V14C-01
An anion-exchange chromatographic study on chlorine isotope effect accompanying hydration
The single-stage separation factor of the chlorine isotopes ($^{35}$Cl and $^{37}$Cl) was determined to be 1.00034 by 4.5 meter-long anion exchange chromatography operated in the reverse breakthrough manner at 25$^{o}$C. The value was in good agreement with those obtained in our previous works. It was confirmed that the lighter isotope ($^{35}$Cl) was preferentially fractionated into the resin phase whereas the heavier one ($^{37}$Cl) into the aqueous phase. This observation, however, contradicted the experimental result on the Cl isotope fractionation during NaCl precipitation and the recent theoretical result on Cl isotope fractionation: The $^{37}$Cl isotope selectively enriched into the solid phase rather than aqueous phase. The discrepancy was discussed on the basis of the theory of isotope distribution between two phases. It was suggested that the chromatographic results were contributed by the isotope effect accompanying hydration more than by the isotope effect due to the phase change, whereas the reverse was the case with the result in the NaCl precipitation study.
V14C-02
Chlorine Isotope Constraints on the Origin and Distribution of Earth's Chlorine
Volatile elements exert a strong influence over the chemical and physical properties of the Earth's mantle. Due to its incompatible, soluble and volatile element chemical characteristics, chlorine is especially valuable in understanding the current and past evolution of the Earth (e.g., melting, recycling, degassing, differentiation). Comprehensive understanding of the exchanges of chlorine among Earth's reservoirs may help constrain the origin and the budget of chlorine and other volatile elements on Earth. Due to the large uncertainties in the estimated range of mantle Cl flux inputs and outputs, we present here the chlorine isotopic compositions (\delta$^{37}$Cl) of mantle and subducted materials as well as those in chondrites with the aim of better understanding the global Cl cycle. {\bf Mantle, subduction and chondrites \delta$^{37}$Cl.} Based on fresh N- and E- MORB samples affected by various degrees of assimilation of seawater-derived materials (e.g., intergranular brines; Bonifacie et al., Chem Geol, 2005), we estimate that the \delta$^{37}$Cl value of the mean upper mantle is inferior or equal to - 1.9\permil. Analyses on HP metaperidotites from the Alps suggest that no Cl-isotopes fractionation occurs during the Cl loss associated with the dehydration of serpentines throughout prograde subduction. Considering HP metaperidotites as suitable candidates for Cl transfer to the mantle, and excluding the possible contribution of sediments, we estimate that the subducted material has \delta$^{37}$Cl values superior or equal to -1.4\permil. Various types of chondrites show relatively homogeneous \delta$^{37}$Cl values (~ -1.7$\mp$\permil). {\bf The global chlorine cycle and implications on the origin of Earth's Cl.} The slight but significant difference between the Cl-isotopic signature of recycled Cl and upper mantle implies that the \delta$^{37}$Cl value of the mantle increased while that of exogenous reservoirs decreased over geological time. Box modeling predicts for early Earth: i/ a large amount of Cl in exogenous reservoirs, and ii/ a significant \delta$^{37}$Cl difference between exogenous reservoirs and mantle. This argues in favor of an early and catastrophic degassing of Cl from the mantle (however not sufficient to explain the predicted early \delta$^{37}$Cl differences) and/or a �late veneer� supply of Cl (i.e. heterogeneous accretion). Our data also show evidence for a \delta$^{37}$Cl difference between the �Upper Earth� (i.e. exogenous reservoirs and upper mantle; -0.3\permil) and chondrites. This might be explained by a preferential loss of $^{35}$Cl during Earth accretion and/or a �late veneer� supply of Cl with positive \delta$^{37}$Cl value.
