Oxygen Isotope Fractionation Factors Research Articles

Composition effects on equilibrium oxygen isotope fractionation factors between garnets and quartz/calcite/olivine: Implication from vibration frequencies of garnets at high temperatures and high pressures

Garnet-group minerals, with their wide range of compositions, play a significant role in Earth's crust and upper mantle, participating in various petrological and geochemical processes. Oxygen isotope fractionation factors between garnets and other minerals hold crucial implications in these contexts. In this work, vibration frequencies were measured via Raman spectra on five garnet mineral samples in the pyrope-almandine-spessartine ternary system and four synthetic pyrope-grossular solid solutions at temperatures up to 1000 °C and pressures up to 17 GPa. Isobaric (γiP) and isothermal (γiT) mode Grüneisen, as well as anharmonic (ai) parameters, were determined for the observed modes. Our study shows that the vibration bands tend to shift to lower frequencies with increasing temperature and to higher frequencies with increasing pressure. Moreover, the anharmonic correction has been found to contribute positively to equilibrium oxygen isotope fractionation β factors in garnets, typically within 0.5 ‰ above 1000 K. The 103∙lnβ(T) factors calculated from vibration spectra (considering Г points only) align closely with those from theoretical calculations (including all frequencies in the Brillouin zone). Contributions from phonon dispersion (including none-Г points) are typically smaller than the uncertainties propagated from frequency measurements. Additionally, the pressure effect on the oxygen isotope fractionation in garnet, evaluated from the isothermal Grüneisen parameters, is found to be insignificant under crustal and upper mantle conditions, consistent with thermodynamic expectations. Using published oxygen isotope fractionation β factors for quartz and calcite, the equilibrium oxygen isotope fractionation factors (103∙lnα) are calculated between garnets of diverse compositions and any of these phases as functions of temperature, pressure and fraction of garnet endmembers. Considering composition effects helps to reduce the discrepancies among experimental and empirical calibrations of oxygen isotope fractionation factors between garnet and quartz. As an example application, oxygen isotope fractionation factors 103∙lnα(P, T) between garnets and forsterite were calculated under upper mantle conditions. Our findings suggest that these fractionation factors are predominantly dependent on temperature and garnet compositions, with minor influence from pressure. Compared with our calculated fractionation factors between garnet and olivine, the oxygen isotope exchange equilibrium may have been reached in some in kimberlite and mantle xenolith samples between garnet and olivine. However, in other samples, the fractionation factors cannot be explained solely by the composition effect, likely due to metasomatism. Our results underscore the importance of considering garnet compositions when interpreting their oxygen isotope compositions.

Oxygen Isotope Fractionation between Carbonate Minerals and Carbonic Acid Systems and Constraints for Environmental Science and Geological Processes

The equilibrium oxygen isotope fractionation factor is widely used in geological thermometry. However, under most natural conditions, the oxygen isotope exchange is rare to reach equilibrium. Especially for the complex water–rock interaction process, the contribution of the H2CO3 solution, CO32− solution, Ca(HCO3)2 solution, and CaCO3 solution to the equilibrium oxygen isotope fractionation factor of this process is poorly understood. In view of this predicament, these key parameters are obtained by ab initio calculations. The results showed that the contributions of different carbonate minerals and different aqueous solutions to the equilibrium oxygen isotope fractionation factor were different. Among all nine carbonate minerals (dolomite, calcite, aragonite, magnesite, siderite, otavite, smithsonite, ankerite, and strontianite), the minerals with the highest and lowest reduced partition function ratios (RPFR) were siderite and strontianite, respectively. At the same time, the RPFR of nitratine, which has the same structure as carbonate, was studied. The RPFRs of the three most widely distributed carbonates in nature (dolomite, calcite, and aragonite) were dolomite > calcite > aragonite. Among the H2CO3 solution, CO32− solution, Ca(HCO3)2 solution, and CaCO3 solution, the H2CO3 solution had the strongest ability to enrich 18O. In addition, the equilibrium oxygen isotope fractionation factors between aqueous solutions and gas phase species (CO2(g), H2O(g), and O2(g), etc.) were calculated systematically. The results showed that the oxygen isotope fractionation factors between solutions and gas phases were often inconsistent with the temperature change direction and that the kinetic effects played a key role. These theoretical parameters obtained in this study will provide key equilibrium oxygen isotope constraints for water-rock interaction processes.

Theoretical estimates of sulfoxyanion triple-oxygen equilibrium isotope effects and their implications

