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.     

Copper isotope fractionation during surface adsorption and intracellular incorporation by bacteria

Copper isotopes may prove to be a useful tool for investigating bacteria-metal interactions recorded in natural waters, soils, and rocks. However, experimental data which attempt to constrain Cu isotope fractionation in biologic systems are limited and unclear. In this study, we utilized Cu isotopes (δ⁶⁵Cu) to investigate Cu-bacteria interactions, including surface adsorption and intracellular incorporation. Experiments were conducted with individual representative species of Gram-positive (Bacillus subtilis) and Gram-negative (Escherichia coli) bacteria, as well as with wild-type consortia of microorganisms from several natural environments. Ph-dependent adsorption experiments were conducted with live and dead cells over the pH range 2.5-6. Surface adsorption experiments of Cu onto live bacterial cells resulted in apparent separation factors (Δ⁶⁵Cu_{solution-solid} = δ⁶⁵Cu_solution − δ⁶⁵Cu_solid) ranging from +0.3‰ to +1.4‰ for B. subtilis and +0.2‰ to +2.6‰ for E. coli. However, because heat-killed bacterial cells did not exhibit this behavior, the preference of the lighter Cu isotope by the cells is probably not related to reversible surface adsorption, but instead is a metabolically-driven phenomenon. Adsorption experiments with heat-killed cells yielded apparent separation factors ranging from +0.3‰ to −0.69‰ which likely reflects fractionation from complexation with organic acid surface functional group sites. For intracellular incorporation experiments the lab strains and natural consortia preferentially incorporated the lighter Cu isotope with an apparent Δ⁶⁵Cu_{solution-solid} ranging from ∼+1.0‰ to +4.4‰. Our results indicate that live bacterial cells preferentially sequester the lighter Cu isotope regardless of the experimental conditions. The fractionation mechanisms involved are likely related to active cellular transport and regulation, including the reduction of Cu(II) to Cu(I). Because similar intracellular Cu machinery is shared by fungi, plants, and higher organisms, the influence of biological processes on the δ⁶⁵Cu of natural waters and soils is probably considerable.     

Exploration potential of Cu isotope fractionation in porphyry copper deposits

We examined the copper isotope ratio of primary high temperature Cu-sulfides, secondary low temperature Cu-sulfides (and Cu-oxides) as well as Fe-oxides in the leach cap, which represent the weathered remains of a spectrum of Cu mineralization, from nine porphyry copper deposits. Copper isotope ratios are reported as δ⁶⁵Cu‰ = ((⁶⁵Cu/⁶³Cu_sample/⁶⁵Cu/⁶³Cu_{NIST 976 standard}) − 1) ? 10³. Errors for all the analyses are ± 0.14‰ (determined by multiple analyses of the samples) and mass bias was corrected through standard-sample-standard bracketing. The overall isotopic variability measured in these samples range from − 16.96‰ to 9.98‰.     

