Fe Isotope Fractionation Factors Research Articles

Iron solid-phase differentiation controls isotopic fractionation during lateritic weathering of basalt

Laterites are important iron (Fe) reservoirs impacting the biogeochemical cycle of Fe at Earth’s surface, and Fe oxides are critical Fe-host minerals in laterites that affect the geochemical behavior of Fe. Whether and how the differentiation of Fe minerals controls Fe isotope fractionation during laterization, however, remains unclear. In this study, the mineralogy, element concentrations, and Fe isotope compositions are reported for a basalt-derived lateritic profile on Hainan Island, China, to investigate the role of Fe solid-phase differentiation in controlling Fe isotope fractionation during lateritic weathering. 57Fe Mössbauer spectroscopy at 13 K is conducted to quantitatively identify the solid-phase Fe, including FeIII-(oxy)hydroxides, FeIII-oxides, and organic/silicate-bound FeIII. The results from soil samples in the lower section show that RMZr,Fe values are positive and that δ56Fe values (−0.01‰ to 0.07‰; 0.05 ± 0.05‰ on average) are identical to those in parent basalt (0.03 ± 0.04‰). The limited variation in δ56Fe values is likely caused by the compaction effect and/or the downward transport of Fe from the upper soils. In contrast, RMZr,Fe values in soils from the upper section are generally negative, and their δ56Fe values (0.01‰ to 0.29‰) display zigzag variations and are linearly correlated with the amount of Fe present in hematite and goethite. The variation in δ56Fe value is likely caused by the atomic distribution of Fe in hematite and goethite, with heavy Fe preferentially incorporated into hematite rather than goethite. Based on the linear regression analysis, we calculated the apparent Fe isotope fractionation factor between hematite and goethite as 0.99 ± 0.18‰ (SD). A comparison of published Fe isotope data shows that the variation in δ56Fe value of silicate-derived soils and rivers may have a potential link with climatic factor (mostly temperature). Our study highlights the importance of Fe-mineral differentiation in controlling Fe isotope fractionation in laterite systems.

Consecutive Fe redox cycles decrease bioreducible Fe(III) and Fe isotope fractionations by eliminating small clay particles

Previous work has shown that about 10% of total clay-bound Fe(III) in unaltered nontronite NAu-1 is bioreducible, although it remains unclear how much of the bioreducible Fe pool persists after repeated oscillations between anoxic and oxic conditions. Here, we report on results from an experiment where we monitored the abundance of bioreducible Fe(III) in NAu-1 over three consecutive redox cycles using chemical extractions and Fe isotope analysis to document the changes in the nature and extent of Fe atom exchange. During each cycle, NAu-1 was reduced biotically by Shewanella oneidensis MR-1 and then re-oxidized abiotically by O2 via aeration. By the third reduction period (RP3), the bacteria were only able to reduce 5.7% of the total clay Fe, that is, 40% less than during the first reduction period (RP1). The decrease in bioreducible Fe(III) is attributed to preferential reductive dissolution of Fe(III) from the finest clay particles. Extrapolation of the observed trend implies that, once the reducible Fe of the finest clay particles is removed, around 4% of the total Fe of the clay remains permanently redox-active, presumably as Fe atoms within the octahedral mineral structure that are accessible to the bacteria. The proposed particle size-dependent evolution of bioreducible Fe(III) from RP1 to RP3 is supported by the observed increasing crystalline domain size, preferential Fe dissolution from the smallest aggregates, and decreasing Fe isotope fractionation factors between aqueous Fe(II) and structural Fe(III) and between solid-bound Fe(II) and structural Fe(III). Our results imply that, in redox dynamic environments, the fraction of insoluble clay-bound Fe that is potentially renewable for use by Fe-reducing bacteria is a function of the evolving size distribution of the clay particles.