Triple-oxygen isotope (δ18O and Δ′17O) analysis of sulfate is becoming a common tool to assess several biotic and abiotic sulfur-cycle processes, both today and in the geologic past. Multi-step sulfur redox reactions often involve intermediate sulfoxyanions such as sulfite, sulfoxylate, and thiosulfate, which may rapidly exchange oxygen atoms with surrounding water. Process-based reconstructions therefore require knowledge of equilibrium oxygen-isotope fractionation factors (18α and 17α) between water and each individual sulfoxyanion. Despite this importance, there currently exist only limited experimental 18α data and no 17α estimates due to the difficulty of isolating and analyzing short-lived intermediate species. To address this, we theoretically estimate 18α and 17α for a suite of sulfoxyanions—including several sulfate, sulfite, sulfoxylate, and thiosulfate species—using quantum computational chemistry. We determine fractionation factors for sulfoxyanion “water droplets” using the B3LYP/6-31G+(d,p) method; we additionally calculate higher-order method (CCSD/aug-cc-pVTZ and MP2/aug-cc-pVTZ) scaling factors, and we qualitatively estimate the importance of anharmonic zero-point energy (ZPE) corrections using a suite of gaseous sulfoxy compounds. Methodological scaling factors greatly impact 18α predictions, whereas ZPE corrections are likely small (i.e., ⩽1‰) at Earth-surface temperatures; existing experimental data best agree with 18α predictions when including redox state-specific CCSD/aug-cc-pVTZ scaling factors. Theoretical pH- and temperature-specific bulk-solution (i.e., abundance-weighted average of all species) 18α values yield root-mean-square errors for sulfate/water, sulfite/water, and thiosulfate/water equilibrium of 4.5‰ (n=18 experimental conditions), 3.7‰ (n=27), and 2.2‰ (n=3), respectively. However, sulfate- and sulfite-system agreement improves considerably when comparing experimental results only to SO3(OH)−/H2O (RMSE = 1.6‰) and SO2(OH)−/H2O (RMSE = 2.2‰) predictions, rather than bulk solutions. This is particularly true for the sulfite system at high and low pH, when SO2(OH)− is not the dominant species. We discuss potential experimental and theoretical biases that may lead to this apparent improvement. By combining 18α and 17α predictions, we additionally estimate that sulfate, sulfite, sulfoxylate, and thiosulfate species can exhibit Δ′17O values as much as 0.199‰, 0.205‰, 0.101‰, and 0.186‰ more negative than equilibrated water at Earth-surface temperatures (reference line slope = 0.5305). This theoretical framework provides a foundation to interpret experimental and observational triple-oxygen isotope results of several sulfur-cycle processes including pyrite oxidation, microbial metabolisms (e.g., sulfate reduction, thiosulfate disproportionation), and hydrothermal anhydrite precipitation. We highlight this with several examples.

Simultaneous Calculation of Chemical and Isotope Equilibria Using the GEOCHEQ_Isotope Software: Oxygen Isotopes

—The GEOCHEQ_Isotope software package, elaborated previously for modeling chemical and carbon isotope equilibria in hydrothermal and hydrogeochemical systems by minimizing the Gibbs energy, is extended to the simultaneous calculation of carbon and oxygen isotopic effects. Similar to what was done for carbon, the β-factor formalism was used to develop algorithms and a database for calculating the isotopic effects of oxygen. According to the developed algorithm, the Gibbs energy of formation of a rare isotopologue, G*(P, T), is calculated through the Gibbs energy of formation of the main isotopologue, the value of the β18O factor of this substance, and the mass ratio of the rare (18O) and main (16O) isotopes. The isotope mixture is assumed to be ideal. The temperature dependence of the β-factor is unified as a polynomial in reciprocal absolute temperature. Necessary information on oxygen isotope equilibria involving important geochemical compounds was critically analyzed, and the available data were reconciled and modified. The temperature dependences of the β18O-factors were correspondingly optimized. The thermodynamic database was updated by adding information on the temperature dependence of β18O-factors specified by polynomial coefficients for each substance. The usage of the GEOCHEQ_Isotope software package and the corresponding database is demonstrated by modeling the dependence of oxygen and carbon isotope fractionation factors on the acidity of the solution (pH) in a carbonate hydrothermal system. The simulation results are in a good agreement with experimental data available from the literature. The enrichment of dissolved carbonates in the 18O heavy oxygen isotope relative to water decreases with increasing pH of the system. At the same time, a pH increase results in a decrease in the negative carbon isotope shift between calcite and dissolved carbonates. At high pH values (~11), the isotope shift inversion and the enrichment of the dissolved carbonate in the heavy carbon isotope relative to calcite are predicted.

Geochemical comparison of ikaite and methane-derived authigenic carbonates recovered from Echigo Bank in the Sea of Japan

This study compares the geochemistry of ikaite (CaCO3·6H2O) with that of nearby (< 500 m in distance) recovered methane-derived authigenic carbonates (MDACs) of ~10 mol% Mg-calcite. The carbon isotopic composition (δ13C) of MDACs (lower than −49‰ Vienna Pee Dee Belemnite or VPDB) suggests a more prominent influence of anaerobic oxidation of methane (AOM) than that of ikaite (approximately −35‰ VPDB). The higher δ13C value of ikaite is attributed to its growth at sediment depths closer to the seafloor, where the proportion of seawater HCO3− relative to methane-derived HCO3− was higher. Assuming that seawater and AOM are the major sources of HCO3− for ikaite growth, the actual 14C age was estimated to be ~14 cal ky before present, close to the age of the sediments observed around the ikaite. For alkalinity increase at shallower sediment depths by the AOM, a higher methane flux is necessary. However, attributing the ikaite growth at shallower sediment depths to higher methane flux alone posed an issue, as no traces of gas hydrate growth were found in the laminated sediments. The oxygen isotopic composition (δ18O) of the freeze-dried ikaite (ikaite-CO3) was ~2‐3‰ VPDB, which is lower than the corresponding value of ~4.0‰ VPDB for the Mg-calibrated MDACs. It is characteristic of ikaite-CO3 to exhibit such low δ18O values, indicating a unique oxygen isotope fractionation factor for ikaite. The preferential H218O uptake of ikaite for growth also influenced the δ18O value of ikaite-CO3. Hence, owing to H218O depletion in the ambient interstitial water, the δ18O value of the ikaite-CO3 tended to decrease from the interior to the margin. The δ18O values of the ikaite-CO3 were archived in yellowish calcite grains that were formed in the pores after patchy dissociation of the ikaite. The δ18O values of the grains were close to those of ikaite-CO3. Therefore, the δ18O-temperature equation for ikaite is necessary for geochemical interpretation of the calcite δ18O values.

