Oxygen Isotope Partitioning Research Articles

Sulfur and Oxygen Isotope Records of Sulfate-Driven Anaerobic Oxidation of Methane in Diffusion-Dominated Marine Sediments

Organoclastic sulfate reduction (OSR) and sulfate-driven anaerobic oxidation of methane (SD-AOM) are the two major microbial pathways for sulfate consumption in marine sulfur cycle. The relative changes of sulfur and oxygen isotope ratios in pore water sulfate are affected by the mode of microbial sulfate reduction and have been applied as an indicator for assessing methane excess environments. However, so far, this isotope proxy fails to distinguish sulfate reduction processes fueled by the oxidation of organic matter or by diffusing methane. To better understand the mechanism of sulfur and oxygen isotope partitioning during OSR and SD-AOM, coupled sulfur and oxygen isotopic compositions of pore water sulfate (δ34SSO4 and δ18OSO4) were investigated from four methane diffusing sites (CL56, CL57, CL59, and CL60) of the South China Sea, supplemented by carbon isotopic composition of dissolved inorganic carbon (DIC) and sulfur isotopic composition of pyrite in bulk sediments. Pore water sulfate and DIC concentrations, as well as calculated net sulfate reduction rates suggest that the sulfate reduction at site CL57 was mainly dominated by OSR, whereas sites CL56, CL59, and CL60 were likely impacted by both OSR and SD-AOM. Furthermore, the trend of cross-plotting δ18OSO4 versus δ34SSO4 values from site CL57 was distinguishable from sites CL56, CL59, and CL60, although all study sites show similar patterns to those derived from methane limited environments. This further indicates the trajectory of sulfur and oxygen isotope partitioning was affected by the mode of sulfate reduction (i.e., OSR vs. SD-AOM). At site CL57, the low net sulfate reduction rate would lead to enhanced oxidation of intermediate sulfur species during OSR, thus leading to a higher slope in the δ18OSO4 vs. δ34SSO4 cross-plot (1.26). In contrast, the higher net sulfate reduction rates at sites CL56, CL59, and CL60 due to the impact from SD-AOM would lead to lower slopes in the δ18OSO4 vs. δ34SSO4 cross-plots (0.78 ± 0.11). This study provides new insights into the sulfur and oxygen isotope systematics during microbial sulfate reduction processes in methane diffusing environments.

Partitioning of oxygen isotopes during the aqueous solvation of nitric acid

Mixtures of water and inorganic acids show differences in isotopic composition between the vapor and liquid phases, reportedly even at an azeotrope. A recent reactive azeotropy analysis rationalized this by allowing for differences in stable isotope composition between bound solvation complexes and the bulk liquid. The objective of this study is to determine whether or not significant differences in stable isotope composition exist between bound solvation complexes involving inorganic acids and the bulk solution. Rayleigh distillation of a solution of nitric acid and water was conducted. During the distillation, it was expected that bulk water would be vaporized first while (bound) water involved with hydrating the acid would be vaporized at later times along with the acid. Samples of both distillate and residual liquid were collected over time intervals and analyzed to determine oxygen and nitrogen stable isotope composition using isotope ratio mass spectrometry. As a reference, comparative samples were collected from Rayleigh distillation of water alone. The measured oxygen fractionation factors suggest that the acid is preferentially hydrated with 18O-enriched water, compared to the unbound, or bulk, water. Preferential coordination of 18O-enriched water with the dissolved acid is consistent with heavier water forming stronger bonds relative to light water. The results, although obtained under non-equilibrum conditions, suggest that isotope-dependent partitioning between bulk solvent and dissolved complexes may prove necessary in developing accurate descriptions of isotope fractionation during vapor-liquid equilibrium.

Glacial–interglacial temperature change in the tropical West Pacific: A comparison of stalagmite-based paleo-thermometers