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

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

Mechanisms and rates of pyrite formation from hydrothermal fluid revealed by iron isotopes

Critical to interpreting iron isotope signatures in hydrothermal sulfide minerals is the accurate knowledge of Fe isotope fractionation factors, both kinetic and at equilibrium, between the fluid and minerals. Here we report Fe isotope fractionation between pyrite and aqueous saline fluid from in-situ pyrite precipitation experiments performed at 400–450 °C and 400–800 bar in a wide range of pH (2.5–4.8) and salt content (0.05–2.0 mol NaCl/KCl per kg of fluid). Our measurements show that while apparent bulk chemical equilibrium (in terms of total dissolved Fe concentration) between the FeCl2-bearing fluid and crystalline pyrite is attained within a few days, the Fe isotope fractionation values between the fluid and mineral (expressed as Δ57Fefluid-pyrite = δ57Fefluid – δ57Fepyrite) exhibit large variations, from about –0.6 to +1.0‰, depending on the elapsed time (1–30 days) and fluid composition (mostly pH, sulfur solubility, and di- and tri-sulfur radical ion concentration). Yet these values remain significantly higher, by 0.5–2‰, than equilibrium isotope fractionation inferred from β-factors of pyrite and the dominant iron chloride species in solution FeCl2(H2O)20(aq), generated in this study using Density Functional Theory (DFT) modeling (Δ57FeFeCl2(H2O)2-pyrite ≈ –1.25 ± 0.25‰ at 300–450 °C). The lack of isotope equilibrium between fluid and pyrite in this study and other recent work demonstrates that the fluid-pyrite Fe isotope exchange is kinetically controlled even at elevated temperatures (≥300 °C), allowing rate patterns to be identified. Our time-series data, combined with other available measurements at 300–350 °C, provide evidence for two major distinct kinetic regimes of fluid-pyrite Fe isotope exchange: (1) a fast and short (<10–20 days) initial step, likely controlled by Fe (poly)sulfide aqueous precursors that transfer their isotope signature to rapidly precipitating pyrite resulting in apparent isotopic disequilibrium between the FeCl2-dominated fluid and pyrite, (2) followed by a slower and longer second step of pyrite recrystallization, growth, and isotope re-equilibration with the fluid. The derived Fe isotope exchange rate constants suggest that more than 1 year is required to isotopically equilibrate pyrite with hydrothermal fluid at temperatures of 300–450 °C. Our new data provide direct interpretation of the Fe isotope ratios in fluids and sulfide minerals from modern submarine hydrothermal vents, which are characterized by fast (hours to days) pyrite precipitation. Furthermore, our results may enable estimating, using Fe isotope signatures, the dynamics of relatively short (<year-scale) mineralizing events in ancient hydrothermal metal sulfide deposits that cannot be assessed using more traditional dating approaches.

Magnesium and iron isotopic evidence of inter-mineral diffusion in ultramafic cumulates of the Peridotite Zone, Stillwater Complex

To better understand the role and extent of inter-mineral diffusion during subsolidus processes, we investigated the iron and magnesium isotopic compositions of olivine, orthopyroxene, and chromite separates from the Peridotite Zone of the Stillwater Complex. Olivine and orthopyroxene show limited Mg isotopic variations, with δ26Mg values ranging from −0.40 to −0.26‰ and from −0.29 to −0.22‰, respectively. In contrast, chromite displays extremely large Mg isotopic variations, with δ26Mg values ranging from −0.05 to +0.84‰. The δ56Fe values of olivine and orthopyroxene range from 0.00 to +0.17‰ and −0.04 to +0.06‰, respectively, whereas those of chromite range from −0.09 to +0.13‰. Most olivine-orthopyroxene pairs in our samples show no clear Fe-Mg isotopic differences, whereas silicate and coexisting chromite pairs plot away from theoretically predicted equilibrium lines, indicating disequilibrium inter-mineral fractionation. Higher δ56Fe values in silicates than in chromite and significantly higher δ26Mg values in chromite than in silicates indicate that subsolidus re-equilibration is responsible for the observed disequilibrium Fe-Mg isotope fractionation. We develop a quantitative model based on the investigated olivine and chromite to reproduce the observed kinetic Fe-Mg isotope fractionations. Our model results indicate a systematic difference between the Fe and Mg isotope fractionation factors, β, of olivine (βFe ≈ 0.27 and βMg ≈ 0.16) and chromite (0.03 < βFe < 0.08 and 0.05 < βMg < 0.16). This study demonstrates that combined investigations of the Fe-Mg isotopic compositions of cumulus minerals provide more recognizable and more reliable information on inter-mineral diffusion among igneous minerals in magma chambers than elemental analyses alone.

