Historical variations in the stable isotope composition of mercury in a sediment core from a riverine lake: Effects of dams,pulp and paper mill wastes,and mercury from a chlor-alkali plant

The Wabigoon River (Ontario, Canada) was affected by dams starting in 1898 and was polluted with pulp and paper mill wastes starting in 1913 and mercury from a chlor-alkali plant from 1962 to 1975. A dated sediment core from a riverine lake was analysed to investigate resultant changes in the biogeochemistry of mercury as revealed by variations in mercury isotope ratios and sediment chemistry. A total mercury maximum formed by the mercury pollution coincided with minimums in the δ-values of the ¹⁹⁸Hg/²⁰²Hg, ¹⁹⁹Hg/²⁰²Hg, ²⁰⁰Hg/²⁰²Hg, and ²⁰¹Hg/²⁰²Hg ratios, and the δ-values decreased in the order δ²⁰¹Hg > δ²⁰⁰Hg > δ¹⁹⁹Hg > δ¹⁹⁸Hg. Thus, mass-dependent fractionation caused depletion in lighter isotopes, implying evaporation of Hg(0) and pollution of the atmosphere as well as the river-lake system. Concurrently, mass-independent fractionation caused ¹⁹⁹Hg enrichment, possibly reflecting an independently documented upsurge in methylmercury production, and ²⁰¹Hg depletion, suggesting removal of methylmercury with anomalously high ²⁰¹Hg/¹⁹⁹Hg ratios by aquatic organisms and accumulation of ²⁰¹Hg-depleted inorganic Hg(II) in sediments. The δ²⁰¹Hg/δ¹⁹⁹Hg ratio rose abruptly when mercury pollution began, reflecting the resultant increase in methylmercury production, and remained high but gradually declined as the pollution abated, paralleling trends shown by methylmercury in aquatic organisms. The δ²⁰¹Hg/δ¹⁹⁹Hg ratio of pre-1962 background mercury increased ca. 1898 and ca. 1913–1929, suggesting accelerated methylmercury production due to stimulation of microbial activities by the damming of the river and the input of pulp and paper mill wastes, respectively. Other variations were linked to economic and technological factors that affected pulp and paper manufacture.     

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.     

Liquid-vapor fractionation of boron and boron isotopes: Experimental calibration at 400°C/23 MPa to 450°C/42 MPa

We experimentally determined the boron partitioning and boron isotope fractionation between coexisting liquid and vapor in the system H₂O−NaCl−B₂O₃. Experiments were performed along the 400 and 450°C isotherms. Pressure conditions ranged from 23 to 28 MPa at 400°C and from 38 to 42 MPa at 450°C. Boron partitions preferentially into the liquid. Its overall liquid-vapor fractionation is, however, weak: Calculated boron distribution coefficients D_B^liquid-vapor are < 2.5 at all run conditions. With decreasing pressure (i.e. increasing opening of the solvus) D_B^liquid-vapor increases along the individual isotherms. Extrapolation to salt saturated conditions yields maximum boron liquid-vapor fractionations of D_B^liquid-vapor = 1.8 at 450°C and D_B^liquid-vapor = 2.7 at 400°C. ¹¹B preferentially fractionates into the vapor. Calculated Δ¹¹B_vapor-liquid = {[(¹¹B/¹⁰B)_vapor - (¹¹B/¹⁰B)_liquid]/(¹¹B/¹⁰B)_{NBS 951}}*1000 are small and range from 0.2 (± 0.7) to 0.9 (± 0.5) ‰ at 450°C and from 0.1 (± 0.6) to 0.7 (± 0.6) ‰ at 400°C. The data indicate increasing isotopic fractionation with decreasing pressure (i.e. increasing opening of the solvus). Extrapolation to salt saturated conditions yields maximum boron isotope liquid-vapor fractionations of Δ¹¹B_vapor-liquid = 1.5 (± 0.7) ‰ at 450°C and Δ¹¹B_vapor-liquid = 1.3 (± 0.6) ‰ at 400°C. The weak boron isotope fractionation suggests similar trigonal speciation in liquid and vapor. Although the boron and boron isotope fractionation between liquid and vapor is only weak, mass balance calculations indicate that for high degrees of fractionation liquid-vapor phase separation in an open system can significantly alter the boron and boron isotope signature of low-salinity hydrous fluids in hydrothermal systems. Comparing the model calculations with natural oceanic hydrothermal fluids, however, indicate that other processes than fluid phase separation dominate the boron geochemistry in oceanic hydrothermal fluids.     