In the tropics, geochemical records from stalagmites have so far mainly been used to qualitatively reconstruct changes in precipitation, but several new methods to reconstruct past temperatures from stalagmite material have emerged recently: i) liquid–vapor homogenization of fluid inclusion water ii) noble gas concentrations in fluid inclusion water, iii) the partitioning of oxygen isotopes between fluid inclusion water and calcite, and iv) the abundance of the 13C18O16O (‘clumped’) isotopologue in calcite. We present, for the first time, a direct comparison of these four paleo-thermometers by applying them to a fossil stalagmite covering nearly two glacial–interglacial cycles (Marine Isotope Stages (MIS) 12–9) and to two modern stalagmites, all from northern Borneo. The temperature estimates from the different methods agree in most cases within errors for both the old and recent samples; reconstructed formation temperatures of the recent samples match within 2-sigma errors with measured cave temperatures. However, slight but systematic deviations are observed between noble gas and liquid–vapor homogenization temperatures. Whereas the temperature sensitivity of fluid inclusion δ18O and clumped isotopes is currently debated, we find that the calibration of Tremaine et al. (2011) for fluid inclusion δ18O and a synthetic calcite-based clumped isotope calibration (Ziegler et al., in prep.) yield temperature estimates consistent with the other methods. All methods (with the potential exception of clumped isotopes) show excellent agreement on the amplitude of glacial–interglacial temperature change, indicating temperature shifts of 4–5 °C. This amplitude is similar to the amplitude of Mg/Ca-based regional sea surface temperature records, when correcting for sea level driven changes in cave elevation. Our reconstruction of tropical temperature evolution over the time period from 440 to 320 thousand years ago (ka) adds support to the view that climate sensitivity to varying greenhouse forcing is substantial also in the deep tropics.

Oxygen isotope fractionation in double carbonates

Oxygen isotope fractionations in double carbonates of different crystal structures were calculated by the increment method. Synthesis experiments were performed at 60 °C and 100 °C to determine oxygen and carbon isotope fractionations involving PbMg[CO3]2. The calculations suggest that the double carbonates of calcite structure are systematically enriched in 18O relative to those of aragonite and mixture structures. Internally consistent oxygen isotope fractionation factors are obtained for these minerals with respect to quartz, calcite and water at a temperature range of 0–1200 °C. The calculated fractionation factors for double carbonate–water systems are generally consistent with the data available from laboratory experiments. The experimentally determined fractionation factors for PbMg[CO3]2, BaMg[CO3]2 and CaMg[CO3]2 against H2O not only fall between fractionation factors involving pure carbonate end-members but are also close to the calculated fractionation factors. In contrast, experimentally determined carbon isotope fractionation factors between PbMg[CO3]2 and CO2 are much closer to theoretical predictions for the cerussite–CO2 system than for the magnesite–CO2 system, similar to the fractionation behavior for BaMg[CO3]2. Therefore, the combined theoretical and experimental results provide insights into the effects of crystal structure and exchange kinetics on oxygen isotope partitioning in double carbonates.

Oxygen isotope partitioning between immiscible silicate melts with H2O, P and S

Differences between the δ18O values of immiscible Si- and Fe-rich melts in the systems Fe2SiO4–Fe3O4–KAlSi2O6–SiO2, Fe3O4–KAlSi2O6–SiO2, and Fe3O4–Fe2O3–KAlSi2O6–SiO2, with H2O, H2O+P or H2O+S have been determined in isothermal, isobaric experiments at 1100 and 1200°C and 200MPa. The Δ18O values for conjugate Fe2SiO4–Fe3O4–KAlSi2O6–SiO2+H2O and, Fe3O4–KAlSi2O6–SiO2·KAlSi2O6–SiO2+H2O melts are only 0.4–0.6‰ and do not differ significantly from those for anhydrous melts of similar composition. The Δ18O values for melts with added H2O+P or S are more variable, ranging from 0.0 to 0.8‰. Partitioning of 18O between the immiscible melts is 0.6–1‰ less than the partitioning reported for melt–mineral and mineral–mineral pairs. The partitioning of 18O in the network modifier-bearing immiscible melts is not controlled by the relative degree of polymerization in the melts or fO2. The upper limit of the range of Δ18O values (<1‰), and the variation in the δ18O values of conjugate melts that occurs with the inclusion of network modifying constituents, suggest that in some cases, oxygen isotope ratios might be useful to distinguish lithologies evolved from coexisting immiscible silicate melts, from lithologies that have evolved by crystal fractionation only.