Cu isotopic fractionation in the supergene environment with and without bacteria

The isotopic composition of dissolved Cu and solid Cu-rich minerals [δ⁶⁵Cu (‰) = (⁶⁵Cu/⁶³Cu_sample/⁶⁵Cu/⁶³Cu_std) - 1)*1000] were monitored in batch oxidative dissolution experiments with and without Thiobacillus ferrooxidans. Aqueous copper in leach fluids released during abiotic oxidation of both chalcocite and chalcopyrite was isotopically heavier (δ⁶⁵Cu = 5.34‰ and δ⁶⁵Cu = 1.90‰, respectively, [±0.16 at 2σ]) than the initial starting material (δ⁶⁵Cu = 2.60 ± 0.16‰ and δ⁶⁵Cu = 0.58 ± 0.16‰, respectively). Isotopic mass balance between the starting material, aqueous copper, and secondary minerals precipitated in these experiments explains the heavier isotopic values of aqueous copper. In contrast, aqueous copper from leached chalcocite and chalcopyrite inoculated with Thiobacillus ferrooxidans was isotopically similar to the starting material. The lack of fractionation of the aqueous copper in the biotic experiments can best be explained by assuming a sink for isotopically heavy copper present in the bacteria cells with δ⁶⁵Cu = 5.59 ± 0.16‰. Consistent with this inference, amorphous Cu-Fe oxide minerals are observed surrounding cell membranes of Thiobacillus grown in the presence of dissolved Cu and Fe.Extrapolating these experiments to natural supergene environments implies that release of isotopically heavy aqueous Cu from oxidative leach caps, especially under abiotic conditions, should result in precipitates in underlying enrichment blankets that are isotopically heavy. Where iron-oxidizing cells are involved, isotopically heavy oxidized Cu entrained in cellular material may become associated with leach caps, causing the released aqueous Cu to be less isotopically enriched in the heavy isotope than predicted for the abiotic system. Rayleigh fractionation trends with fractionation factors calculated from our experiments for both biotic and abiotic conditions are consistent with large numbers of individual abiotic or biotic leaching events, explaining the supergene chalcocites in the Morenci and Silver Bell porphyry copper deposits.     

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.     

Copper isotope fractionation in acid mine drainage

We measured the Cu isotopic composition of primary minerals and stream water affected by acid mine drainage in a mineralized watershed (Colorado, USA). The δ⁶⁵Cu values (based on ⁶⁵Cu/⁶³Cu) of enargite (δ⁶⁵Cu = −0.01 ± 0.10‰; 2σ) and chalcopyrite (δ⁶⁵Cu = 0.16 ± 0.10‰) are within the range of reported values for terrestrial primary Cu sulfides (−1‰ < δ⁶⁵Cu < 1‰). These mineral samples show lower δ⁶⁵Cu values than stream waters (1.38‰ ? δ⁶⁵Cu ? 1.69‰). The average isotopic fractionation (Δ_aq-min = δ⁶⁵Cu_aq − δ⁶⁵Cu_min, where the latter is measured on mineral samples from the field system), equals 1.43 ± 0.14‰ and 1.60 ± 0.14‰ for chalcopyrite and enargite, respectively. To interpret this field survey, we leached chalcopyrite and enargite in batch experiments and found that, as in the field, the leachate is enriched in ⁶⁵Cu relative to chalcopyrite (1.37 ± 0.14‰) and enargite (0.98 ± 0.14‰) when microorganisms are absent. Leaching of minerals in the presence of Acidithiobacillus ferrooxidans results in smaller average fractionation in the opposite direction for chalcopyrite (Δ_aq-min^o=-0.57±0.14‰, where min^o refers to the starting mineral) and no apparent fractionation for enargite (Δ_aq-min^o=0.14±0.14‰). Abiotic fractionation is attributed to preferential oxidation of ⁶⁵Cu⁺ at the interface of the isotopically homogeneous mineral and the surface oxidized layer, followed by solubilization. When microorganisms are present, the abiotic fractionation is most likely not seen due to preferential association of ⁶⁵Cu_aq with A. ferrooxidans cells and related precipitates. In the biotic experiments, Cu was observed under TEM to occur in precipitates around bacteria and in intracellular polyphosphate granules. Thus, the values of δ⁶⁵Cu in the field and laboratory systems are presumably determined by the balance of Cu released abiotically and Cu that interacts with cells and related precipitates. Such isotopic signatures resulting from Cu sulfide dissolution should be useful for acid mine drainage remediation and ore prospecting purposes.     