V14C-03
Stable Chlorine Isotope Fractionation
Chlorine isotope partitioning between different phases is not well understood. Pore fluids can have \delta$^{37}$Cl values as low as -8\permil, with neoform sediments having strongly positive values. Most strikingly, volcanic gases have \delta$^{37}$Cl values that cover a range in excess of 14\permil (Barnes et al., this meeting). The large range is difficult to explain in terms of equilibrium fractionation, which, although calculated to be very large for Cl in different oxidation states, should be less than 2\permil between chloride species (Schauble et al., 2003, GCA). To address the discrepancy between Nature and theory, we have measured Cl isotope fractionation for selected equilibrium and disequilibrium experiments in order to identify mechanisms that might lead to large fractionations. 1) NaCl (s,l) � NaCl (v): NaCl was sealed in an evacuated silica tube and heated at one end, causing vaporization and reprecipitation of NaCl (v) at the cool end of the tube. The fractionation is 0.2\permil at 700\degC (halite- vapor) and 0.7\permil at 800\degC (liquid-vapor), respectively. The larger fractionation at higher temperature may be related to equilibrium fractionation between liquid and gas vs. �stripping' of the solid in the lower T experiments. 2) Sodalite � NaCl(l): Nepheline and excess NaCl were sealed in a Pt crucible at 825\degC for 48 hrs producing sodalite. The measured newly-formed sodalite-NaCl fractionation is -0.2\permil. 3) Volatilization of HCl: Dry inert gas was bubbled through HCl solutions and the vapor was collected in a downstream water trap. There was no fractionation for 12.4M HCl (HCl fuming) � vapor at 25\degC. For a 1 M boiling HCl solution, the HCl-vapor fractionation was ~9\permil. The difference is probably related to the degree of dissociation in the acid, with HCl dissolved in water for the highly acidic solutions, and dissociated H$_{3}$O$^{+}$ and Cl$^{-}$ for lower concentrations. The HCl volatilization experiments are in contrast to earlier vapor-liquid experiments in NaCl-H$_{2}$O system, where fractionation was less than 0.2\permil at 400- 450 \degC (Leibscher et al., Chem. Geol., 2006). The HCl vaporization data provide a mechanism for generating large fractionations under appropriate natural conditions.
V14C-04
Experimental chlorine stable isotope fractionation of perchlorate respiring bacteria
Perchlorate natural occurrences on earth are very limited and seem restricted to extremely arid environments such as nitrate deposits of the Atacama Desert of northern Chile, where perchlorate contents can reach 0.1 to 1%. Anthropogenically sourced perchlorate however is extensively used as a major component of explosives and rocket fuels. Careless disposal of these highly soluble and very stable perchlorates locally led to the contamination of drinking water, now recognised as posing a significant health threat. Recent studies have demonstrated that some microorganisms are able to completely reduce perchlorate to innocuous chloride, and offer a great potential for the bioremediation of contaminated waters. Provided that the isotopic fractionation associated with this reduction is significant, the measurement of the chloride isotopic composition of contaminated water is a powerful tool for monitoring the progress of in-situ remediation. We report here, the characterisation of the isotopic fractionation associated with perchlorate reduction performed by {\it Dechlorosoma suillum} strain PS during 3 culture experiments performed in a batch fermentor (anoxic, $37\deg$�C, pH =7). The basal medium contained acetate as the electron donor and perchlorate as the electron acceptor. When possible, chloride salts were replaced by sulphate salts so as to lower the initial chloride content. The paired chlorine isotopic compositions of chloride and perchlorate in solutions sampled throughout the experiment were measured using the method described in Ader et al. 2001. The fractionation between chloride and perchlorate was calculated independently for each sample, using on the one hand the chloride content and isotopic composition and on the other hand the perchlorate content and isotopic composition. The results show that the fractionation is constant within error throughout the experiment for the 3 experiments with a weighted mean of -14.94$\pm$0.14$\permil$. This value is much lower than the theoretical value expected for an isotopic equilibrium between perchlorate and chloride (-65.8 or -71.3$\permil$; Schauble et al., 2003). Making the assumption that the microbial fractionation reflects that of the rate-limiting reaction step (probably the first electron transfer), and that the theoretical isotopic fractionation for this electron transfer to perchlorate can be deduced from the general relation between the fractionation between perchlorate and the other species depending on their oxidation state, we suggest that the first step of the microbial perchlorate reduction might occur at isotopic equilibrium in the investigated conditions. Ader et al., 2001, Anal. Chem. 73, 4946-4950. Schauble et al., 2003 G.C.A., 67, 3267-3281.