Are oxygen isotope fractionation factors between calcite and water derived from speleothems systematically biased due to prior calcite precipitation (PCP)?

The equilibrium oxygen isotope fractionation factor between calcite and water (18αcalcite/H2O) is an important quantity in stable isotope geochemistry and allows in principle to infer temperature variations from carbonate δ18O if carbonate formation occurred in thermodynamic equilibrium. For this reason, many studies intended to determine the value of the oxygen isotope fractionation factor between calcite and water (18αcalcite/H2O) for a wide range of temperatures using modern cave calcite and the corresponding cave drip water or ancient speleothem carbonate and fluid inclusion samples. However, the picture that emerges from all of these studies indicates that speleothem calcite is not formed in thermodynamic equilibrium but under kinetic conditions, provoking a large variability of determined 18αcalcite/H2O values. Here we present a conceptual framework that can explain the variability of 18αcalcite/H2O values obtained by cave studies. Prior calcite precipitation (PCP) is calcite precipitation before cave drip water is dripping from the cave ceiling and impinges on the surface of a stalagmite or watch glass. Prior to the karst water dripping from the cave ceiling, PCP can occur in the karst above the cave as well as on the cave ceiling, the cave walls and on the surface of stalactites. We argue that PCP leads to increasing the δ18O value of the dissolved HCO3− (δ18OHCO3-), resulting in an oxygen isotope disequilibrium of the δ18OHCO3- values with respect to the δ18O value of water (δ18OH2O). The oxygen isotope disequilibrium between HCO3− and H2O is re-equilibrated by oxygen isotope exchange between H2O and HCO3. Depending on the temperature, the re-equilibration time varies from hours to days and is usually much longer than the residence time of the drip water on stalactites, but much shorter than the time required to percolate through the karst. Therefore, while the oxygen isotope equilibrium between HCO3− and H2O is very likely re-established when PCP occurred in the karst, oxygen isotope disequilibrium conditions between HCO3− and H2O still prevail when PCP occurred inside a cave, e.g., on stalactites. If the oxygen isotope disequilibrium conditions between HCO3− and H2O is not re-established, the precipitated calcite will inherit the elevated δ18O value of the HCO3− and not be in oxygen isotope equilibrium with the corresponding drip water. Consequently, if the 18αcalcite/H2O value is calculated from cave calcite samples affected by PCP, the derived value will be systematically biased.

Chlorine isotope data of chlorides challenge the pore fluid paradigm

In order to examine the seawater-seafloor sediment interactions that influence the chemical composition of seawater through time, we examined hundreds of pore fluid geochemical analyses from 13 clay-rich sedimentary successions drilled by the ODP-IODP. Chemical trends such as monotonous increases in Ca2+, and decreases in Mg2+ and δ18O with depth are traditionally interpreted to result from water-rock interaction. In this view, the release of Ca2+ into fluids and the uptake of Mg2+ and 18O mainly results from the formation of low-temperature clays in the sediment and within underlying basalts. Chloride concentration profiles and isotopic compositions, however, suggest that different processes may influence pore water geochemistry. The data examined here show relatively constant chloride contents but with a systematic decrease in δ37Cl of chlorides with depth from 0 permil (the seawater value) down to −8.5 permil. The δ37Cl data are highly correlated with δ18O (with δ18O down to −5.7 permil).The δ37Cl-depletions of pore fluid chlorides are found in all studied sedimentary piles regardless of tectonic or sedimentary history. These trends cannot be explained by water–rock exchange reactions because minerals formed at low temperature have Cl contents that are too low to compensate for δ37Cl depletions observed in pore fluids. Accordingly, we hypothesize that fluid-specific processes are responsible for the δ37Cl-depletions of the fluids and that δ37Cl-enriched chlorides were expelled out of the sediments into the ocean. After reviewing the fluid-specific processes that are known to change the chlorine isotope ratios in chlorides, we rule out diffusion and gravitational isotope fractionations of chlorides could generate this isotope pattern. The flow of a δ37Cl -depleted fluid from the underlying basaltic basement into the sediments could explain the δ37Cl data. But the mechanism that produces depletion in δ37Cl of the fluid remains unknown. It cannot be chloride exchanges between fluids and rocks.Here we show that compaction-induced ion filtration of chlorides through clay-rich membranes can produce the observed pore fluid δ37Cl-depletions, with isotope fractionation factors ranging from 1.000 to 1.008 between the chlorides of the expelled fluid (the permeate) and those of the residual fluid (the retentate). We find that smectite-rich sediments are associated with higher isotopic fractionation factors, while illite/chlorite-rich sediments are associated with intermediate values and with clay-poor sediments associated with lower values. This suggests that chlorine isotope fractionation might be controlled by surface charge associated with specific clay minerals. Our calculations show that compaction-induced filtration has the capacity to produce 18O-depletion for oxygen isotope fractionation factors between the expelled fluid and the retentate ranging from 1.000 to 1.005. 18O-enrichment in the expelled fluid is in agreement with the experimental data of Haydon and Graf (1986). Overall, although further experimental work on both chlorine and oxygen isotopes is certainly needed, the results of this study indicate that ion-filtration should be considered as a potential mechanism for fractionating isotopic species in sediment pore waters, particularly for oxygen isotope ratios whose variations are often commonly attributed to water–rock exchange.