Isotopic constraints on ice age fluids in active geothermal systems: Reykjanes, Iceland

The Reykjanes geothermal system is located on the landward extension of the Mid-Atlantic Ridge in southwest Iceland, and provides an on-land proxy to high-temperature hydrothermal systems of oceanic spreading centers. Previous studies of elemental composition and salinity have shown that Reykjanes geothermal fluids are likely hydrothermally modified seawater. However, δD values of these fluids are as low as −23‰, which is indicative of a meteoric water component. Here we constrain the origin of Reykjanes hydrothermal solutions by analysis of hydrogen and oxygen isotope compositions of hydrothermal epidote from geothermal drillholes at depths between 1 and 3 km. δD_EPIDOTE values from wells RN-8, -9, -10 and -17 collectively range from −60 to −78‰, and δ¹⁸O_EPIDOTE in these wells are between −3.0 and 2.3‰. The δD values of epidote generally increase along a NE trend through the geothermal field, whereas δ¹⁸O values generally decrease, suggesting a southwest to northeast migration of the geothermal upflow zone with time that is consistent with present-day temperatures and observed hydrothermal mineral zones. For comparative analysis, the meteoric-water dominated Nesjavellir and Krafla geothermal systems, which have a δD_FLUID of ∼ −79‰ and −89‰, respectively, show δD_EPIDOTE values of −115‰ and −125‰. In contrast, δD_EPIDOTE from the mixed meteoric-seawater Svartsengi geothermal system is −68‰; comparable to δD_EPIDOTE from well RN-10 at Reykjanes.Stable isotope compositions of geothermal fluids in isotopic equilibrium with the epidotes at Reykjanes are computed using published temperature dependent hydrogen and oxygen isotope fractionation curves for epidote-water, measured isotope composition of the epidotes and temperatures approximated from the boiling point curve with depth. Calculated δD and δ¹⁸O of geothermal fluids are less than 0‰, suggesting that fluids of meteoric or glacial origin are a significant component of the geothermal solutions. Additionally, δD_FLUID values in equilibrium with geothermal epidote are lower than those of modern-day fluids, whereas calculated δ¹⁸O_FLUID values are within range of the observed fluid isotope composition. We propose that modern δD_EPIDOTE and δD_FLUID values are the result of diffusional exchange between hydrous alteration minerals that precipitated from glacially-derived fluids early in the evolution of the Reykjanes system and modern seawater-derived geothermal fluids. A simplified model of isotope exchange in the Reykjanes geothermal system, in which the average starting δD_ROCK value is −125‰ and the water to rock mass ratio is 0.25, predicts a δD_FLUID composition within 1‰ of average measured values. This model resolves the discrepancy between fluid salinity and isotope composition of Reykjanes geothermal fluids, explains the observed disequilibrium between modern fluids and hydrothermal epidote, and suggests that rock-fluid interaction is the dominant control over the evolution of fluid isotope composition in the hydrothermal system.     

High precision iron isotope measurements of meteoritic material by cold plasma ICP-MS

The first cold plasma ICP-MS (inductively coupled plasma mass spectrometer) Fe isotope study is described. Application of this technique to the analyses of Fe isotopes in a number of meteorites is also reported. The measurement technique relies on reduced temperature operation of the ICP source to eliminate pervasive molecular interferences from Ar complexes associated with conventional ICP-MS. Instrumental mass bias corrections are performed by sample-standard bracketing and using Cu as an external mass bias drift monitor. Repeated measurements of a terrestrial basalt reference sample indicate an external reproducibility of ± 0.06 ‰ for δ⁵⁶Fe and ± 0.25 ‰ for δ⁵⁸Fe (1 σ). The measured iron isotopic compositions of various bulk meteorites, including irons, chondrites and pallasites are identical, within error, to the composition of our terrestrial basalt reference sample suggesting that iron mass fractionation during planet formation and differentiation was non-existent. Iron isotope compositions measured for eight chondrules from the unequilibrated ordinary chondrite Tieschitz range from −0.5 ‰ < δ⁵⁶Fe_chondrules < 0.0 ‰ relative to the terrestrial/meteorite average. Mechanisms for fractionating iron in these chondrules are discussed.     