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

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

Iron isotopic and chemical tracing of basalt alteration and hematite spherule formation in Hawaii: A prospective study for Mars

Hematite spherules from Mauna Kea volcano (Hawaii) resemble Martian hematite spherules at Meridiani Planum in morphology, mineralogy, and geochemistry, and they both formed in association with acid-sulfate alteration of basalt. To gain insights into the processes involved in the formation of these spherules, we have measured the Fe isotopic compositions of unaltered basalts and tephras, as well as a variety of weathering/alteration products including altered basalts, hematite-rich breccia and individual hematite spherules, jarosite-bearing tephras, palagonitic tephras, calcined tephras (high-T dry oxidation) and steam-vent tephras.Compared to other alteration processes that induce little Fe mobility, acid-sulfate alteration produces large Fe isotopic variations (δ56Fe values ranging from −0.15 to +0.94‰). We used rock slabs from a variably altered basaltic rock to constrain the Fe isotope fractionation factor during acid-sulfate alteration of basaltic rocks. We found that the process followed a Rayleigh distillation with an instantaneous fractionation factor of −0.24‰ between the fluid and the residue for δ56Fe, resulting in a heavy Fe isotope enrichment in the alteration residue. However, the constant Mn/Fe ratios of the slabs suggest that acid-sulfate leaching does not significantly fractionate Mn/Fe ratios in basalts, which is consistent with the mobilization of iron as Fe(II).We also measured the Fe isotopic compositions and Mn/Fe ratios (using NanoSIMS) of individual hematite spherules from a spherule-rich breccia HWMK745R. The spherules show heavy Fe isotopic enrichments and low Mn/Fe ratios compared to unaltered samples, which is similar to the fractionations encountered in iron formations. A Rayleigh fractionation model suggests that both Fe isotopic compositions and Mn/Fe ratios of the spherules are consistent with ∼80% Fe oxidation and precipitation from an acid-sulfate fluid. Therefore, direct Fe precipitation from a fluid can readily explain the hematite spherule formation, and a more complex scenario involving jarosite precipitation followed by conversion to hematite is not needed. Combined thermodynamic and kinetic constraints also suggest that precipitation of jarosite would be hindered. The study presents a roadmap for studying future returned Martian samples in the laboratory.