Oxygen isotope fractionation between water and the aqueous hydroxide ion

The stable oxygen isotope fractionation factor between water and the aqueous hydroxide ion is a fundamental geochemical parameter widely used in the Earth Sciences and other fields. Studies from the 1960s reported α(H2O(l)-OH-(aq)), the fractionation factor between liquid H2O and aqueous OH−, theoretically as 1.046 and ∼1.048 (46‰ and 48‰ at 15 °C and 25 °C) and experimentally as 1.045 (45‰) at 15 °C. These, or similar values have been used in the literature for almost 60 years. Here I present quantum-chemical (QC) calculations, showing that the above theoretical values are fundamentally wrong as they pertain to free OH− (incorrectly assumed equivalent to aqueous OH−) and ignore intermolecular vibrational modes in solution. As a result, the theoretical values from the 1960s are off by a factor of ∼2 (when expressed in ‰), suggesting that the experimental value is also wrong. QC computations of OH−-water clusters with up to n=22 water molecules demonstrate that hydrogen bonding in solution significantly affects the fundamental vibrational modes associated with OH− and substantially reduces the oxygen isotope fractionation between water and OH-(aq), compared to the fractionation between water and free OH−. The most accurate QC methods tested here yield values for the fractionation factor between H2O and OH− in water clusters with n=7 to 22 water molecules of ∼1.019 to ∼1.024 (19 to 24‰) at 25 °C based on the harmonic approximation. Estimated effects due to anharmonicity (from numerically demanding computations) could add uncertainties of up to ∼3‰ to these values.

Oxygen isotope fractionation between gypsum and its formation waters: Implications for past chemistry of the Kawah Ijen volcanic lake, Indonesia

Abstract Gypsum (CaSO4·2H2O) provides an opportunity to obtain information from both the oxygen isotopic composition of the water and sulfate of its formation waters, where these components are commonly sourced from different reservoirs (e.g., meteoric vs. magmatic). Here, we present δ18O values for gypsum and parent spring waters fed by the Kawah Ijen crater lake in East Java, Indonesia, and from these natural samples derive gypsum-fluid oxygen isotope fractionation factors for water and sulfate group ions of 1.0027 ± 0.0003‰ and 0.999 ± 0.001‰, respectively. Applying these fractionation factors to a growth-zoned gypsum stalactite that records formation waters from 1980 to 2008 during a period of passive degassing, and gypsum cement extracted from the 1817 eruption tephra fall deposit, shows that these fluids were in water-sulfate oxygen isotopic equilibrium. However, the 1817 fluid was &amp;gt;5‰ lighter. This indicates that the 1817 pre-eruption lake was markedly different, and had either persisted for a much shorter duration or was more directly connected to the underlying magmatic-hydrothermal system. This exploratory study highlights the potential of gypsum to provide a historical record of both the δ18Owater and δ18Osulfate of its parental waters, and provides insights into the processes acting on volcanic crater lakes or any other environment that precipitates gypsum.

Nitrogen and Oxygen Isotope Signatures of Nitrogen Compounds during Anammox in the Laboratory and a Wastewater Treatment Plant.

Isotopic fractionation factors against 15N and 18O during anammox (anaerobic ammonia oxidization by nitrite) are critical for evaluating the importance of this process in natural environments. We performed batch incubation experiments with an anammox-dominated biomass to investigate nitrogen (N) and oxygen (O) isotopic fractionation factors during anammox and also examined apparent isotope fractionation factors during anammox in an actual wastewater treatment plant. We conducted one incubation experiment with high δ18O of water to investigate the effects of water δ18O. The N isotopic fractionation factors estimated from incubation experiments and the wastewater treatment plant were similar to previous values. We also found that the N isotopic effect (15εNXR of –77.8 to –65.9‰ and 15ΔNXR of –31.3 to –30.4‰) and possibly O isotopic effect (18εNXR of –20.6‰) for anaerobic nitrite oxidation to nitrate were inverse. We applied the estimated isotopic fractionation factors to the ordinary differential equation model to clarify whether anammox induces deviations in the δ18O vs δ15N of nitrate from a linear trajectory of 1, similar to heterotrophic denitrification. Although this deviation has been attributed to nitrite oxidation, the O isotopic fractionation factor for anammox is crucial for obtaining a more detailed understanding of the mechanisms controlling this deviation. In our model, anammox induced the trajectory of the δ18O vs δ15N of nitrate during denitrification to less than one, which strongly indicates that this deviation is evidence of nitrite oxidation by anammox under denitrifying conditions.