Intra- and Intercrystalline Oxygen Isotope Variations in Minerals from Basalts and Peridotites

Igneous phenocrysts commonly exhibit zoning in major and trace element composition, reflecting (and potentially constraining) the differentiation and/or mixing histories of their parent melts. To date, little work has been done characterizing zonation of oxygen isotopes in minerals from mafic and ultramafic rocks. We present 259 ion probe (CAMECA ims-1280) measurements of δ^(18)O in 34 natural magmatic and mantle olivines and pyroxenes from five hand samples from diverse igneous environments. We compare δ^(18)O variations with zonation in other elements [especially P; analyzed by electron microprobe analysis (EMPA) and nano-secondary ionization mass spectrometry (nanoSIMS)]. There is generally a close (average within ~0·1–0·2 ‰) agreement between average δ^(18)O values of olivines measured by SIMS (standardized against San Carlos olivine) and independently known values for bulk separates from the same samples measured by laser fluorination. These data demonstrate that current ion microprobe techniques are not only precise but also accurate enough for study of sub-per-mil oxygen isotope variations in silicates (within ~0·2 ‰), provided samples are prepared and analyzed following strict guidelines. All but one of the 34 studied grains are homogeneous in δ^(18)O within a small multiple of analytical precision [estimated ±0·2‰, 1σ for most data; poorer for a subset of measurements made on small (~5 µm) spots]. This population of isotopically homogeneous grains includes some with oscillatory micrometer-scale P banding. The lack of δ^(18)O variations suggests that whatever factors lead to this common mode of trace element zonation have no detectable effect on melt–crystal partitioning of oxygen isotopes. Large (2‰) oxygen isotope variations are observed in one olivine glomerocryst from Mauna Kea, Hawaii. This glomerocryst contains P-rich domains that are either equant or skeletal or feathery in outline, and these P-rich domains are systematically low in δ^(18)O compared with adjacent, later-grown, P-poor olivine. This unusual oxygen isotope zonation pattern might reflect a kinetic fractionation during nucleation and growth of the cores of some olivine phenocrysts. We tested this hypothesis through measurements of δ^(18)O distributions in synthetic olivines grown at a range of rates and exhibiting diverse patterns of P zoning. These synthetic olivines are homogeneous in δ^(18)O, within the limits of our analyses (± 0·3–0·4‰ in this case) and show no connection between P zonation and oxygen isotope heterogeneity. We therefore think it more plausible that unusual O isotope zonation in the Mauna Kea glomerocryst reflects addition of a low-δ^(18)O component to some Hawaiian magmas just before nucleation of olivine. More generally, this study demonstrates the utility of modern SIMS techniques for in situ study of the subtle (~1‰ range) oxygen isotope variations characteristic of common mafic and ultramafic rocks.

Experimental calibration of oxygen isotope fractionation between quartz and zircon

We report the results of an experimental calibration of oxygen isotope fractionation between quartz and zircon. Data were collected from 700 to 1000 °C, 10–20 kbar, and in some experiments the oxygen fugacity was buffered at the fayalite–magnetite–quartz equilibrium. Oxygen isotope fractionation shows no clear dependence on oxygen fugacity or pressure. Unexpectedly, some high-temperature data (900–1000 °C) show evidence for disequilibrium oxygen isotope partitioning. This is based in part on ion microprobe data from these samples that indicate some high-temperature quartz grains may be isotopically zoned. Excluding data that probably represent non-equilibrium conditions, our preferred calibration for oxygen isotope fractionation between quartz and zircon can be described by: 1000 ln α qtz – zrc = ( 2.33 ± 0.24 ) × 10 6 / T 2 ( in K ) This relationship can be used to calculate fractionation factors between zircon and other minerals. In addition, results have been used to calculate WR/melt–zircon fractionations during magma differentiation. Modeling demonstrates that silicic magmas show relatively small changes in δ 18O values during differentiation, though late-stage mafic residuals capable of zircon saturation contain elevated δ 18O values. However, residuals also have larger predicted melt–zircon fractionations meaning zircons will not record enriched δ 18O values generally attributed to a granitic protolith. These results agree with data from natural samples if the zircon fractionation factor presented here or from natural studies is applied.

Oxygen isotope partitioning during oxidation of pyrite by H 2O 2 and its dependence on temperature