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

Here we compare new experimental studies with theoretical predictions of equilibrium iron isotopic fractionation among aqueous ferric chloride complexes (Fe(H₂O)₆³⁺, FeCl(H₂O)₅²⁺, FeCl₂(H₂O)₄⁺, FeCl₃ (H₂O)₃, and FeCl₄^-), using the Fe-Cl-H₂O system as a simple, easily-modeled example of the larger variety of iron-ligand compounds, such as chlorides, sulfides, simple organic acids, and siderophores. Isotopic fractionation (⁵⁶Fe/⁵⁴Fe) among naturally occuring iron-bearing species at Earth surface temperatures (up to ∼3‰) is usually attributed to redox effects in the environment. However, theoretical modeling of reduced isotopic partition functions among iron-bearing species in solution also predicts fractionations of similar magnitude due to non-redox changes in speciation (i.e., ligand bond strength and coordination number). In the present study, fractionations are measured in a series of low pH ([H⁺] = 5 M) solutions of ferric chloride (total Fe = 0.0749 mol/L) at chlorinities ranging from 0.5 to 5.0 mol/L. Advantage is taken of the unique solubility of FeCl₄^- in immiscible diethyl ether to create a separate spectator phase, used to monitor changing fractionation in the aqueous solution. Δ⁵⁶Fe_aq-eth = δ⁵⁶Fe (total Fe remaining in aqueous phase)−δ⁵⁶Fe (FeCl₄^- in ether phase) is determined for each solution via MC-ICPMS analysis.Both experiments and theoretical calculations of Δ⁵⁶Fe_aq-eth show a downward trend with increasing chlorinity: Δ⁵⁶Fe_aq-eth is greatest at low chlorinity, where FeCl₂(H₂O)₄⁺ is the dominant species, and smallest at high chlorinity where FeCl₃(H₂O)₃ is dominant. The experimental Δ⁵⁶Fe_aq-eth ranges from 0.8‰ at [Cl^-] = 0.5 M to 0.0‰ at [Cl^-] = 5.0 M, a decrease in aqueous-ether fractionation of 0.8‰. This is very close to the theoretically predicted decreases in Δ⁵⁶Fe_aq-eth, which range from 1.0 to 0.7‰, depending on the ab initio model.The rate of isotopic exchange and attainment of equilibrium are shown using spiked reversal experiments in conjunction with the two-phase aqueous-ether system. Equilibrium under the experimental conditions is established within 30 min.The general agreement between theoretical predictions and experimental results points to substantial equilibrium isotopic fractionation among aqueous ferric chloride complexes and a decrease in ⁵⁶Fe/⁵⁴Fe as the Cl^-/Fe³⁺ ion ratio increases. The effects on isotopic fractionation shown by the modeling of this simple iron-ligand system imply that ligands present in an aqueous environment are potentially important drivers of fractionation, are indicative of possible fractionation effects due to other speciation effects (such as iron-sulfide systems or iron bonding with organic ligands), and must be considered when interpreting iron isotope fractionation in the geological record.     

Stable Cu and Zn isotope ratios as tracers of sources and transport of Cu and Zn in contaminated soil

Copper and Zn metals are produced in large quantities for different applications. During Cu production, large amounts of Cu and Zn can be released to the environment. Therefore, the surroundings of Cu smelters are frequently metal-polluted. We determined Cu and Zn concentrations and Cu and Zn stable isotope ratios (δ⁶⁵Cu, δ⁶⁶Zn) in three soils at distances of 1.1, 3.8, and 5.3 km from a Slovak Cu smelter and in smelter wastes (slag, sludge, ash) to trace sources and transport of Cu and Zn in soils. Stable isotope ratios were measured by multicollector inductively coupled plasma mass spectrometry (MC-ICP-MS) in total digests. Soils were heavily contaminated with concentrations up to 8087 μg g⁻¹ Cu and 2084 μg g⁻¹ Zn in the organic horizons. The δ⁶⁵Cu values varied little (−0.12‰ to 0.36‰) in soils and most wastes and therefore no source identification was possible. In soils, Cu became isotopically lighter with increasing depth down to 0.4 m, likely because of equilibrium reactions between dissolved and adsorbed Cu species during transport of smelter-derived Cu through the soil. The δ⁶⁶Zn_IRMM values were isotopically lighter in ash (−0.41‰) and organic horizons (−0.85‰ to −0.47‰) than in bedrock (−0.28‰) and slag (0.18‰) likely mainly because of kinetic fractionation during evaporation and thus allowed for separation of smelter-Zn from native Zn in soil. In particular in the organic horizons large variations in δ⁶⁶Zn values occur, probably caused by biogeochemical fractionation in the soil-plant system. In the mineral horizons, Zn isotopes showed only minor shifts to heavier δ⁶⁶Zn values with depth mainly because of the mixing of smelter-derived Zn and native Zn in the soils. In contrast to Cu, Zn isotope fractionation between dissolved and adsorbed species was probably only a minor driver in producing the observed variations in δ⁶⁶Zn values. Our results demonstrate that metal stable isotope ratios may serve as tracer of sources, vertical dislocation, and biogeochemical behavior in contaminated soil.     