Respiration in rivers fractionates stable isotopes of dissolved oxygen; a global investigation on the influences of temperature and flow

Quantifying ecosystem respiration remains challenging in aquatic ecosystems. Most investigators assume that nighttime and daytime respiration are equal. Recent studies suggest measuring dissolved oxygen isotopes during periods with and without photosynthesis can account for variations in daytime and nighttime respiration. These models are extremely sensitive to the oxygen isotopic fractionation factor (α) value used for respiration, yet almost nothing is known about the variability of α and factors driving that variability. We quantified how α varies with temperature and flow velocity using field measurements, laboratory experiments, and a modeling approach. We measured α in the field using sealed recirculating chambers in 16 rivers from different biomes (temperate, tropical, and sub-arctic) to assess a range of possible α values. The α values were widely variable, and variation was higher among sites in the same biome or ecoregion (e.g. 0.9780 ± 0.005 to 0.9898 ± 0.002 among six desert sites) than across different biomes. Our data revealed that both temperature, flow, and biofilm characteristics produced variations in α, with temperature decreasing and flow increasing it, until leveling off at high flow velocities. Biological and physical processes occurring in the diffusion boundary layer produced variations in α. Our results highlight that environmental conditions produce variable α values, the need for site-specific α measurements, and practical implications for consideration when measuring α in the field. More generally we illustrate an array of factors that can influence isotopic fractionation associated with metabolic activity of biologically active layers that could be important in any diffusion-limited environment.

A novel application of triple oxygen isotope ratios of speleothems

We present triple oxygen isotope data from speleothems obtained by an O2-CO2 Pt-catalyzed oxygen-isotope equilibration method. The high precision (9 per meg or better, 1σ SD) of our new speleothem Δ17O (carbonate 17O anomaly) data is sufficient to resolve subtle hydroclimatic signals. In addition, we determined triple oxygen isotope fractionation factors through two sets of modern paired carbonate-dripwater samples collected at a temperature of 17 ± 1 °C, which is prerequisite to the calculation of triple oxygen isotope compositions of parent meteoric waters at cave sites from speleothem proxy data. Based on this calibration, we back calculated triple oxygen isotope compositions of parent waters across well-characterized climate transitions using speleothem proxy data from three regions. Resulting oxygen-isotope data closely track the Global Meteoric Water Line (GMWL), providing a preliminary validation of the method. Our speleothem Δ17O data indicate a 21 per meg difference between Marine Isotope Stage 5d and 5e in samples from Central Asia and a 15 per meg difference between the Middle Holocene and Last Glacial Maximum in samples from the eastern Mediterranean, suggesting a shift in moisture source and/or fractionation history. Unexpectedly, there were no measurable Δ17O differences between glacial and interglacial samples from both the South American (western Amazon) and Asian (southern China) monsoon domains, implying consistent moisture-source conditions across glacial and interglacial cycles, at least in terms of relative humidity. Remarkably, Δ17O values from the western Amazonian samples are significantly higher (∼20 per meg) than those from Asian monsoon regions, suggesting lower relative humidity along moisture trajectories in the western Amazon during the Middle Holocene and Last Glacial period. Similarly, Δ17O values of the eastern Mediterranean samples are significantly higher (19–55 per meg) than those of coeval samples from Central Asia, implying different hydrological environments or moisture sources despite being in the same westerly circulation domain. Speleothem Δ17O data may thus provide new and important constraints for understanding regional and global hydroclimate dynamics.

Simulating speleothem growth in the laboratory: Determination of the stable isotope fractionation (δ13C and δ18O) between H2O, DIC and CaCO3

Here we present novel cave-analogue experiments directly investigating stable carbon and oxygen isotope fractionation between the major involved species of the carbonate system (HCO3−, CO2, CaCO3 and H2O). In these experiments, which were performed under controlled conditions inside a climate box, a thin film of solution flew down an inclined marble or glass plate. After different distances of flow and, thus, residence times on the plate, pH, electrical conductivity, supersaturation with respect to calcite, precipitation rate as well as the δ18O and δ13C values of the dissolved inorganic carbon (DIC) and the precipitated CaCO3 were obtained.Progressive precipitation of CaCO3 along the plate is accompanied by degassing of CO2 and stable isotope fractionation, and the system is driven out of isotope equilibrium. We observe a strong enrichment of the δ13C values with increasing residence time and a smaller enrichment in δ18O. The temporal evolution of the δ18O and δ13C values of both the DIC and the precipitated CaCO3 can be explained by a Rayleigh fractionation model, but the observed enrichment in δ13C values is much larger than expected based on isotope equilibrium fractionation factors.Our setup enables to determine the fractionation between CaCO3 and HCO3−, i.e., εCaCO3/HCO3−. Carbon isotope fractionation, 13εCaCO3/HCO3−, is strongly negative for all experiments and much lower than equilibrium isotope fractionation (0–1‰). In addition, 13εCaCO3/HCO3− decreases with increasing residence time on the plate, and thus decreasing supersaturation with respect to calcite. Thus, isotope fractionation depends on precipitation rate and consequently occurs under kinetic conditions. This is in contrast to previous studies, which found no rate-dependence and no or even a positive carbon isotope fractionation between CaCO3 and HCO3−. Oxygen isotope fractionation, 18εCaCO3/HCO3−, is also negative and dependent on precipitation rate. Since no literature values for 18εCaCO3/HCO3− are available, we calculated 18εCaCO3/HCO3− using equilibrium oxygen isotope fractionation factors between water and calcite and water and HCO3−, respectively. At the beginning of the plate, the fractionation is in agreement with the fractionation calculated using fractionation factors determined in cave systems.The observed fractionation between CaCO3 and water, 1000ln18α, is also in good agreement with the values determined in cave systems and shows a very similar temperature dependence 1000ln18α=16.516±1.267∗103T−26.141±4.356. However, with progressive precipitation of CaCO3 along the plate, the system is forced out of isotope equilibrium with the water, and 1000ln18α increases.The large, negative, rate-dependent isotope fractionations observed in this study suggest that precipitation of speleothem calcite is strongly kinetically controlled and may, thus, have a large effect on speleothem δ18O and δ13C values. Since these values may erroneously be interpreted as reflecting changes in past temperature, precipitation and/or vegetation density, these results have important implications for paleoclimate reconstructions from speleothems.