A detailed experimental study was conducted to investigate mechanisms of pyrite oxidation by determining product yields and oxygen isotopic fractionation during reactions between powdered pyrite (FeS 2) with aqueous hydrogen peroxide (H 2O 2). Sealed silica-tube experiments utilized aliquots of pyrite that were reacted with 0.2 M H 2O 2 for 7 to 14 days at 4 to 150 °C. No volatile sulfur species were detected in any experiment. The only gaseous product recovered was elemental oxygen inferred to result from decomposition of H 2O 2. Aqueous sulfate (S aq) was the only sulfur product recovered from solution. Solid hydrated ferric iron sulfates (i.e., water-soluble sulfate fraction, S ws) were recovered from all experiments. Ferric oxide (hematite) was detected only in high temperature experiments. Reactants were selected with large differences in initial δ 18O values. The oxygen isotopic compositions of oxygen-bearing reactants and products were analyzed for each experiment. Subsequent isotopic mass-balances were used to identify sources of oxygen for reaction products and to implicate specific chemical reaction mechanisms. δ 18O of water did not show detectable change during any experiment. δ 18O of sulfate was similar for S aq and S ws and indicated that both H 2O and H 2O 2 were sources of oxygen in sulfate. Low-temperature experiments suggest that H 2O-derived oxygen was incorporated into sulfate via Fe 3+ oxidation, whereas H 2O 2-derived oxygen was incorporated into sulfate via oxidation by hydroxyl radicals (HO ). These two competing mechanisms for oxygen incorporation into sulfate express comparable influences at 25 °C. With increasing reaction temperatures from 4 to 100 °C, it appears that accelerated thermal decomposition and diminished residence time of H 2O 2 limit the oxygen transfer from H 2O 2 into sulfate and enhance the relative importance of H 2O-derived oxygen for incorporation into sulfate. Notably, at temperatures between 100 and 150 °C there is a reversal in the lower temperature trend resulting in dominance of H 2O 2-derived oxygen over H 2O-derived oxygen. At such high temperatures, complete thermal decomposition of H 2O 2 to water and molecular oxygen (O 2) occurs within minutes in mineral-blank experiments and suggests little possibility for direct oxidation of pyrite by H 2O 2 above 100 °C. We hypothesize that a Fe–O 2 mechanism is responsible for oxygenating pyrite to sulfate using O 2 from the preceding thermal decomposition of H 2O 2.

Spatial and temporal zoning of hydrothermal alteration and mineralization in the Sossego iron oxide–copper–gold deposit, Carajás Mineral Province, Brazil: paragenesis and stable isotope constraints

The Sossego iron oxide–copper–gold deposit (245 Mt @ 1.1% Cu, 0.28 g/t Au) in the Carajas Mineral Province of Brazil consists of two major groups of orebodies (Pista–Sequeirinho–Baiano and Sossego–Curral) with distinct alteration assemblages that are separated from each other by a major high angle fault. The deposit is located along a regional WNW–ESE-striking shear zone that defines the contact between metavolcano–sedimentary units of the ∼2.76 Ga Itacaiunas Supergroup and tonalitic to trondhjemitic gneisses and migmatites of the ∼2.8 Ga Xingu Complex. The deposit is hosted by granite, granophyric granite, gabbro, and felsic metavolcanic rocks. The Pista–Sequeirinho–Baiano orebodies have undergone regional sodic (albite–hematite) alteration and later sodic–calcic (actinolite-rich) alteration associated with the formation of massive magnetite–(apatite) bodies. Both these alteration assemblages display ductile to ductile–brittle fabrics. They are cut by spatially restricted zones of potassic (biotite and potassium feldspar) alteration that grades outward to chlorite-rich assemblages. The Sossego–Curral orebodies contain weakly developed early albitic alteration and very poorly developed subsequent calcic–sodic alteration. These orebodies contain well-developed potassic alteration assemblages that were formed during brittle deformation that resulted in the formation of breccia bodies. Breccia matrix commonly displays coarse mineral infill suggestive of growth into open space. Sulfides in both groups of deposits were precipitated first with potassic alteration and more importantly with a later assemblage of calcite–quartz–epidote–chlorite. In the Sequeirinho orebodies, sulfides range from undeformed to deformed; sulfides in the Sossego–Curral orebodies are undeformed. Very late, weakly mineralized hydrolytic alteration is present in the Sossego/Currral orebodies. The sulfide assemblage is dominated by chalcopyrite with subsidiary siegenite, and millerite. Pyrrhotite and pyrite are minor constituents of ore in the Sequerinho orebodies while pyrite is relatively abundant in the Sossego–Curral bodies. Oxygen isotope partitioning between mineral pairs constrains temperatures in the deposit spatially and through time. In the Sequeirinho orebody, the early sodic–calcic alteration stage was characterized by temperatures exceeding 500°C and $$ \delta ^{{{\text{18}}}} {\text{O}}_{{{\text{H}}_{{\text{2}}} {\text{O}}}} $$ values for the alteration fluid of 6.9 ± 0.9‰. Temperature declines outward and upward from the zone of most intense alteration. Paragenetically later copper–gold mineralization displays markedly lower temperatures (<300°C) and was characterized by the introduction of 18O-depleted hydrothermal fluids −1.8 ± 3.4‰. The calculated δDH2O and $$ \delta ^{{{\text{18}}}} {\text{O}}_{{{\text{H}}_{{\text{2}}} {\text{O}}}} $$ values suggest that the fluids that formed the early calcic–sodic alteration assemblage were of formational/metamorphic or magmatic origin. The decrease of $$ \delta ^{{{\text{18}}}} {\text{O}}_{{{\text{H}}_{{\text{2}}} {\text{O}}}} $$ values through time may reflect influx of surficially derived waters during later alteration and mineralization events. Influx of such fluids could be related to episodic fluid overpressure, resulting in dilution and cooling of the metalliferous fluid, causing deposition of metals transported as metal chloride complexes.