Stable Cu isotope fractionation in soils during oxic weathering and podzolization

Copper stable isotope ratios are fractionated during various biogeochemical processes and may trace the fate of Cu during long-term pedogenetic processes. We assessed the effects of oxic weathering (formation of Cambisols) and podzolization on Cu isotope ratios (δ⁶⁵Cu). Two Cambisols (oxic weathered soils without strong vertical translocations of soil constituents) and two Podzols (soils showing vertical translocation of organic matter, Fe and Al) were analyzed for Cu concentrations, partitioning of Cu in seven fractions of a sequential extraction and δ⁶⁵Cu values in bulk soil. Cu concentrations in the studied soils were low (1.4-27.6 μg g⁻¹) and Cu was mainly associated with strongly bound Fe oxide- and silicate-associated forms. Bulk δ⁶⁵Cu values varied between −0.57‰ and 0.44‰ in all studied horizons. The O horizons had on average significantly lighter Cu isotope compositions (−0.21‰) than the A horizons (0.13‰) which can either be explained by Cu isotope fractionation during cycling through the plants or deposition of isotopically light Cu from the atmosphere. Oxic weathering without pronounced podzolization in both Cambisols and a weakly developed Podzol (Haplic Podzol 2) caused no significant isotope fractionation in the single profiles, while a slight tendency to lower δ⁶⁵Cu values with depth was visible in all four profiles. This is the opposite depth distribution of δ⁶⁵Cu values to that we observed in hydromorphic soils (soils which show indication of redox changes because of the influence of water saturation) in a previous study. In a more pronounced Podzol (Haplic Podzol 1), δ⁶⁵Cu values and Cu concentrations decreased from Ah to E horizons and increased again deeper in the soil. Humus-rich sections of the Bhs horizon had higher Cu concentrations (2.8 μg g⁻¹) and a higher δ⁶⁵Cu value (−0.18‰) than oxide-rich sections (1.9 μg g⁻¹, −0.35‰) suggesting Cu translocation between E and B horizons as organo-Cu complexes. The different depth distributions in oxic weathered and hydromorphic soils and the pronounced vertical differences in δ⁶⁵Cu values in Haplic Podzol 1 indicate a promising potential of δ⁶⁵Cu values to improve our knowledge of the fate of Cu during long-term pedogenetic processes.     

Zn and Cu isotopic variations in chondrites and iron meteorites: Early solar nebula reservoirs and parent-body processes