Triple oxygen isotope fractionation in the DIC-H2O-CO2 system: A numerical framework and its implications

Triple oxygen isotope composition of carbonate minerals reflects the isotope composition of the water from which carbonates precipitate, and is emerging as a promising proxy for constraining past changes in Earth hydroclimate and surface environment. However, quantitative interpretation of this proxy is not straightforward when carbonate minerals do not form under isotope equilibrium with water. For these carbonates, isotope composition of dissolved inorganic carbon (DIC) in the precipitating solution usually exerts the dominant control on their isotope composition. Here we examine the systematics of triple oxygen isotope fractionation in the DIC-H2O-CO2 system by deriving the fundamental equilibrium and kinetic triple oxygen isotope fractionation factors in this system and simulating the evolution of triple oxygen isotope composition of DIC during three common isotope fractionation processes, i.e., DIC-H2O isotope exchange, CO2 degassing and CO2 absorption.We show that under thermodynamic equilibrium dissolved HCO3– and CO32– both exhibit similar Δ′17O as calcite and aragonite (within 10 per meg at 25 °C), with temperature dependences around 0.61 per meg/°C. However, kinetic isotope fractionations associated with CO2 hydration and hydroxylation and their reverse reactions can produce a range of disequilibrium triple oxygen isotope effects in DIC during all three processes we simulated. The magnitudes of these disequilibrium effects vary with both time and the physicochemical conditions of the solution, e.g., temperature, pH and initial composition. Particularly, we predict correlated enrichments in Δ′17O and δ18O of DIC during CO2 degassing but depletions during CO2 absorption. The slopes of these correlations vary mainly as a function of solution pH but not temperature, DIC concentration or air pCO2, yielding values of 8.8 and 12.0 per meg/‰ at pH = 8 and 9, respectively, at 25 °C. Such disequilibrium isotope effects in DIC are expected to be inherited by carbonate minerals that form from these solutions (e.g., speleothem, coral skeleton, and high pH travertine) and, if not accounted for, could lead to inaccurate estimates of the triple oxygen isotope compositions of the parent water and carbonate formation temperatures. Our numerical model provides a quantitative framework for interpreting triple oxygen isotope composition of DIC and for correcting disequilibrium triple oxygen isotope effects in natural carbonates.

Oxygen isotope fractionation during smithsonite formation from aqueous solutions

Oxygen isotope fractionation between carbonate minerals and water is used as an environmental proxy to estimate mineral formation temperatures or isotopic composition of precipitating fluids. To date no experimental data on the oxygen isotope fractionation factor between smithsonite (ZnCO3) and water, α(18O)smithsonite-water = (18O/16O)smithsonite / (18O/16O)water, exist. Therefore, in the present study experimental work on smithsonite synthesis in the temperature range between 25 and 80 °C is coupled with ab-initio based theoretical calculations.Laboratory precipitation experiments took place in titanium reactors at elevated pCO2 (~10 atm) in order to induce the formation of smithsonite from hydrozincite (Zn5(CO3)2(OH)6), which is the precursor phase initially formed at 25 °C and low pCO2 (pH ~ 6.8). The constant α(18O)smithsonite-water = (1000 + δ18Osmithonite) / (1000 + δ18Owater) value reached at a reaction time ≥ 10,000 min (7 days) suggests near equilibrium conditions. Based on the experimentally obtained temperature relation of α(18O)smithsonite-water at 25, 40, 60, and 80 °C the integrated equation can be linearly described by the function:103lnα18Osmithsonite−water=2.79∗106/T2–0.95±0.06∗106/T2+0.60where the temperature is in Kelvin. The ab initio calculations suggest that this relation can be described in the temperature range from 0 to 100 °C as:103lnαsmithsonite−liquid water=3.21∗106/T2–3.63±0.025∗106/T2+0.90The α(18O)smithsonite-water values from the experimental approach fit within error to the theoretical relationship from the literature and the above ab initio calculations. Difference in slope between the experiment and theoretical obtained equation likely reflects modeling inaccuracies, whereas kinetic effects cannot be completely ruled out in the experimental approach.The obtained α(18O)smithsonite-water values match with the general sequence of Zn2+ < Fe2+ < Mn2+ < Ca2+ for mono-cation trigonal Me-carbonate minerals suggesting incorporation of lighter oxygen isotopes in the carbonate mineral at increasing cation radius as indicated from thermodynamic considerations. Potential applications of oxygen isotope fractionation during smithsonite formation for natural aqueous surroundings are discussed.

