the Fe(II) and Fe(III) isotopic effects in citrates,nicotianamine, and phytosiderophore,and new Fe isotopic measurements in higher plants

Iron is one of the most abundant transition metal in higher plants and variations in its isotopic compositions can be used to trace its utilization. In order to better understand the effect of plant-induced isotopic fractionation on the global Fe cycling, we have estimated by quantum chemical calculations the magnitude of the isotopic fractionation between different Fe species relevant to the transport and storage of Fe in higher plants: Fe(II)-citrate, Fe(III)-citrate, Fe(II)-nicotianamine, and Fe(III)-phytosiderophore. The ab initio calculations show firstly, that Fe(II)-nicotianamine is ～3‰ (⁵⁶Fe/⁵⁴Fe) isotopically lighter than Fe(III)-phytosiderophore; secondly, even in the absence of redox changes of Fe, change in the speciation alone can create up to ～1.5‰ isotopic fractionation. For example, Fe(III)-phytosiderophore is up to 1.5‰ heavier than Fe(III)-citrate₂ and Fe(II)-nicotianamine is up to 1‰ heavier than Fe(II)-citrate. In addition, in order to better understand the Fe isotopic fractionation between different plant components, we have analyzed the iron isotopic composition of different organs (roots, seeds, germinated seeds, leaves and stems) from six species of higher plants: the dicot lentil (Lens culinaris), and the graminaceous monocots Virginia wild rye (Elymus virginicus), Johnsongrass (Sorghum halepense), Kentucky bluegrass (Poa pratensis), river oat (Uniola latifolia), and Indian goosegrass (Eleusine indica). The calculations may explain that the roots of strategy-II plants (Fe(III)-phytosiderophore) are isotopically heavier (by about 1‰ for the δ⁵⁶Fe) than the upper parts of the plants (Fe transported as Fe(III)-citrate in the xylem or Fe(II)-nicotianamine in the phloem). In addition, we suggest that the isotopic variations observed between younger and older leaves could be explained by mixing of Fe received from the xylem and the phloem.     

Iron isotope fractionation and atom exchange during sorption of ferrous iron to mineral surfaces

The application of stable Fe isotopes as a tracer of the biogeochemical Fe cycle necessitates a mechanistic knowledge of natural fractionation processes. We studied the equilibrium Fe isotope fractionation upon sorption of Fe(II) to aluminum oxide (γ-Al₂O₃), goethite (α-FeOOH), quartz (α-SiO₂), and goethite-loaded quartz in batch experiments, and performed continuous-flow column experiments to study the extent of equilibrium and kinetic Fe isotope fractionation during reactive transport of Fe(II) through pure and goethite-loaded quartz sand. In addition, batch and column experiments were used to quantify the coupled electron transfer-atom exchange between dissolved Fe(II) (Fe(II)_aq) and structural Fe(III) of goethite. All experiments were conducted under strictly anoxic conditions at pH 7.2 in 20 mM MOPS (3-(N-morpholino)-propanesulfonic acid) buffer and 23 °C. Iron isotope ratios were measured by high-resolution MC-ICP-MS. Isotope data were analyzed with isotope fractionation models. In batch systems, we observed significant Fe isotope fractionation upon equilibrium sorption of Fe(II) to all sorbents tested, except for aluminum oxide. The equilibrium enrichment factor, , of the Fe(II)_sorb-Fe(II)_aq couple was 0.85 ± 0.10‰ (±2σ) for quartz and 0.85 ± 0.08‰ (±2σ) for goethite-loaded quartz. In the goethite system, the sorption-induced isotope fractionation was superimposed by atom exchange, leading to a δ^56/54Fe shift in solution towards the isotopic composition of the goethite. Without consideration of atom exchange, the equilibrium enrichment factor was 2.01 ± 0.08‰ (±2σ), but decreased to 0.73 ± 0.24‰ (±2σ) when atom exchange was taken into account. The amount of structural Fe in goethite that equilibrated isotopically with Fe(II)_aq via atom exchange was equivalent to one atomic Fe layer of the mineral surface (∼3% of goethite-Fe). Column experiments showed significant Fe isotope fractionation with δ^56/54Fe(II)_aq spanning a range of 1.00‰ and 1.65‰ for pure and goethite-loaded quartz, respectively. Reactive transport of Fe(II) under non-steady state conditions led to complex, non-monotonous Fe isotope trends that could be explained by a combination of kinetic and equilibrium isotope enrichment factors. Our results demonstrate that in abiotic anoxic systems with near-neutral pH, sorption of Fe(II) to mineral surfaces, even to supposedly non-reactive minerals such as quartz, induces significant Fe isotope fractionation. Therefore we expect Fe isotope signatures in natural systems with changing concentration gradients of Fe(II)_aq to be affected by sorption.     