Variation in oxygen isotope fractionation during cellulose synthesis: intramolecular and biosynthetic effects

The oxygen isotopic composition of plant cellulose is commonly used for the interpretations of climate, ecophysiology and dendrochronology in both modern and palaeoenvironments. Further applications of this analytical tool depends on our in-depth knowledge of the isotopic fractionations associated with the biochemical pathways leading to cellulose. Here, we test two important assumptions regarding isotopic effects resulting from the location of oxygen in the carbohydrate moiety and the biosynthetic pathway towards cellulose synthesis. We show that the oxygen isotopic fractionation of the oxygen attached to carbon 2 of the glucose moieties differs from the average fractionation of the oxygens attached to carbons 3-6 from cellulose by at least 9%, for cellulose synthesized within seedlings of two different species (Triticum aestivum L. and Ricinus communis L.). The fractionation for a given oxygen in cellulose synthesized by the Triticum seedlings, which have starch as their primary carbon source, is different than the corresponding fractionation in Ricinus seedlings, within which lipids are the primary carbon source. This observation shows that the biosynthetic pathway towards cellulose affects oxygen isotope partitioning, a fact heretofore undemonstrated. Our findings may explain the species-dependent variability in the overall oxygen isotope fractionation during cellulose synthesis, and may provide much-needed insight for palaeoclimate reconstruction using fossil cellulose.

Oxygen isotope exchange and disequilibrium between calcite and tremolite in the absence and presence of an experimental C?O?H fluid

Oxygen isotope exchange between minerals during metamorphism can occur in either the presence or the absence of aqueous fluids. Oxygen isotope partitioning among minerals and fluid is governed by both chemical and isotopic equilibria during these processes, which progress by intragranular and intergranular diffusion as well as by surface reactions. We have carried out isotope exchange experiments in two- and three-phase systems, respectively, between calcite and tremolite at high temperatures and pressures. The two-phase system experiments were conducted without fluid either at 1 GPa and 680 °C for 7 days or at 500 MPa and 560 °C for 20 days. Extrapolated equilibrium fractionations between calcite and tremolite are significantly lower than existing empirical estimates and experimental determinations in the presence of small amounts of fluid, but closely match calculated fractionations by means of the increment method for framework oxygen in tremolite. The small fractionations measured in the direct calcite–tremolite exchange experiments are interpreted by different rates of oxygen isotope exchange between hydroxyl oxygen, framework oxygen and calcite during the solid–solid reactions where significant recrystallization occurs. The three-phase system experiments were accomplished in the presence of a large amount of fluid (CO2+H2O) at 500 MPa and 560 °C under conditions of phase equilibrium for 5, 10, 20, 40, 80, 120, 160, and 200 days. The results show that oxygen isotope exchange between minerals and fluid proceeds in two stages: first, through a mechanism of dissolution-recrystallization and very rapidly; second, through a mechanism of diffusion and very slowly. Synthetic calcite shows a greater rate of isotopic exchange with fluid than natural calcite in the first stage. The rate of oxygen diffusion in calcite is approximately equal to or slightly greater than that in tremolite in the second stage. A calculation using available diffusion coefficients for calcite suggests that grain boundary diffusion, rather than volume diffusion, has been the dominant mechanism of oxygen transport between the fluid and the mineral grains in the later stage.

