Dissociation constants and speciation in aqueous Li2SO4 and K2SO4 from measurements of electrical conductance to 673 K and 29 MPa

The electrical conductivities of aqueous solutions of Li₂SO₄ and K₂SO₄ have been measured at 523-673 K at 20-29 MPa in dilute solutions for molalities up to 2 × 10⁻² mol kg⁻¹. These conductivities have been fitted to the conductance equation of Turq, Blum, Bernard, and Kunz with a consensus mixing rule and mean spherical approximation activity coefficients. In the temperature interval 523-653 K, where the dielectric constant, ε, is greater than 14, the electrical conductance data can be fitted by a solution model which includes ion association to form , , and , where M is Li or K. The adjustable parameters of this model are the first and second dissociation constants of the M₂SO₄. For the 673 K and 300 kg m⁻³ state point where the Coulomb interactions are the strongest (dielectric constant, ε = 5), models with more extensive association give good fits to the data. In the case of the Li₂SO₄ model, including the multi-ion associate, , gave an extremely good fit to the conductance data.     

Free energy of formation for green rust sodium sulphate (NaFe6Fe3(OH)18(SO4)2(s))

In a recent study, sulphate-bearing green rust (GR_SO4) was shown to incorporate Na⁺ in its structure (NaFe^II₆Fe^III₃(OH)₁₈(SO₄)_2(s); GR_Na,SO4). The compound was synthesised by aerial oxidation of Fe(OH)_2(s) in the presence of NaOH. This paper reports on its free energy of formation .Freshly synthesised GR_Na,SO4 was titrated with 0.5 M H₂SO₄ in an inert atmosphere at 25 °C, producing dissolved Fe²⁺ and magnetite or goethite. Solution concentrations, PHREEQC and the MINTEQ database were used to calculate reaction constants for the reactions:

     

Uranophane dissolution and growth in CaCl2-SiO2(aq) test solutions

The dissolution and growth of uranophane [Ca(UO₂)₂(SiO₃OH)₂·5H₂O] have been examined in Ca- and Si-rich test solutions at low temperatures (20.5 ± 2.0 °C) and near-neutral pH (∼6.0). Uranium-bearing experimental solutions undersaturated and supersaturated with uranophane were prepared in matrices of ∼10⁻² M CaCl₂ and ∼10⁻³ M SiO₂(aq). The experimental solutions were reacted with synthetic uranophane and analyzed periodically over 10 weeks. Interpretation of the aqueous solution data permitted extraction of a solubility constant for the uranophane dissolution reaction and standard state Gibbs free energy of formation for uranophane ( kJ mol⁻¹).     

Dawsonite synthesis and reevaluation of its thermodynamic properties from solubility measurements: Implications for mineral trapping of CO2

Over the last decade, a significant research effort has focused on determining the feasibility of sequestering large amounts of CO₂ in deep, permeable geologic formations to reduce carbon dioxide emissions to the atmosphere. Most models indicate that injection of CO₂ into deep sedimentary formations will lead to the formation of various carbonate minerals, including the common phases calcite (CaCO₃), dolomite (CaMg(CO₃)₂), magnesite (MgCO₃), siderite (FeCO₃), as well as the far less common mineral, dawsonite (NaAlCO₃(OH)₂). Nevertheless, the equilibrium and kinetics that control the precipitation of stable carbonate minerals are poorly understood and few experiments have been performed to validate computer codes that model CO₂ sequestration.In order to reduce this uncertainty we measured the solubility of synthetic dawsonite according to the equilibrium: , from under- and oversaturated solutions at 50-200 °C in basic media at 1.0 mol · kg⁻¹ NaCl. The solubility products (Q_s) obtained were extrapolated to infinite dilution to obtain the solubility constants (. Combining the fit of these values and fixing  at 25 °C, which was derived from the calorimetric data of Ferrante et al. [Ferrante, M.J., Stuve, J.M., and Richardson, D.W., 1976. Thermodynamic data for synthetic dawsonite. U.S. Bureau of Mines Report Investigation, 8129, Washington, D.C., 13p.], the following thermodynamic parameters for the dissolution of dawsonite were calculated at 25 °C: , and . Subsequently, we were able to derive values for the Gibbs energy of formation (, enthalpy of formation ( and entropy ( of dawsonite. These results are within the combined experimental uncertainties of the values reported by Ferrante et al. (1976). Predominance diagrams are presented for the dawsonite/boehmite and dawsonite/bayerite equilibria at 100 °C in the presence of a saline solution with and without silica-containing minerals.     