Iron isotope fractionation between pyrite (FeS2), hematite (Fe2O3) and siderite (FeCO3): 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 α⁵⁷Fe/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. Acta71, 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.     

Influence of pH and dissolved Si on Fe isotope fractionation during dissimilatory microbial reduction of hematite

Microbial dissimilatory iron reduction (DIR) has been identified as a mechanism for production of aqueous Fe(II) that has low ⁵⁶Fe/⁵⁴Fe ratios in modern and ancient suboxic environments that contain ferric oxides or hydroxides. These studies suggest that DIR could have played an important role in producing distinct Fe isotope compositions in Precambrian banded iron formations or other marine sedimentary rocks. However, the applicability of experimental studies of Fe isotope fractionation produced by DIR in geochemically simple systems to ancient marine environments remains unclear. Here we report Fe isotope fractionations produced during dissimilatory microbial reduction of hematite by Geobacter sulfurreducens in the presence and absence of dissolved Si at neutral and alkaline pH. Hematite reduction was significantly decreased by Si at alkaline (but not neutral) pH, presumably due to Si polymerization at the hematite surface. The presence of Si altered Fe isotope fractionation factors between aqueous Fe(II) or sorbed Fe(II) and reactive Fe(III), reflecting changes in bonding environment of the reactive Fe(III) component at the oxide surface. Despite these changes in isotopic fractionations, our results demonstrate that microbial Fe(III) oxide reduction produces Fe(II) with negative δ⁵⁶Fe values under conditions of variable pH and dissolved Si, similar to the large inventory of negative δ⁵⁶Fe in Neoarchean and Paleoproterozoic age marine sedimentary rocks.     

Structural constraints of ferric (hydr)oxides on dissimilatory iron reduction and the fate of Fe(II)

Due to the strong reducing capacity of ferrous Fe, the fate of Fe(II) following dissimilatory iron reduction will have a profound bearing on biogeochemical cycles. We have previously observed the rapid and near complete conversion of 2-line ferrihydrite to goethite (minor phase) and magnetite (major phase) under advective flow in an organic carbon-rich artificial groundwater medium. Yet, in many mineralogically mature environments, well-ordered iron (hydr)oxide phases dominate and may therefore control the extent and rate of Fe(III) reduction. Accordingly, here we compare the reducing capacity and Fe(II) sequestration mechanisms of goethite and hematite to 2-line ferrihydrite under advective flow within a medium mimicking that of natural groundwater supplemented with organic carbon. Introduction of dissolved organic carbon upon flow initiation results in the onset of dissimilatory iron reduction of all three Fe phases (2-line ferrihydrite, goethite, and hematite). While the initial surface area normalized rates are similar (∼10⁻¹¹ mol Fe(II) m⁻² g⁻¹), the total amount of Fe(III) reduced over time along with the mechanisms and extent of Fe(II) sequestration differ among the three iron (hydr)oxide substrates. Following 16 d of reaction, the amount of Fe(III) reduced within the ferrihydrite, goethite, and hematite columns is 25, 5, and 1%, respectively. While 83% of the Fe(II) produced in the ferrihydrite system is retained within the solid-phase, merely 17% is retained within both the goethite and hematite columns. Magnetite precipitation is responsible for the majority of Fe(II) sequestration within ferrihydrite, yet magnetite was not detected in either the goethite or hematite systems. Instead, Fe(II) may be sequestered as localized spinel-like (magnetite) domains within surface hydrated layers (ca. 1 nm thick) on goethite and hematite or by electron delocalization within the bulk phase. The decreased solubility of goethite and hematite relative to ferrihydrite, resulting in lower Fe(III)_aq and bacterially-generated Fe(II)_aq concentrations, may hinder magnetite precipitation beyond mere surface reorganization into nanometer-sized, spinel-like domains. Nevertheless, following an initial, more rapid reduction period, the three Fe (hydr)oxides support similar aqueous ferrous iron concentrations, bacterial populations, and microbial Fe(III) reduction rates. A decline in microbial reduction rates and further Fe(II) retention in the solid-phase correlates with the initial degree of phase disorder (high energy sites). As such, sustained microbial reduction of 2-line ferrihydrite, goethite, and hematite appears to be controlled, in large part, by changes in surface reactivity (energy), which is influenced by microbial reduction and secondary Fe(II) sequestration processes regardless of structural order (crystallinity) and surface area.     