Kinetic mechanism of oxygen isotope disequilibrium in precipitated witherite and aragonite at low temperatures: an experimental study

To study what dictates oxygen isotope equilibrium fractionation between inorganic carbonate and water during carbonate precipitation from aqueous solutions, a direct precipitation approach was used to synthesize witherite, and an overgrowth technique was used to synthesize aragonite. The experiments were conducted at 50 and 70°C by one- and two-step approaches, respectively, with a difference in the time of oxygen isotope exchange between dissolved carbonate and water before carbonate precipitation. The two-step approach involved sufficient time to achieve oxygen isotope equilibrium between dissolved carbonate and water, whereas the one-step approach did not. The measured witherite-water fractionations are systematically lower than the aragonite-water fractionations regardless of exchange time between dissolved carbonate and water, pointing to cation effect on oxygen isotope partitioning between the barium and calcium carbonates when precipitating them from the solutions. The two-step approach experiments provide the equilibrium fractionations between the precipitated carbonates and water, whereas the one-step experiments do not. The present experiments show that approaching equilibrium oxygen isotope fractionation between precipitated carbonate and water proceeds via the following two processes: 1.Oxygen isotope exchange between [CO3]2− and H2O: (1)[C16O3]2−+2H218O=[C18O216O]2−+2H216O2.A combination of divalent metal cation M2+ with the [CO3]2− to form carbonate: (2)M2++[C18O216O]2−+H218O=MC18O3+H216OReaction 1 is the rate-limiting step for equilibrium oxygen isotope fractionation between carbonate and water, and reaction 2 bears the structural effect of carbonate crystallization on oxygen isotope fractionation during carbonate precipitation. The present results show that the time of [CO3]2−-H2O isotope exchange in the precipitation experiments is of critical importance in dictating the extent of isotopic equilibration between precipitated carbonate and water. If the time of oxygen isotope exchange between [CO3]2− and H2O is long enough to attain equilibrium before carbonate precipitation, the precipitated carbonate is able to achieve isotopic equilibration with water.

Brachiopod Shells: Recorders of the Present and Keys to the Past

The Brachiopod shell provides more than protection for the organism dwelling within: it provides a structural record of growth and a chemical record of fluctuations in the environment in which it formed. Typically, a shell formed by an accretionary process will have growth lines visible on the shell surface that can be used to measure chronological age directly. This is common in many fossil and modern bivalve molluscs where growth bands are prominent (see, e.g., Lutz and Rhoads, 1980; Jones, 1983; Tanabe, 1988). However, growth increments are not always so easily identified in other shelled organisms. For instance, growth lines in brachiopod shells are often indistinct, and the time interval represented by each growth line is uncertain. When growth information can't be obtained directly, we must turn to an alternative method (Jones and Gould, 1999). Fortunately, periodicities in the chemical composition of a shell may serve as proxies for age (e.g., Buening and Carlson, 1992). The shell chemistry of brachiopods reflects, for the most part, the changes in chemistry of the surrounding seawater. For instance, partitioning of oxygen isotopes (18O/16O ratio) in shell carbonate is sensitive to temperature (e.g., Hoefs, 1997). Thus, if brachiopods precipitate shells in water that is affected by seasonal temperature variations, then fluctuations of the18O/16O ratio of the shell should provide a record of annual temperature variation in ambient seawater. Consequently, this seasonal record may be used in conjunction with structural growth lines to reconstruct the life history of the individual, as well as the local environmental conditions in which the brachiopod lived.

Chemical, mineralogical and isotope behavior, and phase transformation during the precipitation of calcium carbonate minerals from intermediate ionic solution at 25°C

A set of time-series experiments was performed in which CaCO3 was precipitated (25°C, 1 atm) from aliquots of an NaHCO3 (0.330 M)/CaCl2 (0.0023 M) solution in a closed system to evaluate the effect of phase transformation on calcium carbonate precipitation kinetics and isotope behavior.Monohydrocalcite precipitated initially at a solution saturation state (Ωmhc) slightly greater than one, whereas calcite crystallized later at a significant higher saturation state (Ωcl ≅ 14.4). The precipitation of calcite promoted the dissolution of monohydrocalcite at a rate that exceeded calcite precipitation, producing anomalous behavior in the chemical and isotope composition of the system.The carbon isotope fractionation factors (103 lnα) for monohydrocalcite-HCO3(aq)- and monohydrocalcite-CO2(g) were 0.36 ± 0.01‰ and 8.35 ± 0.01‰, respectively. The oxygen isotope fractionation factor for monohydrocalcite-H2O was 27.8 ± 0.1‰.The carbon isotope fractionation factors for calcite-HCO3 (aq)− and calcite-CO2(g) were 0.94 ± 0.06‰ and 8.93 ± 0.06‰, respectively, whereas the oxygen isotope fractionation factor for calcite-H2O(l) was 28.0 ± 0.2‰. A carbon isotope fractionation factor of 0.58 ± 0.07‰ was determined for the mineral pair calcite-monohydrocalcite, but no fractionation was observed for oxygen isotopes over time steps when both minerals co-precipitated.Fractionation factors for calcite were independent of precipitation rate over the range in rates of 103.96 to 105.63 μmol/m2h. These results extend the upper limit of characterization for the relationship between precipitation rate and isotope partitioning of carbon between calcite, HCO3 (aq)− and CO2(g), and quantitatively document for the first time the independence between precipitation rate and oxygen isotope partitioning in the calcite-H2O(l) system.