High-precision Zn isotopic variations are reported for carbonaceous chondrites (CC), equilibrated (EOC) and unequilibrated (UOC) ordinary chondrites, iron meteorites from the IAB-IIICD (nonmagmatic) and IIIA (magmatic) groups, and metal from the Brenham pallasite. For irons, δ⁶⁵Cu values are also reported. Data have also been obtained on a coarse-grained type-B calcium-, aluminum-rich refractory inclusion (CAI) from Allende and on acid leaches of Allende (CV3), Krymka (LL3), and Charsonville (H6). Variations expressed as δ⁶⁶Zn (deviation in parts per thousand of ⁶⁶Zn/⁶⁴Zn in samples relative to a standard) spread over a range of 0.3‰ for carbonaceous chondrites, 2‰ for ordinary chondrites, and 4‰ for irons.The measured ⁶⁶Zn/⁶⁴Zn, ⁶⁷Zn/⁶⁴Zn, and ⁶⁸Zn/⁶⁴Zn ratios vary linearly with mass difference and define a common isotope fractionation line with terrestrial samples, which demonstrates that Zn was derived from an initially single homogeneous reservoir. The δ⁶⁶Zn values are correlated with meteorite compositions and slightly decrease in the order CI, CM, CV-CO, and to UOC. The isotopically light Zn of Allende CAI and the acid-resistant residues of Allende and Krymka show that the light component is associated with refractory material, presumably minerals from the spinel-group. This, together with the reverse correlation between relative abundances of light Zn isotopes and volatile element abundances, suggests that Zn depletion in planetary bodies with respect to CI cannot be ascribed to devolatilization of CI-like material. These observations rather suggest that refractory material reacted with a gas phase enriched in the lighter Zn isotopes. Alternatively, chondrules with their associated rims should carry a light Zn isotopic signature. The δ⁶⁶Zn values of unequilibrated chondrites are rather uniform, whereas equilibrated chondrites show distinctly more isotopic variability.The values of δ⁶⁵Cu-δ⁶⁶Zn in irons define two trends. The moderate and positively correlated Cu and Zn isotope variations in IIIA and pallasite samples probably reflect crystallization of silicate, sulfide, and solid metal from the liquid metal. The range of δ⁶⁶Zn values of the IAB-IIICD group is large (>3‰) and contrasts with the moderate fractionation of Cu isotopes. We interpret this feature and the negative δ⁶⁶Zn-δ⁶⁵Cu correlation as reflecting mixing, possibly achieved by percolation, between metals from a regolith devolatilized at low temperature (enriched in heavy zinc) and metallic liquids formed within the parent body.     

Iron and magnesium isotopic compositions of peridotite xenoliths from Eastern China

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

Iron isotope fractionation 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.     

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.     

Ion-exchange fractionation of copper and zinc isotopes

Whether transition element isotopes can be fractionated at equilibrium in nature is still uncertain. Standard solutions of Cu and Zn were eluted on an anion-exchange resin, and the isotopic compositions of Cu (with respect to Zn) of the eluted fractions were measured by multiple-collector inductively coupled plasma mass spectrometry. It was found that for pure Cu solutions, the elution curves are consistent with a ⁶³Cu/⁶⁵Cu mass fractionation coefficient of 0.46‰ in 7 mol/L HCl and 0.67‰ in 3 mol/L HCl between the resin and the solution. Batch fractionation experiments confirm that equilibrium fractionation of Cu between resin and 7 mol/L HCl is ∼0.4‰ and therefore indicates that there is no need to invoke kinetic fractionation during the elution. Zn isotope fractionation is an order of magnitude smaller, with a ⁶⁶Zn/⁶⁸Zn fractionation factor of 0.02‰ in 12 mol/L HCl. Cu isotope fractionation results determined from a chalcopyrite solution in 7 mol/L HCl give a fractionation factor of 0.58‰, which indicates that Fe may interfere with Cu fractionation.Comparison of Cu and Zn results suggests that the extent of Cu isotopic fractionation may signal the presence of so far unidentified polynuclear complexes in solution. In contrast, we see no compelling reason to ascribe isotope fractionation to the coexistence of different oxidation states. We further suggest that published evidence for iron isotopic fractionation in nature and in laboratory experiments may indicate the distortion of low-spin Fe tetrahedral complexes.The isotope geochemistry of transition elements may shed new light on their coordination chemistry. Their isotopic fractionation in the natural environment may be interpreted using models of thermodynamic fractionation.     

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.     