Theoretical investigation of iron isotope fractionation between Fe(H2O)6 and Fe(H2O)6: Implications for iron stable isotope geochemistry

The magnitude of equilibrium iron isotope fractionation between Fe(H₂O)₆³⁺ and Fe(H₂O)₆²⁺ 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 (10³ · ln(β)), and hence, the equilibrium isotope fractionation factor (10³ · 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, 10³ · 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 ν₄ mode (O-Fe-O deformation) of Fe(H₂O)₆³⁺ 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 10³ · 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(H₂O)₆³⁺ − Fe(H₂O)₆²⁺ fractionation factor and demonstrate the utility of DFT for applications in “heavy” stable isotope geochemistry.     

The effect of early diagenesis on the Fe isotope compositions of porewaters and authigenic minerals in continental margin sediments

Iron isotope compositions in marine pore fluids and sedimentary solid phases were measured at two sites along the California continental margin, where isotope compositions range from δ⁵⁶Fe = −3.0‰ to +0.4‰. At one site near Monterey Canyon off central California, organic matter oxidation likely proceeds through a number of diagenetic pathways that include significant dissimilatory iron reduction (DIR) and bacterial sulfate reduction, whereas at our other site in the Santa Barbara basin DIR appears to be comparatively small, and production of sulfides (FeS and pyrite) was extensive. The largest range in Fe isotope compositions is observed for Fe(II)_aq in porewaters, which generally have the lowest δ⁵⁶Fe values (minimum: −3.0‰) near the sediment surface, and increase with burial depth. δ⁵⁶Fe values for FeS inferred from HCl extractions vary between ∼−0.4‰ and +0.4‰, but pyrite is similar at both stations, where an average δ⁵⁶Fe value of −0.8 ± 0.2‰ was measured. We interpret variations in dissolved Fe isotope compositions to be best explained by open-system behavior that involves extensive recycling of Feflux. This study is the first to examine Fe isotope variations in modern marine sediments, and the results show that Fe isotopes in the various reactive Fe pools undergo isotopic fractionation during early diagenesis. Importantly, processes dominated by sulfide formation produce high-δ⁵⁶Fe values for porewaters, whereas the opposite occurs when Fe(III)-oxides are present and DIR is a major pathway of organic carbon respiration. Because shelf pore fluids may carry a negative δ⁵⁶Fe signature it is possible that the Fe isotope composition of ocean water reflects a significant contribution of shelf-derived iron to the open ocean. Such a signature would be an important means for tracing iron sources to the ocean and water mass circulation.     