Empirical determination of oxygen isotope fractionation factors for titanite with respect to zircon and quartz

Measurements of oxygen isotope fractionations between coexisting igneous titanite (Ttn) and zircon (Zrc) have been used to formulate a self-consistent, empirical calibration of equilibrium oxygen isotope partitioning: 1000 lnα(Zrc−Ttn)=1.02×106/(T2) This calibration is based on the average measured Δ(Zrc-Igneous Ttn) = 1.2 ± 0.3‰ (n = 27 rocks) and a closure temperature of titanite to oxygen diffusion of approximately 650°C. The average measured fractionation between zircon and metamorphic titanite is 2.1 ± 0.4‰ (n = 5 rocks). These results show that Δ(Zrc-Ttn) can be used to distinguish igneous vs. metamorphic (or hydrothermal) titanite.The new zircon-titanite fractionation has been combined with published experimental and empirical data to provide a new quartz-titanite fractionation curve: 1000 lnα(Qtz−Ttn)=3.57×106/(T2) This new calibration is consistent with Δ(Qtz-Igneous Ttn) data from coexisting mineral pairs in granitic rocks. Modeling diffusional exchange of oxygen during cooling in a typical granitic rock, using the new calibration, results in the same Δ(Qtz-Igneous Ttn) as measured (4.5 ± 0.43‰, n = 25 rocks). Both measured Δ(Zrc-Igneous Ttn) and Δ(Qtz-Igneous Ttn) differ from fractionations calculated using the semiempirical increment method. The consistency of empirical Δ(Zrc-Igneous Ttn) in a range of geologic environments allows the calculation of a more accurate fractionation factor for titanite.

An explanation of the effect of seawater carbonate concentration on foraminiferal oxygen isotopes

Abstract Stable oxygen isotope ratios of foraminiferal calcite are widely used in paleoceanography to provide a chronology of temperature changes during ocean history. It was recently demonstrated that the stable oxygen isotope ratios in planktonic foraminifera are affected by changes of the seawater chemistry carbonate system: the δ 18 O of the foraminiferal calcite decreases with increasing CO 3 2− concentration or p H. This paper provides a simple explanation for seawater chemistry dependent stable oxygen isotope variations in the planktonic foraminifera Orbulina universa which is derived from oxygen isotope partitioning during inorganic precipitation. The oxygen isotope fractionation between water and the dissolved carbonate species S = [H 2 CO 3 ] + [HCO 3 − ] + [CO 3 2− ] decreases with increasing p H. Provided that calcium carbonate is formed from a mixture of the carbonate species in proportion to their relative contribution to S , the oxygen isotopic composition of CaCO 3 also decreases with increasing p H. The slope of shell δ 18 O vs. [CO 3 2− ] of Orbulina universa observed in culture experiments is −0.0022‰ (μmol kg −1 ) −1 (Spero et al., 1997) , whereas the slope derived from inorganic precipitation is −0.0024‰ (μmol kg −1 ) − . The theory also provides an explanation of the nonequilibrium fractionation effects in synthetic carbonates described by Kim and O’Neil (1997) which can be understood in terms of equilibrium fractionation at different p H. The results presented here emphasize that the oxygen isotope fractionation between calcium carbonate and water does not only depend on the temperature but also on the p H of the solution from which it is formed.

Oxygen isotope fractionation in carbonate and sulfate minerals.