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

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

Elemental and isotopic fractionation of Type B CAI-like liquids by evaporation

Vacuum evaporation experiments with Type B CAI-like starting compositions were carried out at temperatures of 1600, 1700, 1800, and 1900 °C to determine the evaporation kinetics and evaporation coefficients of silicon and magnesium as a function of temperature as well as the kinetic isotope fractionation factor for magnesium. The vacuum evaporation kinetics of silicon and magnesium are well characterized by a relation of the form J = J_oe^−E/RT with J_o = 4.17 × 10⁷ mol cm⁻² s⁻¹, E = 576 ± 36 kJ mol⁻¹ for magnesium, J_o = 3.81 × 10⁶ mol cm⁻² s⁻¹, E = 551 ± 63 kJ mol⁻¹ for silicon. These rates only apply to evaporation into vacuum whereas the actual Type B CAIs were almost certainly surrounded by a finite pressure of a hydrogen-dominated gas. A more general formulation for the evaporation kinetics of silicon and magnesium from a Type B CAI-like liquid that applies equally to vacuum and conditions of finite hydrogen pressure involves combining our determinations of the evaporation coefficients for these elements as a function of temperature (γ = γ₀e^−E/RT with γ₀ = 25.3, E = 92 ± 37 kJ mol⁻¹ for γ_Si; γ₀ = 143, E = 121 ± 53 kJ mol⁻¹ for γ_Mg) with a thermodynamic model for the saturation vapor pressures of Mg and SiO over the condensed phase. High-precision determinations of the magnesium isotopic composition of the evaporation residues from samples of different size and different evaporation temperature were made using a multicollector inductively coupled plasma mass spectrometer. The kinetic isotopic fractionation factors derived from this data set show that there is a distinct temperature effect, such that the isotopic fractionation for a given amount of magnesium evaporated is smaller at lower temperature. We did not find any significant change in the isotope fractionation factor related to sample size, which we interpret to mean that recondensation and finite chemical diffusion in the melt did not affect the isotopic fractionations. Extrapolating the magnesium kinetic isotope fractionations factors from the temperature range of our experiments to temperatures corresponding to partially molten Type B CAI compositions (1250-1400 °C) results in a value of α_Mg ≈ 0.991, which is significantly different from the commonly used value of .     

The effect of aqueous diffusion on the fractionation of chlorine and bromine stable isotopes

Diffusive isotopic fractionation factors are important in order to understand natural processes and have practical application in radioactive waste storage and carbon dioxide sequestration. We determined the isotope fractionation factors and the effective diffusion coefficients of chloride and bromide ions during aqueous diffusion in polyacrylamide gel. Diffusion was determined as functions of temperature, time and concentration. The effect of temperature is relatively large on the diffusion coefficient (D) but only small on isotope fractionation. For chlorine, the ratio, D_³⁵Cl/D_³⁷Cl varied from 1.00128 ± 0.00017 (1σ) at 2 °C to 1.00192 ± 0.00015 at 80 °C. For bromine, D_⁷⁹Br/D_⁸¹Br varied from 1.00098 ± 0.00009 at 2 °C to 1.0064 ± 0.00013 at 21 °C and 1.00078 ± 0.00018 (1σ) at 80 °C. There were no significant effects on the isotope fractionation due to concentration. The lack of sensitivity of the diffusive isotope fractionation to anything at the most common temperatures (0 to 30 °C) makes it particularly valuable for application to understanding processes in geological environments and an important natural tracer in order to understand fluid transport processes.     

Solubility of cyclooctasulfur in pure water and sea water at different temperatures

The solubility of cyclooctasulfur in water and sea water at various temperatures in the range between 4 and 80 °C was determined. Cyclooctasulfur in equilibrium with rhombic sulfur reacted with hot acidic aqueous potassium cyanide to form thiocyanate anion which was measured by anion chromatography. Sulfur solubility in pure water was found to increase with temperature by more than 78 times: from 6.1 nM S₈ at 4 °C to 478 nM S₈ at 80 °C. The following thermodynamic values for solubilisation of S₈ in water were calculated from the experimental data: K° = 3.01 ± 1.04 × 10⁻⁸, ΔG_r° = 42.93 ± 0.73 kJ mol⁻¹, ΔH_r° = 47.4 ± 3.6 kJmol⁻¹, ΔS_r° = 15.0 ± 11.7 J mol⁻¹ K⁻¹). Solubility of cyclooctasulfur in sea water was found to be 61 ± 13% of the solubility in pure water regardless of the temperature.     