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

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

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

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

Controls on the iron isotopic composition of global arc magmas

We determined the iron isotope composition of 130 mafic lavas from 15 arcs worldwide with the hypothesis that the results would reflect the relatively high oxidation state of arc magmas. Although this expectation was not realized, this Fe isotope data set reveals important insights into the geodynamic controls and style of the melting regimes in the sub-arc mantle. Samples are from oceanic arcs from the circum-Pacific, the Indonesian Sunda–Banda islands, Scotia and the Lesser Antilles as well as from the eastern Pacific Cascades. Their mean δ57Fe value is +0.075 ± 0.05‰, significantly lighter than MORB (+0.15 ± 0.03‰). Western Pacific arcs extend to very light δ57Fe (Kamchatka = −0.11 ± 0.04‰). This is contrary to expectation, because Fe isotope fractionation factors (Sossi et al., 2016, 2012) and the incompatibility of ferric versus ferrous iron during mantle melting, predict that melts of more oxidized sources will be enriched in heavy Fe isotopes. Subducted oxidation capacity flux may correlate with hydrous fluid release from the slab. If so, a positive correlation between each arc's thermal parameter (ϕ) and δ57Fe is predicted. On the contrary, the sampled arcs mostly contribute to a negative array with the ϕ value. High ϕ arcs, largely in the western Pacific, have primary magmas with lower δ57Fe values than the low ϕ, eastern Pacific arcs.Arcs with MORB-like Sr-, Nd- and Pb-isotopes, show a large range of δ57Fe from heavy MORB-like values (Scotia or the Cascades) to very light values (Kamchatka, Tonga). Although all basalts with light δ57Fe values have MORB-like Pb-, Nd- and Sr-isotope ratios some, particularly those from eastern Indonesia, have heavier δ57Fe and higher Pb- and Sr- and lower Nd-isotope ratios reflecting sediment contamination of the mantle wedge. Because basalts with MORB-like radiogenic isotopes range all the way from heavy to light δ57Fe values this trend is process-, not source composition-driven. Neither the slab-derived influx of fluids with light iron or sediment-derived melts with heavier iron can drive the iron isotopic shifts. The trend to light iron isotopes is partly the result of repeated, hydrous flux-driven, fO2-buffered, melting of initially normal-DMM-like mantle. However the most negative δ57Fe must also reflect re-melting of sources that have experienced prior diffusive (disequilibrium) stripping of heavy Fe isotopes due to rapid melt extraction and metasomatism.Data from intra-arc to back-arc rifts in the western Pacific show that these arc signatures are rapidly dispersed by influx of DMM or OIB mantle once intra- and back-arc rifting and slab rollback gains momentum. We suggest that the characteristic light arc signatures only form when the source is lodged under arcs where sub-arc mantle undergoes corner flow forming an isolated roll. This process of heavy iron depletion is most efficient in the high ϕ arcs of the western Pacific and least prevalent in the low ϕ arcs of the eastern Pacific where δ57Fe values are MORB-like. This implies that there is a fundamental change in character of sub-arc mantle melting between east and west Pacific, percolative and fluid fluxed in the west and diapiric and decompressional in the east.

Geochemical and iron isotopic insights into hydrothermal iron oxyhydroxide deposit formation at Loihi Seamount

Low-temperature hydrothermal vents, such as those encountered at Loihi Seamount, harbor abundant microbial communities and provide ideal systems to test hypotheses on biotic versus abiotic formation of hydrous ferric oxide (FeOx) deposits at the seafloor. Hydrothermal activity at Loihi Seamount produces abundant microbial mats associated with rust-colored FeOx deposits and variably encrusted with Mn-oxyhydroxides. Here, we applied Fe isotope systematics together with major and trace element geochemistry to study the formation mechanisms and preservation of such mineralized microbial mats. Iron isotope composition of warm (<60 °C), Fe-rich and H2S-depleted hydrothermal fluids yielded δ56Fe values near +0.1‰, indistinguishable from basalt values. Suspended particles in the vent fluids and FeOx deposits recovered nearby active vents yielded systematically positive δ56Fe values. The enrichment in heavy Fe isotopes between +1.05‰ and +1.43‰ relative to Fe(II) in vent fluids suggest partial oxidation of Fe(II) during mixing of the hydrothermal fluid with seawater. By comparing the results with experimentally determined Fe isotope fractionation factors, we determined that less than 20% of Fe(II) is oxidized within active microbial mats, although this number may reach 80% in aged or less active deposits. These results are consistent with Fe(II) oxidation mediated by microbial processes considering the expected slow kinetics of abiotic Fe oxidation in low oxygen bottom water at Loihi Seamount. In contrast, FeOx deposits recovered at extinct sites have distinctly negative Fe-isotope values down to −1.77‰ together with significant enrichment in Mn and occurrence of negative Ce anomalies. These results are best explained by the near-complete oxidation of an isotopically light Fe(II) source produced during the waning stage of hydrothermal activity under more oxidizing conditions. Light Fe isotope values of FeOx are therefore generated by subsurface precipitation of isotopically heavy Fe-oxides rather than by the activity of dissimilatory Fe reduction in the subsurface. Overall, Fe-isotope compositions of microbial mats at Loihi Seamount display a remarkable range between −1.2‰ and +1.6‰ which indicate that Fe isotope compositions of hydrothermal Fe-oxide precipitates are particularly sensitive to local environmental conditions where they form, and are less sensitive to abiotic versus biotic origins. It follows that FeOx deposits at Loihi Seamount provides important modern analogues for ancient seafloor Fe-rich deposits allowing for testing hypotheses about the biogeochemical cycling of Fe isotopes on early Earth.