Oxygen isotope fractionations involving carbonates and sulfates have been controversial for a long time. There are important unresolved conflicts among the results of theoretical calculations, experimental measurements and empirical estimates. In this paper, the increment method is adapted to systematically evaluate oxygen isotope fractionations in the carbonates and sulfates. The following sequence of 18-enrichment in carbonate minerals is obtained: siderite > ankerite > magnesite ≥ dolomite > calcite > aragonite > strontianite > cerussite ≥ witherite. The sequence of 18O-enrichment in sulfate minerals is predicted as follows: anhydrite > celestite > barite > anglesite. The internally consistent fractionation factors for the systems carbonate-water and sulfate-water are acquired for a temperature range of 0 to 1200°C, which are in fair agreement with existing experimental and/or empirical data. The present calculations suggest that dolomite should behave isotopically like calcite; equilibrium fractionation between dolomite and calcite is only 0.56‰ at 25°C. Aragonite is predicted to be significantly depleted in 18O relative to calcite; equilibrium fractionation between calcite and aragonite is 4.47‰ at 25°C. It is possible that polymorphic transition from aragonite to calcite could proceed through an essentially intact oxygen structure without isotopic resetting. As a result, the temperature dependence of oxygen isotope partitioning in aragonite could be conveyed to calcite. Oxygen isotope inheritance in calcite formation by the polymorphic transition may be of critical importance in attempts to resolve dilemma involving fractionations in aragonite-calcite-dolomite-water systems.

Oxygen isotope partitioning between phosphate and carbonate in mammalian apatite

The oxygen isotope compositions of phosphate and structural carbonate in mammalian enamel and bone apatite are linked to that of body water at constant body temperature near 37°C, but the isotope systematics of oxygen in structural carbonate are not well understood. Using coupled measurements of the oxygen isotope composition of structural carbonate and phosphate from horse tooth enamel, the apparent oxygen isotope fractionation factor between structural carbonate and body water is estimated to be 1.0263 ± 0.0014. These estimates provide a quantitative basis for using the oxygen isotope composition of structural carbonate in mammalian biogenic apatite for ecological, climatological, and physiological reconstruction.

Oxygen isotope partitioning between rhyolitic glass/melt and CO 2: An experimental study at 550–950°C and 1 bar

Oxygen isotope partitioning between gaseous C02 and a natural rhyolitic glass and melt (77.7 wt% SiO 2, 0.16 wt% H 2O total) has been measured at 550–950°C and approximately 1 bar. Equilibrium oxygen isotope fractionation factors (α CO 2−rhyolite = ( 18O/ 16O) CO 2 /( 18O/ 16O) rhyolite) determined in exchange experiments of 100–255 day duration are: (°C) 1000 1n α CO 2-rhyohte 550 5.08 ± 0.13 650 4.62 ± 0.14 750 3.99 ± 0.19 850 3.71 ± 0.19 950 2.95 ± 0.16 These values agree well with predictions based on experimentally determined oxygen isotope fractionation factors for CO 2-silica glass (Stolper and Epstein, 1991) and CO 2-albitic glass/melt (Matthews et al., 1994), if the rhyolitic glass is taken to be a simple mixture of normative silica and alkali feldspar components. The results indicate that oxygen isotope partitioning in felsic glasses and melts can be modeled by linear combinations of endmember silicate constituents. Rates of oxygen isotope exchange observed in the partitioning experiments are consistent with control by diffusion of molecular H 2O dissolved in the glass/melt and are three orders of magnitude faster than predicted for rate control solely by diffusion of dissolved molecular CO 2 under the experimental conditions. Additional experiments using untreated and dehydrated (0.09 wt% H 2O total) rhyolitic glass quanitatively support these interpretations. We conclude that diffusive oxygen isotope exchange in rhyolitic glass/melt, and probably other polymerized silicate materials, is controlled by the concentrations and diffusivities of dissolved oxygen-bearing volatile species rather than diffusion of network oxygen under all but the most volatile-poor conditions.

Oxygen isotope fractionations involving apatites: Application to paleotemperature determination

The increment method is adapted to calculate oxygen isotope fractionations in phosphate minerals by taking calcite as a reference. The results show a significant difference in oxygen isotope partitioning between OH-apatite, Cl-apatite and F-apatite. The substitution of halogen for hydroxyl in apatites can thus have a considerable influence on oxygen isotope properties. The obtained order of 18O enrichment is: Cl-apatite < F-apatite < OH-apatite < monazite < xenotime. At 25°C, the oxygen isotope fractionation between OH-apatite and F-apatite is 0.7‰, and that between OH-apatite and Cl-apatite is 1.8‰. Three sets of internally consistent fractionation factors for the systems phosphate-water, quartz-phosphate and calcite-phosphate are acquired, respectively, for the temperature range of 0° to 1200°C. The inspection of existing isotope data demonstrates that the theoretical calibrations involving apatites are applicable to the paleotemperature determination of ancient oceans.