Quantifying Li isotope fractionation during smectite formation and implications for the Li cycle

Tri-octahedral Li-Mg smectites (hectorites) were synthesized at temperatures ranging from 25 to 250 °C, in the presence of solutions highly enriched in lithium. After removing all the exchangeable lithium from the synthesized clays, Li isotope fractionation (Δ⁷Li_{clay-solution}) was determined. This fractionation was linked to Li incorporation into the structural octahedral site, substituting for Mg²⁺. As predicted, experimental Δ⁷Li_{clay-solution} inversely correlates with temperature, and ranges from −1.6‰ ± 1.3‰ at 250 °C to −10.0‰ ± 1.3‰ at 90 °C, and then stays relatively constant down to 25 °C. The relatively constant isotope fractionation factor below 90 °C may be due to high concentrations of edge octahedra in low crystallinity smectites. The isotopic fractionation factor (α), for a given temperature, does not depend on the solution matrix, nor on the amount of structural Li incorporated into the clay. Empirical linear laws for α as a function of 1/T (K) were inferred. Smectite Li contents and smectite-solution distribution coefficients (D_Li/Mg) increase with temperature, as expected for a substitution process. The fractions of dissolved Li incorporated into the smectite octahedral sites are small and do not depend on the duration of the experiment. In a seawater-like matrix solution, less Li is incorporated into the smectites, probably as a result of competition with dissolved Mg²⁺ ions for incorporation into the octahedral sites. The high Li contents observed in marine smectites are therefore best explained either by a significant contribution from basalts, by adsorption processes, or by the influence of seawater chemical composition on distribution coefficients. We also calculate, using present-day estimates of hydrothermal water and river fluxes, that a steady-state ocean would require a relatively large global clay-water Li isotope fractionation (−12‰ to −21‰). This study demonstrates the ability of laboratory experiments to quantify the impact of secondary phases on the Li geochemical cycle and associated isotope fractionations.     

Sources and biological fractionation of Silicon isotopes in the Eastern Equatorial Pacific

Silicon isotopes in dissolved silicic acid were measured in the upper four kilometers between 4°N and 3°S latitude at 110°W longitude in the eastern Equatorial Pacific. Silicon isotopes became progressively heavier with silicic acid depletion of surface water as expected from biological fractionation. The value of ε estimated by applying a steady-state isotope fractionation model to data from all stations between 4°N and 3°S was −0.77 ± 0.12‰ (std. err.). When the analysis was restricted to those stations whose temperature and salinity profiles indicated that they were directly influenced by upwelling of the Equatorial Undercurrent (EUC), the resulting value of ε was −1.08 ± 0.27‰ (std. err.) similar to the value established in culture studies (−1.1‰). When the non steady state Rayleigh model was applied to the same restricted data set the resulting value of ε was significantly more positive, −0.61 ± 0.16‰ (std. err.). To the extent that the equatorial system approximates a steady state these results support a value of −1.1‰ for the fractionation factor for isotopes of Si in the sea. Without the assumption of steady state the value of ε can only be constrained to be between −0.6 and −1.1‰. Silicic acid in Equatorial Pacific Deep Water below 2000 m had a near constant δ³⁰Si of +1.32 ± 0.05‰. That value is significantly more positive than obtained for North Pacific Deep Water at similar depths at stations to the northwest of our study area (0.9-1.0‰) and it is slightly less positive than new measures of the δ³⁰Si of silicic acid from the silicic acid plume centered over the Cascadia basin in the Northeast Pacific (Si(OH)₄ > 180  μM, δ³⁰Si = +1.46 ± 0.12‰ (SD, n = 4). We show that the data from the equator and Cascadia basin fit a general trend of increasing δ³⁰Si(OH)₄ with increasing silicic acid concentration in the deep sea, but that the isotope values from the Northeast Pacific are anomalously light. The observed level of variation in the silicon isotope composition of deep waters from this single ocean basin is considerably larger than that predicted by current models based on fractionation during opal formation with no isotope effect during dissolution. Confirmation of such high variability in deep water δ³⁰Si(OH)₄ within individual ocean basins will require reassessment of the mechanisms controlling the distribution of isotopes of silicon in the sea.     