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

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

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

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

Iron Isotope Characteristics of Hot Springs at Chocolate Pots, Yellowstone National Park

Chocolate Pots Hot Springs in Yellowstone National Park is a hydrothermal system that contains high aqueous ferrous iron [∼0.1 mM Fe(II)] at circumneutral pH conditions. This site provides an ideal field environment in which to test our understanding of Fe isotope fractionations derived from laboratory experiments. The Fe(III) oxides, mainly produced through Fe(II) oxidation by oxygen in the atmosphere, have high ⁵⁶Fe/⁵⁴Fe ratios compared with the aqueous Fe(II). However, the degree of fractionation is less than that expected in a closed system at isotopic equilibrium. We suggest two explanations for the observed Fe isotope compositions. One is that light Fe isotopes partition into a sorbed component and precipitate out on the Fe(III) oxide surfaces in the presence of silica. The other explanation is internal regeneration of isotopically heavy Fe(II) via dissimilatory Fe(III) reduction farther down the flow path as well as deeper within the mat materials. These findings provide evidence that silica plays an important role in governing Fe isotope fractionation factors between reduced and oxidized Fe. Under conditions of low ambient oxygen, such as may be found on early Earth or Mars, significantly larger Fe isotope variations are predicted, reflecting the more likely attainment of Fe isotope equilibrium associated with slower oxidation rates under low-O₂ conditions.

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

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

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 (δ 57Fe; 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 (δ 65Cu; permil deviation in 65Cu/ 63Cu relative to the pure copper standard NIST 976) of both metals (δ 65Cu M) and sulphides (δ 65Cu FeS) span much wider ranges (−9.30 to 0.99‰ and −8.90 to 0.63‰, respectively) than reported previously. Metal–troilite fractionation factors (Δ 65Cu M–FeS = δ 65Cu M − δ 65Cu 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 (δ 57Fe FeS) isotope compositions and between metal–troilite Cu and Fe (Δ 57Fe 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 δ 65Cu FeS and δ 57Fe FeS and calculated Cu metal–sulphide partition coefficients (D Cu = [Cu] metal/[Cu] FeS) are also present. A strong negative correlation exists between Δ 57Fe M–FeS and D Cu; a more diffuse positive array is defined by Δ 65Cu 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 Δ 57Fe M–FeS and the lowest Δ 65Cu 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 Δ 57Fe 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 Δ 57Fe M–FeS and largest D Cu values. The range in Δ 65Cu 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 δ 65Cu 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 Δ 65Cu M–FeS and Δ 57Fe 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 Δ 65Cu M–FeS and Δ 57Fe 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.