V14C-05
Isotope Fractionation During Microbial Metal Assimilation
The possibility that metal stable isotopes record the influence of microbes on metal geochemical cycling has motivated much recent research on �non-traditional� stable isotopes, particularly Fe. The initial wave of research on biogenic metal isotope effects focused on Fe isotope fractionation during microbially-mediated dissimilatory reduction or oxidation of Fe. Although isotope variations arising from biogenic effects have been reported in laboratory systems it is difficult to ascribe comparable variations in nature to biology because of pervasive and significant abiotic fractionation. As an alternative approach, we are investigating isotope fractionation during microbial assimilation of transition metals. Assimilation occurs because a large number of metals are essential intracellular constituents. Although assimilatory isotope fractionation is not likely to be unique in direction or magnitude compared to other processes, the large number of elements potentially involved greatly broadens the number of elements that can be examined for biogenic isotope effects in materials of interest. This raises the possibility of multi-element isotope �fingerprints� of biological metal processing. In experiments with {\it Azotobacter vinelandii}, a nitrogen-fixing soil bacterium that does not use Fe or other metals in dissimilatory respiration, fractionation of both Fe and Mo isotopes are observed. The two systems exhibit opposite sense fractionation: preferential assimilation of heavy isotopes is observed for Fe, while Mo assimilation favors uptake of light isotopes. Rayleigh-type behavior is seen in both cases; $\alpha$ = 1.0011 and 0.9997, respectively. The Fe isotope results are most readily interpreted in terms of an equilibrium fractionation between inorganic Fe complexes and strongly bound Fe-siderophore complexes that are taken into the cell. In contrast, the Mo isotope results may reflect a kinetic isotope effect. However, it is alternatively possible that Mo isotope fractionation occurs when octahedral Mo-catecholate complexes in solution equilibrate with tetrahedral molybdate complexes, the latter taken into the cell via an ABC transport system. Regardless of the mechanistic details, the data suggest that the coupled isotope systematics of two or more elements may prove useful in studying biological use of metals in recent or ancient sediments.
V14C-06
Chlorophyll-a Photosynthesis and Mg Isotope Fractionation
Mg is the metal center of all the chlorophyll pigments and therefore at the center of the process of photosynthesis. Chlorophyll (Chl) is often used as a biomarker of photosynthesis and is an enormous contributor to the global carbon cycle. Biosynthetic processes fractionate isotopes of light elements and this led us to examine the isotopic composition of Mg in Chl, as another potential biomarker. Here we detail the Mg isotopic composition of Chl-a, extracted from cultures of Synechococcus elongatus, and the culture medium (Black et al., 2006). After Chl extraction, the Mg was liberated from Chl and purified on cation- exchange columns, with a final yield of 100 \pm 5%. $^{26}$Mg/$^{24}$Mg and $^{25}$Mg/$^{24}$Mg, were measured relative to Cambridge 1 and DSM3 standards by a standard-sample-bracketing technique on an MC- ICP-MS (Nu Instruments Ltd). We have measured the average isotopic fractionation of Mg from six samples of Chl-a from early growth phase and 4 samples from late growth phase, 9 samples of the culture medium and the Cambridge 1 Std, all relative to the DSM3 Std. We demonstrate for the first time that there is a clearly resolved depletion in the heavy isotopes of Mg in Chl-a relative to the culture medium (\Delta$^{26}$Mg =-0.61\permil; \Delta$^{25}$Mg =-0.30\permil). The heavy isotope depletion observed may be caused by chelation effects during the biosynthesis of Chl-a. We are now evaluating two hypotheses about the cause of the fractionation. One hypothesis is that the insertion step induces a fractionation via the Mg-chelatase enzyme. The second is that transport into the cell, such as via an ion channel, causes the fractionation. In either case, no difference between Chl-a and Chl-b is anticipated. Experiments and field studies are underway to examine these ideas. References Black, J., Yin, Q.-Z., Casey, W.H., 2006. Geochim. Cosmochim. Acta, 70, 4072-4079.