Intra-shell boron isotope ratios in the symbiont-bearing benthic foraminiferan Amphistegina lobifera: Implications for δB vital effects and paleo-pH reconstructions

The boron isotope composition of marine carbonates is considered to be a seawater pH proxy. Nevertheless, the use of δ¹¹B has some limitations such as the knowledge of the fractionation factor (α_4-3) between boric acid and the borate ion and the amplitude of “vital effects” on this proxy that are not well constrained. Using secondary ion mass spectrometry (SIMS) we have examined the internal variability of the boron isotope ratio in the shallow water, symbionts bearing foraminiferan Amphistegina lobifera. Specimens were cultured at constant temperature (24 ± 0.1 °C) in seawater with pH ranging between 7.90 and 8.45. Intra-shell boron isotopes showed large variability with an upper limit value of ≈30‰. Our results suggest that the fractionation factor α_4-3 of 0.97352 (Klochko et al., 2006) is in better agreement with our experiments and with direct pH measurements in seawater vacuoles associated with the biomineralization process in these foraminifera. Despite the large variability of the skeletal pH values in each cultured specimen, it is possible to link the lowest calculated pH values to the experimental culture pH values while the upper pH limit is slightly below 9. This variability can be interpreted as follows: foraminifera variably increase the pH at the biomineralization site to about 9. This increase above ambient seawater pH leads to a range in δ¹¹B (Δ¹¹B) for each seawater pH. This Δ¹¹B is linearly correlated with the culture seawater pH with a slope of −13.1 per pH unit, and is independent of the fractionation factor α_4-3, or the δ¹¹B_sw through time. It may also be independent of the pK_B (the dissociation constant of boric acid) value. Therefore, Δ¹¹B in foraminifera can potentially reconstruct paleo-pH of seawater.     

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).     

High-temperature Si isotope fractionation between iron metal and silicate

The Si stable isotope fractionation between metal and silicate has been investigated experimentally at 1800, 2000, and 2200 °C. We find that there is a significant silicon stable isotope fractionation at high temperature between metal and silicate in agreement with Shahar et al. (2009). Further we find that this fractionation is insensitive to the structure and composition of the silicate as the fractionation between silicate melt and olivine is insignificant within the error of the analyses. The temperature-dependent silicon isotope fractionation is Δ³⁰Si_{silicate-metal} = 7.45 ± 0.41 × 10⁶/T². We also demonstrate the viability of using laser ablation MC-ICPMS as a tool for measuring silicon isotope ratios in high pressure and temperature experiments.     

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.     

Silicon isotopic fractionation of CAI-like vacuum evaporation residues

Calcium-, aluminum-rich inclusions (CAIs) are often enriched in the heavy isotopes of magnesium and silicon relative to bulk solar system materials. It is likely that these isotopic enrichments resulted from evaporative mass loss of magnesium and silicon from early solar system condensates while they were molten during one or more high-temperature reheating events. Quantitative interpretation of these enrichments requires laboratory determinations of the evaporation kinetics and associated isotopic fractionation effects for these elements. The experimental data for the kinetics of evaporation of magnesium and silicon and the evaporative isotopic fractionation of magnesium is reasonably complete for Type B CAI liquids (Richter F. M., Davis A. M., Ebel D. S., and Hashimoto A. (2002) Elemental and isotopic fractionation of Type B CAIs: experiments, theoretical considerations, and constraints on their thermal evolution. Geochim. Cosmochim. Acta66, 521-540; Richter F. M., Janney P. E., Mendybaev R. A., Davis A. M., and Wadhwa M. (2007a) Elemental and isotopic fractionation of Type B CAI-like liquids by evaporation. Geochim. Cosmochim. Acta71, 5544-5564.). However, the isotopic fractionation factor for silicon evaporating from such liquids has not been as extensively studied. Here we report new ion microprobe silicon isotopic measurements of residual glass from partial evaporation of Type B CAI liquids into vacuum. The silicon isotopic fractionation is reported as a kinetic fractionation factor, α_Si, corresponding to the ratio of the silicon isotopic composition of the evaporation flux to that of the residual silicate liquid. For CAI-like melts, we find that α_Si = 0.98985 ± 0.00044 (2σ) for ²⁹Si/²⁸Si with no resolvable variation with temperature over the temperature range of the experiments, 1600-1900 °C. This value is different from what has been reported for evaporation of liquid Mg₂SiO₄ (Davis A. M., Hashimoto A., Clayton R. N., and Mayeda T. K. (1990) Isotope mass fractionation during evaporation of Mg₂SiO₄. Nature347, 655-658.) and of a melt with CI chondritic proportions of the major elements (Wang J., Davis A. M., Clayton R. N., Mayeda T. K., and Hashimoto A. (2001) Chemical and isotopic fractionation during the evaporation of the FeO-MgO-SiO₂-CaO-Al₂O₃-TiO₂-REE melt system. Geochim. Cosmochim. Acta65, 479-494.). There appears to be some compositional control on α_Si, whereas no compositional effects have been reported for α_Mg. We use the values of α_Si and α_Mg, to calculate the chemical compositions of the unevaporated precursors of a number of isotopically fractionated CAIs from CV chondrites whose chemical compositions and magnesium and silicon isotopic compositions have been previously measured.     