Rare-earth element and stable isotope signatures of kaolin from a Pliocene lateritic weathering profile at mid-latitude region (Andalusia, Spain): Implications for paleoweathering and paleoclimatic reconstructions

This study reports the impact of weathering on behavior of rare earth elements (REE) in a coastal lateritic profile developed on Pliocene sediments at mid-latitude location (~37° N), in the Guadalquivir Basin (Spain). It also explores the stable isotope signature of kaolinite providing constraints on paleotemperature and isotopic composition of the ancient meteoric water. The paleoweathering profile (~20 m thick) comprises a white sandy clayey saprolite overlain by a red-mottled clay zone, which in turn is capped by a pisolitic ferricrete. The kaolinitic regolith was practically reduced to a mixture of kaolinite and quartz by intense chemical weathering under warm and seasonally humid climate.The kaolinized material is markedly depleted in total REE (mean value of 77.30 mg kg−1 in the white saprolite and 72.70 mg kg−1 in the mottled zone) relative to the parent rock (168.96 mg kg−1). Development of acidic and oxidizing conditions promoted a suitable soil environment for REE release and leaching with percolating rainwater, leading to an apparent loss of REE from the profile (up to 87%). The parent rock-normalized REE patterns display concave-upward shapes tracing a remarkable depletion in middle REE (MREE), with no significant anomalies. The REE concentrations normalized against the North American Shale Composite (NASC) showed consistently similar fractionation patterns. The MREE-depleted signature (NASC-normalized ratios of La/Sm up to 3.32 and Gd/Lu as low as 0.45) probably arises from reductive dissolution of iron oxyhydroxide phases due to seasonal water saturation, as indicated from redoximorphic features.The <2 μm kaolinite separates showed δ18O values ranging from 17.3‰ to 20.0‰, and δD values between −71‰ and −60‰. The δ18O and δD mean values indicated a crystallization temperature of about 24 °C, which is higher than the local present-day annual mean air temperature (~18 °C). For the calculated temperature, the oxygen isotope fractionation factor between kaolinite and water (αk-w = 1.0248) implies that kaolinite formed in equilibrium with meteoric waters (δ18O = −5.4‰ and δD = −33‰) that closely reflect the regional meteoric precipitation.The lateritic weathering profiles evolved during the mid-Pliocene warm period in the southwestern Iberian margin could be useful for predicting potential future environmental effects of increased atmospheric CO2 on the Earth's critical zone.

Oxygen isotope analysis of bacterial and fungal manganese oxidation.

The ability of micro-organisms to oxidize manganese (Mn) from Mn(II) to Mn(III/IV) oxides transcends boundaries of biological clade or domain. Many bacteria and fungi oxidize Mn(II) to Mn(III/IV) oxides directly through enzymatic activity or indirectly through the production of reactive oxygen species. Here, we determine the oxygen isotope fractionation factors associated with Mn(II) oxidation via various biotic (bacteria and fungi) and abiotic Mn(II) reaction pathways. As oxygen in Mn(III/IV) oxides may be derived from precursor water and molecular oxygen, we use a twofold approach to determine the isotope fractionation with respect to each oxygen source. Using both 18 O-labeled water and closed-system Rayleigh distillation approaches, we constrain the kinetic isotope fractionation factors associated with O atom incorporation during Mn(II) oxidation to -17.3‰ to -25.9‰ for O2 and -1.9‰ to +1.8‰ for water. Results demonstrate that stable oxygen isotopes of Mn(III/IV) oxides have potential to distinguish between two main classes of biotic Mn(II) oxidation: direct enzymatic oxidation in which O2 is the oxidant and indirect enzymatic oxidation in which superoxide is the oxidant. The fraction of Mn(III/IV) oxide-associated oxygen derived from water varies significantly (38%-62%) among these bio-oxides with only weak relationship to Mn oxidation state, suggesting Mn(III) disproportionation may account for differences in the fraction of mineral-bound oxygen from water and O2 . Additionally, direct incorporation of molecular O2 suggests that Mn(III/IV) oxides contain a yet untapped proxy of δ18OO2 of environmental O2 , a parameter reflecting the integrated influence of global respiration, photorespiration, and several other biogeochemical reactions of global significance.