V14C-07 INVITED
The equilibrium fractionation factor between CaCO$_3$ and Ca$^{2+}$ (aq): Fractionation mechanisms and diagenetic Ca isotope effects
The Ca isotopic compositions ($\delta^{44}$Ca) of 30 high-purity nannofossil ooze and chalk and 7 pore fluid samples from ODP Site 807A (Ontong Java Plateau), combined with reactive transport modeling, are used to determine the equilibrium Ca isotope fractionation factor ($\alpha_{s-f}$) between calcite and dissolved Ca$^{2+}$. The value of $\alpha_{s-f}$ at equilibrium is inferred to be 1.0000 $\pm$ 0.0001, which is significantly different from the value (0.9985) inferred from abiotic calcite precipitation in the laboratory. Since calcite precipitation rates at 807A (constrained by Ca and Sr elemental and isotopic data) are ~10--14 orders of magnitude slower than laboratory rates, and the pore fluids are close to calcite saturation, the Ca isotopic composition of calcite should reflect precipitation at isotopic equilibrium. Consequent modeling of diagenesis produces a maximum shift in $\delta^{44}$Ca of $+$0.15$\permil$ at 807A; however, diagenesis will have a larger impact in sections where sedimentation rates are low, seawater circulates through the sediment pile, or there are prolonged depositional hiatuses. We propose that adsorption and diffusion control Ca isotopic fractionation when precipitation rates and oversaturation levels are high, such as in biogenic calcite formation and in laboratory precipitation experiments. At high rates of precipitation, calcite growth may be limited by diffusion to the crystal surface; this may result in no isotopic fractionation due to the hypothesized equality of the aqueous diffusivities for $^{44}$Ca$^{2+}$ and $^{40}$Ca$^{2+}$. At intermediate precipitation rates, growth is controlled primarily by adsorption, producing a surface layer that has low $^{44}$Ca/$^{40}$Ca via preferential adsorption of $^{40}$Ca$^{2+}$. Rapid precipitation incorporates these layers into the crystal before they equilibrate with the solution, resulting in observed isotope fractionation between crystal and solution. At very low precipitation rates, the adsorbed surface layer is in contact with the bulk solution for a relatively long time. As a result, the surface layer can equilibrate with the bulk solution before it is removed from the reacting system by the deposition of successive layers.
V14C-08
11,10B Isotopic Fractionation Between B(OH)3 and B(OH)4- and an Inorganic Mechanism for their Incorporation Into Calcite
It has recently been established experimentally by Byrne, et al. (2006) and Klochko, et al. (2006) that the equilbrium constant for the isotopic exchange reaction: 10B(OH)3 + 11B(OH)4- = 11B(OH)3 + 10B(OH)4- (1) has a value around 1.027 for seawater at 25$^{circ}$C, for total B concentrations from 0.01 to 0.05 molal. These experimental studies involved essentially the accurate determination of the small pKa difference between the 11B and 10B isotopomers of boric acid. This equilibrium constant value is significantly higher than the traditional value of 1.0194 from Kakihana, et al. (1977). It agrees well with calculated values from Liu and Tossell (2005) but disagrees with a considerably smaller value obtained from spectral studies by Sanchez-Valle, et al. (2005). We will present additional calculations supporting and extending the study of Liu and Tossell (2005) and we will explain the discrepancy with the study of Sanchez-Valle, et al. (2005) and discuss the general unsuitability of methods employing experimental spectral data. We will decompose the free energy change for reaction (1) into its enthalpic and entropic components and will discuss isotopic fractionation between B(OH)3 (aq) and various vapor species. We will also present a model for the incorporation of B(OH)3 or B(OH)4- into calcite by a chemical reaction with HCO3- on the surface of hydrated calcite which produces a B(OH)2CO3- complex, in which the B may be either 3- or 4-coordinate. References Byrne, et al. Deep-Sea Research I (2006) 53, 684-688. Kakihana, et al. Bull. Chem. Soc. Jpn. (1977) 50, 158-163. Klochko, et al. Earth Planet. Sci. Lett. (2006) 248, 276-285. Liu and Tossell Geochim. Cosmochim. Acta. (2005) 69, 3995-4006. Sanchez-Valle, et al. Geochim. Cosmochim. Acta (2005) 69, 4301-4313.