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

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

Experimental tests of the effects of Al substitution on the goethite-water D/H fractionation factor

Twelve goethite samples with different degrees of substitution of Al for Fe were synthesized at 22-48 °C and pH values of 1.5-14 under closed system conditions and used to study the effects of Al substitution on the hydrogen isotopic fractionation between goethite and its ambient water. The syntheses followed two pathways: (1) Fe³⁺ hydrolysis in high pH aqueous solutions; (2) oxidation of Fe²⁺ to Fe³⁺ in mid to low pH solutions. XRD and SEM analyses indicated that, irrespective of temperature and pH, goethite was the predominant product of the syntheses in all of the experiments (with degrees of Al substitution as high as ∼13 mol %). “High temperature nonstoichiometric” (HTN) water is present in all of the samples and rapidly exchanges D/H with ambient vapor at room temperature. Uncertainties in the value of the apparent D/H fractionation factor (α_e-v) between HTN water and ambient exchange water at 22 °C lead to significant uncertainties in determinations of the δD values of structural hydrogen (δD_s) in goethites which contain high proportions of HTN water. As determined for the samples of this study, α_e-v has a nominal value of 0.942 (±0.02). δD_s values determined using an α_e-v value of 0.942 indicate that Al substitution increases the δD value of structural hydrogen in goethite by about 1.4 (±0.4)‰ for each increase in Al of 1 mol %. This dependence on Al is of the same sign as, but somewhat larger in magnitude than, the effect of Al predicted by a published model (∼0.7‰ per mol % Al). The overall uncertainties in the current results suggest that an increase of ∼1‰ per mol % Al, as adopted by previous studies, may be a reasonable estimate with which to adjust δ D_s values of natural goethites to those of the pure FeOOH endmember and could be valid for degrees of Al substitution of up to at least 15 mol %. These synthesis experiments also yield a hydrogen isotopic fractionation factor (^Dα_G-W) between pure goethite (α-FeOOH) and liquid water of 0.900 (±0.006), which is analytically indistinguishable from the published value of 0.905 (±0.004). Thus, use of an ^Dα_G-W value of 0.905 in applications to the FeOOH component of natural goethites is supported by the current study.     

The geochemical behavior and isotopic composition of Hg in a mid-Pleistocene western Mediterranean sapropel

The concentrations of mercury (Hg) and other trace metals (Ni, Cu, Zn, Mo, Ba, Re, U) and the Hg isotopic composition were examined across a dramatic redox and productivity transition in a mid-Pleistocene Mediterranean Sea sapropel sequence. Characteristic trace metal enrichment in organic-rich layers was observed, with organic-rich sapropel layers ranging in Hg concentration from 314 to 488 ng/g (avg = 385), with an average enrichment in Hg by a factor of 5.9 compared to organic-poor background sediments, which range from 39 to 94 ng/g Hg (avg = 66). Comparison of seawater concentrations and sapropel accumulations of trace metals suggests that organic matter quantitatively delivers Hg to the seafloor. Near complete scavenging of Hg from the water column renders the sapropel Hg isotopic composition representative of mid-Pleistocene Mediterranean seawater. Sapropels have an average δ²⁰²Hg value of −0.91‰ ± 0.15‰ (n = 5, 1 SD) and Δ¹⁹⁹Hg value of 0.11‰ ± 0.03‰ (n = 5, 1 SD). Background sediments have an average δ²⁰²Hg of −0.76‰ ± 0.16‰ (n = 5, 1 SD) and Δ¹⁹⁹Hg of 0.05‰ ± 0.01‰ (n = 5, 1 SD), which is indistinguishable from the sapropel values. We suggest that the sapropel isotopic composition is most representative of the mid-Pleistocene Tyrrhenian Sea.     