Fe isotopic fractionation during mineral dissolution with and without bacteria

Fe released into solution is isotopically lighter (enriched in the lighter isotope) than hornblende starting material when dissolution occurs in the presence of the siderophore desferrioxamine mesylate (DFAM). In contrast, Fe released from goethite dissolving in the presence of DFAM is isotopically unchanged. Furthermore, Δ⁵⁶Fe_{solution-hornblende} for Fe released to solution in the presence of ligands varies with the affinity of the ligand for Fe. The extent of isotopic fractionation of Fe released from hornblende also increases when experiments are agitated continuously. The Fe isotope fractionation observed during hornblende dissolution with organic ligands is attributed predominantly to retention of ⁵⁶Fe in an altered surface layer, while the lack of isotopic fractionation during goethite dissolution in DFAM is consistent with the lack of an altered layer. When a siderophore-producing soil bacterium is added to the system (without added organic ligands), Fe released to solution from both hornblende and goethite differs isotopically from Fe in the bulk mineral: Δ⁵⁶Fe_{solution-starting material} = −0.56 ± 0.19 (hornblende) and −1.44 ± 0.16 (goethite). Increased isotopic fractionation is attributed in this case to the fact that as bacterial respiration depletes the system in oxygen and aqueous Fe is reduced, equilibration between aqueous ferrous and ferric iron creates a pool of isotopically heavy ferric iron that is assimilated by bacterial cells. Adsorption of isotopically heavy ferrous iron (Fe(II) enriched in the heavier isotope) or precipitation of isotopically heavy Fe minerals may also contribute to observed fractionations.To test whether these Fe isotope signatures are recorded in natural systems, we also investigated extractions of samples of soils from which the bacteria were isolated. These extractions show variability in the isotopic signatures of exchangeable Fe and Fe oxyhydroxide fractions from one soil sample to another, but exchangeable Fe is observed to be lighter than Fe in soil Fe oxyhydroxides and hornblende. This observation is consistent with isotopically light Fe-organic complexes in soil pore water derived from the Fe-silicate starting materials in the presence of growing microorganisms, as documented in experiments reported here. The contributions from phenomena including organic ligand-promoted nonstoichiometric dissolution of Fe silicates, uptake of ferric iron by organisms, adsorption of isotopically heavy ferrous iron, and precipitation of iron minerals should create complex isotopic signatures in soils. Better understanding of these processes and the timescales over which they contribute to fractionation is needed.     

Selenium isotope fractionation during reduction by Fe(II)-Fe(III) hydroxide-sulfate (green rust)

We have determined the extent of Se isotope fractionation induced by reduction of selenate by sulfate interlayered green rust (GR_SO4), a Fe(II)-Fe(III) hydroxide-sulfate. This compound is known to reduce selenate to Se(0), and it is the only naturally relevant abiotic selenate reduction pathway documented to date. Se reduction reactions, when they occur in nature, greatly reduce Se mobility and bioavailability. Se stable isotope analysis shows promise as an indicator of Se reduction, and Se isotope fractionation by various Se reactions must be known in order to refine this tool. We measured the increase in the ⁸⁰Se/⁷⁶Se ratio of dissolved selenate as lighter isotopes were preferentially consumed during reduction by GR_SO4. Six different experiments that used GR_SO4 made by two methods, with varying solution compositions and pH, yielded identical isotopic fractionations. Regression of all the data yielded an instantaneous isotope fractionation of 7.36 ± 0.24‰. Selenate reduction by GR_SO4 induces much greater isotopic fractionation than does bacterial selenate reduction. If selenate reduction by GR_SO4 occurs in nature, it may be identifiable on the basis of its relatively large isotopic fractionation.     

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.     

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.     