First-principles calculations of equilibrium fractionation of O and Si isotopes in quartz, albite, anorthite, and zircon

In this study, we used first-principles calculations based on density functional theory to investigate silicon and oxygen isotope fractionation factors among the most abundant major silicate minerals in granites, i.e., quartz and plagioclase (including albite and anorthite), and an important accessory mineral zircon. Combined with previous results of minerals commonly occurring in the crust and upper mantle (orthoenstatite, clinoenstatite, garnet, and olivine), our study reveals that the Si isotope fractionations in minerals are strongly correlated with SiO4 tetrahedron volume (or average Si–O bond length). The 30Si enrichment order follows the sequence of quartz > albite > anorthite > olivine ≈ zircon > enstatite > diopside, and the 18O enrichment follows the order of quartz > albite > anorthite > enstatite > zircon > olivine. Our calculation predicts that measurable fractionation of Si isotopes can occur among crustal silicate minerals during high-temperature geochemical processes. This work also allows us to evaluate Si isotope fractionation between minerals and silicate melts with variable compositions. Trajectory for δ30Si variation during fractional crystallization of silicate minerals was simulated with our calculated Si isotope fractionation factors between minerals and melts, suggesting the important roles of fractional crystallization to cause Si isotopic variations during magmatic differentiation. Our study also predicts that δ30Si data of ferroan anorthosites of the Moon can be explained by crystallization and aggregation of anorthite during lunar magma ocean processes. Finally, O and Si isotope fractionation factors between zircon and melts were estimated based on our calculation, which can be used to quantitatively account for O and Si isotope composition of zircons crystallized during magma differentiation.

Experimental calibration of silicon and oxygen isotope fractionations between quartz and water at 250 °C by in situ microanalysis of experimental products and application to zoned low δ30Si quartz overgrowths

The interpretation of silicon isotope data for quartz is hampered by the lack of experimentally determined fractionation factors between quartz and fluid. Further, there is a large spread in published oxygen isotope fractionation factors at low temperatures, primarily due to extrapolation from experimental calibrations at high temperature. We present the first measurements of silicon isotope ratios from experimentally precipitated quartz and estimate the equilibrium fractionation vs. dissolved silica using a novel in situ analysis technique applying secondary ion mass spectrometry to directly analyze experimental products. These experiments also yield a new value for oxygen isotope fractionation. Quartz overgrowths up to 235μm thick were precipitated in silica–H2O–NaOH–NaCl fluids, at pH12–13 and 250°C. At this temperature, 1000lnα30Si(Qtz–fluid)=0.55±0.10‰ and 1000lnα18O(Qtz–fluid)=10.62±0.13‰, yielding the relations 1000lnα30Si(Qtz–fluid)=(0.15±0.03)*106/T2 and 1000lnα18O(Qtz–fluid)=(2.91±0.04)*106/T2 when extended to zero fractionation at infinite temperature. Values of δ30Si(Qtz) from diagenetic cement in sandstones from the basal Cambrian Mt. Simon Formation in central North America range from 0 to −5.4‰. Paired δ18O and δ30Si values from individual overgrowths preserve a record of Precambrian weathering and fluid transport. The application of the experimental quartz growth results to observations from natural sandstone samples suggests that precipitation of quartz at low temperatures in nature is dominated by kinetic, rather than equilibrium, processes.

Oxygen isotope fractionation in the portlandite–water and brucite–water systems from 125 to 450 °C, 50 MPa

Equilibrium oxygen isotope fractionation factors were determined for the portlandite–water and brucite–water systems from 125 to 425°C, 50MPa using the partial exchange technique. Reagent grade cryptocrystalline Ca(OH)2 and amorphous Mg(OH)2 were reacted with three waters having different initial δ18O compositions. Isotope exchange occurred via recrystallization with exchange varying from 40% to 95% at 200 to 425°C, respectively. Equilibrium 18O brucite–water fractionation factors (103lnα) increase from −4.7±3.5‰ at 200°C to −3.5±2.5‰ at 425°C. These data connect smoothly with previous experimental calibrations at lower and higher temperatures to define a single function valid from 15 to 450°C, as follows:103lnαbrucite-water=4.39×106T2-16.95×103T+11.19where T is temperature in Kelvin. These results confirm the existence of a broad minimum in the fractionation factor for brucite at ∼250°C. The equilibrium 18O fractionation factor for portlandite–water varies from −11.1±2.7‰ at 125°C to −6.6±0.1‰ at 425°C, and can be described by the following function:103lnαportlandite-water=5.61×106T2-26.29×103T+19.72where T is temperature in Kelvin. These experimental results indicate that brucite favors 18O relative to portlandite with brucite–portlandite fractionation decreasing from 8‰ to 3‰ from 125 to 425°C. A significant temperature dependent cation mass effect is therefore indicated for cation–OH bonds in hydroxide minerals. The observed fractionation is consistent with quantum theory which predicts that bonds with less massive cations have higher vibrational frequencies and will display a relative affinity for 18O to stabilize the structure. Brucite–portlandite 18O fractionation predicted using the increment method is extremely small, opposite in sign (−0.1‰ to −0.2‰), and shows very little dependence on temperature, in poor agreement with the experimental calibration. This indicates that the method does not adequately account for the effect of cation mass on 18O fractionation within hydroxide minerals. It is suggested that cation-specific parameters within the increment method could be fit to the experimental calibrations reported here to improve prediction of fractionation factors for hydroxides and hydroxyl-bearing aluminosilicates, particularly at low temperate where the cation-mass effect is more significant.