Exchange and fractionation of Mg isotopes between epsomite and saturated MgSO4 solution

The equilibrium Mg isotope fractionation factor between epsomite and aqueous MgSO₄ solution has been measured using the three isotope method in recrystallization experiments conducted at 7, 20, and 40 °C. Complete or near-complete isotopic exchange was achieved within 14 days in all experiments. The Mg isotope exchange rate between epsomite and MgSO₄ solution is dependent on the temperature, epsomite seed crystal grain size, and experimental agitation method. The Mg isotope fractionation factors (Δ²⁶Mg_eps-sol) at 7, 20, and 40 °C are 0.63 ± 0.07‰, 0.58 ± 0.16‰, and 0.56 ± 0.03‰, respectively. These values are indistinguishable within error, indicating that the Mg isotope composition of epsomite is relatively insensitive to temperature. The magnitude of the isotope fractionation factor (Δ²⁶Mg_eps-sol = ca. 0.6‰ between 7 and 40 °C) indicates that significant Mg isotope variations can be produced in evaporite sequences, and Mg isotopes may therefore, constrain the degree of closed-system behavior, paleo-humidity, and hydrological history of evaporative environments.     

Oxygen isotope fractionation between synthetic aragonite and water: Influence of temperature and Mg concentration

Aragonite was precipitated in the laboratory at 0, 5, 10, 25, and 40 °C to determine the temperature dependence of the equilibrium oxygen isotope fractionation between aragonite and water. Forced CO₂ degassing, passive CO₂ degassing, and constant addition methods were employed to precipitate aragonite from supersaturated solutions, but the resulting aragonite-water oxygen isotope fractionation was independent of the precipitation method. In addition, under the experimental conditions of this study, the effect of precipitation rate on the oxygen isotope fractionation between aragonite and water was almost within the analytical error of ±∼0.13‰ and thus insignificant. Because the presence of Mg²⁺ ions is required to nucleate and precipitate aragonite from Na-Ca-Cl-HCO₃ solutions under these experimental conditions, the influence of the total Mg²⁺ concentration (up to ∼0.9 molal) on the aragonite-water oxygen isotope fractionation was examined at 25 °C. No significant Mg²⁺ ion effect, or oxygen isotope salt effect, was detected up to 100 mmolal total Mg²⁺ but a noticeable isotope salt effect was observed at ∼0.9 molal total Mg²⁺.On the basis of results of the laboratory synthesis experiments, a new expression for the aragonite-water fractionation is proposed over the temperature range of 0-40 °C:

1000lnα_{aragonite-water}=17.88±0.13(10³/T)-31.14±0.46     

An experimental study of magnesium-isotope fractionation in chlorophyll-a photosynthesis

Measurements are presented of the magnesium isotopic composition of chlorophyll-a, extracted from cyanobacteria, relative to the isotopic composition of the culture medium in which the cyanobacteria were grown. Yields of 50-93% chlorophyll-a were achieved from the pigment extracts of Synechococcus elongatus, a unicellular cyanobacteria. This material was then digested using concentrated nitric acid to extract magnesium. Separation was accomplished using columns of cation-exchange resin, which achieved a 103 ± 10% yield of magnesium from chlorophyll-a. This procedure ensured accurate measurement of the magnesium-isotopic ratios without isobaric interferences using a multi-collector inductively coupled plasma mass spectrometry (MC-ICP-MS). We find a slight depletion in the heavier isotopes of magnesium in chlorophyll-a relative to culture medium, early growth phase: Δ²⁶Mg = −0.71(±0.35)‰ and Δ²⁵Mg = −0.37(±0.18)‰; late growth phase: Δ²⁶Mg = −0.53(±0.20)‰ and Δ²⁵Mg = −0.26(±0.11)‰, due to an apparent mass-dependent fractionation. We suggest that the small fractionation results from chelation during intracellular processes. A likely candidate for this chelation step involves the magnesium-chelatase enzyme, which mediates the insertion of magnesium to the tetrapyrrole ring during chlorophyll-a biosynthesis. Proof of this hypothesis can be tested with biological controls whereby steps in the enzymatic pathways of chlorophyll synthesis are selectively suppressed.