Adsorption as a cause for iron isotope fractionation in reduced groundwater

Iron isotopes were used to investigate iron transformation processes during an in situ field experiment for removal of dissolved Fe from reduced groundwater. This experiment provided a unique setting for exploring Fe isotope fractionation in a natural system. Oxygen-containing water was injected at a test well into an aquifer containing Fe(II)-rich reduced water, leading to oxidation of Fe(II) and precipitation of Fe(III)(hydr)oxides. Subsequently, groundwater was extracted from the same well over a time period much longer than the injection time. Since the surrounding water is rich in Fe(II), the Fe(II) concentration in the extracted water increased over time. The increase was strongly retarded in comparison to a conservative tracer added to the injected solution, indicating that adsorption of Fe(II) onto the newly formed Fe(III)(hydr)oxides occurred. A series of injection-extraction (push-pull) cycles were performed at the same well. The δ⁵⁷Fe/⁵⁴Fe of pre-experiment background groundwater (−0.57 ± 0.17 ‰) was lighter than the sediment leach of Fe(III) (−0.24 ± 0.08 ‰), probably due to slight fractionation (only ∼0.3 ‰) during microbial mediated reductive dissolution of Fe(III)(hydr)oxides present in the aquifer. During the experiment, Fe(II) was adsorbed from native groundwater drawn into the oxidized zone and onto Fe(III)(hydr)oxides producing a very light groundwater component with δ⁵⁷Fe/⁵⁴Fe as low as −4 ‰, indicating that heavier Fe(II) is preferentially adsorbed to the newly formed Fe(III)(hydr)oxides surfaces. Iron concentrations increased with time of extraction, and δ⁵⁷Fe/⁵⁴Fe linearly correlated with Fe concentrations (R² = 0.95). This pattern was reproducible over five individual cycles, indicating that the same process occurs during repeated injection/extraction cycles. We present a reactive transport model to explain the observed abiotic fractionation due to adsorption of Fe(II) on Fe(III)(hydr)oxides. The fractionation is probably caused by isotopic differences in the equilibrium sorption constants of the various isotopes (K_ads) and not by sorption kinetics. A fractionation factor α_57/54 of 1.001 fits the observed fractionation.     

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.     

Oxygen and sulfur isotope systematics of sulfate produced by bacterial and abiotic oxidation of pyrite

To better understand reaction pathways of pyrite oxidation and biogeochemical controls on δ¹⁸O and δ³⁴S values of the generated sulfate in acid mine drainage (AMD) and other natural environments, we conducted a series of pyrite oxidation experiments in the laboratory. Our biological and abiotic experiments were conducted under aerobic conditions by using O₂ as an oxidizing agent and under anaerobic conditions by using dissolved Fe(III)_aq as an oxidant with varying δ¹⁸O_H2O values in the presence and absence of Acidithiobacillus ferrooxidans. In addition, aerobic biological experiments were designed as short- and long-term experiments where the final pH was controlled at ∼2.7 and 2.2, respectively. Due to the slower kinetics of abiotic sulfide oxidation, the aerobic abiotic experiments were only conducted as long term with a final pH of ∼2.7. The δ³⁴S_SO4 values from both the biological and abiotic anaerobic experiments indicated a small but significant sulfur isotope fractionation (∼−0.7‰) in contrast to no significant fractionation observed from any of the aerobic experiments. Relative percentages of the incorporation of water-derived oxygen and dissolved oxygen (O₂) to sulfate were estimated, in addition to the oxygen isotope fractionation between sulfate and water, and dissolved oxygen. As expected, during the biological and abiotic anaerobic experiments all of the sulfate oxygen was derived from water. The percentage incorporation of water-derived oxygen into sulfate during the oxidation experiments by O₂ varied with longer incubation and lower pH, but not due to the presence or absence of bacteria. These percentages were estimated as 85%, 92% and 87% from the short-term biological, long-term biological and abiotic control experiments, respectively. An oxygen isotope fractionation effect between sulfate and water (ε¹⁸O_SO4-H2O) of ∼3.5‰ was determined for the anaerobic (biological and abiotic) experiments. This measured value was then used to estimate the oxygen isotope fractionation effects between sulfate and dissolved oxygen in the aerobic experiments which were −10.0‰, −10.8‰, and −9.8‰ for the short-term biological, long-term biological and abiotic control experiments, respectively. Based on the similarity between δ¹⁸O_SO4 values in the biological and abiotic experiments, it is suggested that δ¹⁸O_SO4 values cannot be used to distinguish biological and abiotic mechanisms of pyrite oxidation. The results presented here suggest that Fe(III)_aq is the primary oxidant for pyrite at pH < 3, even in the presence of dissolved oxygen, and that the main oxygen source of sulfate is water-oxygen under both aerobic and anaerobic conditions.