Fractionation of Cu and Zn isotopes during adsorption onto amorphous Fe(III) oxyhydroxide: Experimental mixing of acid rock drainage and ambient river water

Fractionation of Cu and Zn isotopes during adsorption onto amorphous ferric oxyhydroxide is examined in experimental mixtures of metal-rich acid rock drainage and relatively pure river water and during batch adsorption experiments using synthetic ferrihydrite. A diverse set of Cu- and Zn-bearing solutions was examined, including natural waters, complex synthetic acid rock drainage, and simple NaNO₃ electrolyte. Metal adsorption data are combined with isotopic measurements of dissolved Cu (⁶⁵Cu/⁶³Cu) and Zn (⁶⁶Zn/⁶⁴Zn) in each of the experiments. Fractionation of Cu and Zn isotopes occurs during adsorption of the metal onto amorphous ferric oxyhydroxide. The adsorption data are modeled successfully using the diffuse double layer model in PHREEQC. The isotopic data are best described by a closed system, equilibrium exchange model. The fractionation factors (α_soln-solid) are 0.99927 ± 0.00008 for Cu and 0.99948 ± 0.00004 for Zn or, alternately, the separation factors (Δ_soln-solid) are −0.73 ± 0.08‰ for Cu and −0.52 ± 0.04‰ for Zn. These factors indicate that the heavier isotope preferentially adsorbs onto the oxyhydroxide surface, which is consistent with shorter metal-oxygen bonds and lower coordination number for the metal at the surface relative to the aqueous ion. Fractionation of Cu isotopes also is greater than that for Zn isotopes. Limited isotopic data for adsorption of Cu, Fe(II), and Zn onto amorphous ferric oxyhydroxide suggest that isotopic fractionation is related to the intrinsic equilibrium constants that define aqueous metal interactions with oxyhydroxide surface sites. Greater isotopic fractionation occurs with stronger metal binding by the oxyhydroxide with Cu > Zn > Fe(II).     

Coupled lithium- and iron isotope fractionation during magmatic differentiation

In this study, we investigated Fe and Li isotope fractionation between mineral separates of olivine pheno- and xenocrysts (including one clinopyroxyene phenocryst) and their basaltic hosts. Samples were collected from the Canary Islands (Teneriffa, La Palma) and some German volcanic regions (Vogelsberg, Westerwald and Hegau). All investigated bulk samples fall in a tight range of Li and Fe isotope compositions (δ⁵⁶Fe_wr = 0.06–0.17‰ and δ⁷Li_ma = 2.5–5.2‰, assuming δ⁷Li of the olivine-free matrix is virtually identical to that of the bulk sample for mass balance reasons). In contrast, olivine phenocrysts display highly variable, but generally light Fe and mostly light Li isotope compositions compared to their respective olivine-free basaltic matrix, which was considered to represent the melt (with δ⁵⁶Fe_ol = ? 0.24 to 0.14‰ and δ⁷Li_ol = ? 10.5 to + 6.5‰, respectively). Single olivine crystals from one sample display even a larger range of δ⁵⁶Fe_ol between ? 0.7 and + 0.1‰. One single clinopyroxene phenocryst displays the lightest Li isotope composition (δ⁷Li_cpx = ? 17.7‰), but no Fe isotope fractionation relative to melt. The olivine phenocrysts show variable Mg# and Ni (correlated in most cases) that range between 0.89 and 0.74 and between 300 and 3000 μg/g, respectively. These olivines likely grew by fractional crystallization in an evolving magma. One sample from the Vogelsberg volcano contained olivine xenocrysts (Mg# > 0.89 and Ni > 3000 μg/g), in addition to olivine phenocrysts. This sample displays the highest Li- and the second highest Fe-isotope fractionation between olivine and melt (Δ⁷Li_ol-melt = ? 13; Δ⁵⁶Fe_ol-melt = ? 0.29).Our data, i.e. the variable olivine- at constant whole rock and matrix isotope compositions, strongly indicate disequilibrium, i.e. kinetic Fe and Li isotope fractionation between olivine and melt (for Li also between cpx and melt) during fractional crystallization. Δ⁷Li_ol-melt is correlated with the Li partitioning between olivine and melt (i.e. with Li_ol/Li_melt), indicating Li isotope fractionation due to preferential (faster) diffusion of ⁶Li into olivine during fractional crystallization. Olivine with low Δ⁷Li_ol-melt, also have low Δ⁵⁶Fe_ol-melt, indicating that Fe isotope fractionation is also driven by diffusion of isotopically light Fe into olivine, potentially, as Fe–Mg inter-diffusion. The lowest Δ⁵⁶Fe_ol-melt (? 0.40) was observed in a sample from Westerwald (Germany) with abundant magnetite, indicating relatively oxidizing conditions during magma differentiation. This may have enhanced equilibrium Fe isotope fractionation between olivine and melt or fine dispersed magnetite in the basalt matrix may have shifted its Fe isotope composition towards higher δ⁵⁶Fe. The decoupling of Li- and Fe isotope fractionation in cpx is likely due to faster diffusion of Li relative to Fe in cpx, implying that the large investigated cpx phenocryst resided in the magma for only a short period of time which was sufficient for Li- but not for Fe diffusion. The absence of any equilibrium Fe isotope fractionation between the investigated cpx phenocryst and its basaltic host may be related to the similar Fe^3 +/Fe^2 + of cpx and melt. In contrast to cpx, the generally light Fe isotope composition of all investigated olivine separates implies the existence of equilibrium- (in addition to diffusion-driven) isotope fractionation between olivine and melt, on the order of 0.1‰.     