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

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

Iron isotope fractionation between aqueous ferrous iron and goethite

The equilibrium Fe isotope fractionation factor between aqueous Fe(II) and goethite has been experimentally measured to be −1.05±0.08‰ in 56Fe/54Fe (2σ) at 22°C, using the three-isotope method. Experiments were done using two sizes of goethite (81×11nm and 590×42nm), and the experimental products were subjected to serial extraction using acid partial dissolution techniques to determine if surface Fe(III) atoms have different isotopic properties than the bulk goethite. These experiments indicate that the interaction of Fe(II)aq and goethite is dynamic and results in complete or near-complete Fe isotope exchange over 30days, involving at least four components: Fe(II)aq, goethite, sorbed Fe(II), and Fe(III)surface. The equilibrium fractionation factor between Fe(II)aq and Fe(II)sorb is the same for both sizes of goethite, at Δ56FeFe(II)aq–Fe(II)sorb=−1.24±0.14‰; this fractionation factor is significantly different than the results of previous studies on Fe(II) sorption to goethite. The proportion of the Fe(III)surface component is greatest in the experiments that used the smallest goethite, and the Fe(III)surface–Fe(II)aq fractionation is estimated to be at least +2.1‰. The high Fe(III)surface–Fe(II)aq fractionation may exert a significant influence on the Fe isotope compositions of aqueous Fe(II) in natural systems that contain nanoparticulate goethite, including those involving bacterial iron reduction. These results demonstrate that the isotopic properties of nano-scale minerals may be distinct from micron-scale or larger minerals, as is the case for other thermodynamic properties of nanoparticles.

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 δ 56Fe values under conditions of variable pH and dissolved Si, similar to the large inventory of negative δ 56Fe in Neoarchean and Paleoproterozoic age marine sedimentary rocks.

Iron and lithium isotope systematics of the Hekla volcano, Iceland — Evidence for Fe isotope fractionation during magma differentiation

In this study potential iron isotope fractionation by magmatic processes in the Earth's crust was systematically investigated. High precision iron isotope analyses by MC-ICP-MS were performed on a suite of rock samples representative for the volcanic evolution of the Hekla volcano, Iceland. The whole series of Hekla's rocks results from several processes. (i) Basaltic magmas rise and induce partial melting of meta-basalts in the lower part of the Icelandic crust. The resulting dacitic magma evolves to rhyolitic composition through crystal fractionation. During this differentiation the δ 56/54Fe IRMM-014 values increase successively from 0.051 ± 0.021‰ for the primitive dacites to 0.168 ± 0.021‰ for the rhyolites. This increase can be described by a Rayleigh fractionation model using a constant bulk fractionation factor between all mineral phases (M) and the silicate liquid (L) of Δ 56/54Fe M–L = − 0.1‰. (ii) The basaltic magma itself differentiates by crystal fractionation to basaltic andesite composition. No Fe isotope fractionation was found in this series. All basalts and basaltic andesites have an average δ 56/54Fe IRMM-014 value of 0.062 ± 0.042‰ (2SD, n = 9), identical to mean terrestrial basaltic values reported in previous studies. This observation is consistent with the limited removal of iron from the remaining silicate melt through crystal fractionation and small mineral-melt Fe isotope fractionation factors expected at temperatures in excess of 1050 °C. (iii) Andesites are produced by mixing of basaltic andesite with dacitic melts. The iron isotope composition of the andesites is matching that of the basaltic andesites and the less evolved dacites, in agreement with a mixing process. In the Hekla volcanic suite Li concentrations are positively correlated with indicators of magma differentiation. All Hekla rocks have δ 7Li values typical for the upper mantle and demonstrate the absence of resolvable Li isotope fractionation during crystal fractionation. As a fluid-mobile trace element, Li concentrations and isotopes are a potential tracer of magma/fluid interaction. At Hekla, Li concentrations and isotope compositions do not indicate any extensive fluid exsolution. Hence, the heavy Fe isotope composition of the dacites and rhyolites can be predominately attributed to fractional crystallisation. Iron isotope analyses on single samples from other Icelandic volcanoes (Torfajökull, Vestmannaeyjar) confirm heavy Fe isotope enrichment in evolving magmas. Our results suggest that the iron isotope composition of highly evolved crust can be slightly modified by magmatic processes.