Copper isotopes as monitors of redox processes in hydrothermal mineralization

The stable copper isotope composition of 79 samples of primary and secondary copper minerals from hydrothermal veins in the Schwarzwald mining district, South Germany, shows a wide variation in δ⁶⁵Cu ranging from −2.92 to 2.41‰. We investigated primary chalcopyrite, various kinds of fahlores and emplectite, as well as supergene native copper, malachite, azurite, cuprite, tenorite, olivenite, pseudomalachite and chrysocolla. Fresh primary Cu(I) ores have at most localities copper isotope ratios (δ⁶⁵Cu values) of 0 ± 0.5‰ despite the fact that the samples come from mineralogically different types of deposits covering an area of about 100 by 50 km and that they formed during three different mineralization events spanning the last 300 Ma. Relics of the primary ores in oxidized samples (i.e., chalcopyrite relics in an iron oxide matrix with an outer malachite coating) display low isotope ratios down to −2.92‰. Secondary Cu(I) minerals such as cuprite have high δ⁶⁵Cu values between 0.4 and 1.65‰, whereas secondary Cu(II) minerals such as malachite show a range of values between −1.55 and 2.41‰, but typically have values above +0.5‰. Within single samples, supergene oxidation of fresh chalcopyrite with a δ value of 0‰ causes significant fractionation on the scale of a centimetre between malachite (up to 1.49‰) and relict chalcopyrite (down to −2.92‰). The results show that—with only two notable exceptions—high-temperature hydrothermal processes did not lead to significant and correlatable variations in copper isotope ratios within a large mining district mineralized over a long period of time. Conversely, low-temperature redox processes seriously affect the copper isotope compositions of hydrothermal copper ores. While details of the redox processes are not yet understood, we interpret the range in compositions found in both primary Cu(I) and secondary Cu(II) minerals as a result of two competing controls on the isotope fractionation process: within-fluid control, i.e., the fractionation during the redox process among dissolved species, and fluid-solid control, i.e., fractionation during precipitation involving reactions between dissolved Cu species and minerals. Additionally, Rayleigh fractionation in a closed system may be responsible for some of the spread in isotope compositions. Our study indicates that copper isotope variations may be used to decipher details of natural redox processes and therefore may have some bearing on exploration, evaluation and exploitation of copper deposits. On the other hand, copper isotope analyses of single archeological artefacts or geological or biological objects cannot be easily used as reliable fingerprint for the source of copper, because the variation caused by redox processes within a single deposit is usually much larger than the inter-deposit variation.