Solubility of NaCl in CO2 at high pressure and temperature: First experimental measurements

NaCl solubility in gaseous carbon dioxide has been measured in the pressure range from 30 to 70 MPa at 623 and 673 K. Our originally-designed high pressure apparatus allows in situ sampling of a portion of the fluid phase for chemical analysis. The results indicate that the solubility of NaCl increases with both temperature and pressure, and is about 4-5 orders of magnitude higher than saturated NaCl pressure values at the same temperature conditions (6.02 × 10⁻¹² at 623 K and 1.51 × 10⁻¹⁰ at 673 K). It is also 1-2 orders of magnitude greater than predictions according to the Equation of State of the ternary H₂O-CO₂-NaCl system by Duan, Moeller and Weare [Duan, Z., Moller, N., and Weare, J. H. (1995) Equation of state for the NaCl-H₂O-CO₂ system: prediction of phase equilibria and volumetric properties. Geochim. Cosmochim. Acta59, 2869] and has the opposite pressure dependence. The activity values of NaCl in the vapor phase, calculated from the experiments (with pure molten NaCl as a standard state in the vapor), have been fitted to the Darken Quadratic Formalism: , where, x_NaCl,v is mole the fraction of NaCl in the vapor phase, , , where P is the pressure in MPa and T the absolute temperature. Caution should be exerted while extrapolating this empirical equation far beyond the experimental P-T-compositional range.     

Coupled phase and aqueous species equilibrium of the H2O-CO2-NaCl-CaCO3 system from 0 to 250 °C, 1 to 1000 bar with NaCl concentrations up to saturation of halite

A model is developed for the calculation of coupled phase and aqueous species equilibrium in the H₂O-CO₂-NaCl-CaCO₃ system from 0 to 250 °C, 1 to 1000 bar with NaCl concentrations up to saturation of halite. The vapor-liquid-solid (calcite, halite) equilibrium together with the chemical equilibrium of H⁺, Na⁺, Ca²⁺, , Ca(OH)⁺, OH⁻, Cl⁻, , , CO_2(aq) and CaCO_3(aq) in the aqueous liquid phase as a function of temperature, pressure, NaCl concentrations, CO_2(aq) concentrations can be calculated, with accuracy close to those of experiments in the stated T-P-m range, hence calcite solubility, CO₂ gas solubility, alkalinity and pH values can be accurately calculated. The merit and advantage of this model is its predictability, the model was generally not constructed by fitting experimental data.One of the focuses of this study is to predict calcite solubility, with accuracy consistent with the works in previous experimental studies. The resulted model reproduces the following: (1) as temperature increases, the calcite solubility decreases. For example, when temperature increases from 273 to 373 K, calcite solubility decreases by about 50%; (2) with the increase of pressure, calcite solubility increases. For example, at 373 K changing pressure from 10 to 500 bar may increase calcite solubility by as much as 30%; (3) dissolved CO₂ can increase calcite solubility substantially; (4) increasing concentration of NaCl up to 2 m will increase calcite solubility, but further increasing NaCl solubility beyond 2 m will decrease its solubility.The functionality of pH value, alkalinity, CO₂ gas solubility, and the concentrations of many aqueous species with temperature, pressure and NaCl_(aq) concentrations can be found from the application of this model. Online calculation is made available on www.geochem-model.org/models/h2o_co2_nacl_caco3/calc.php.     

Calcite dissolution kinetics in the system CaCO3-H2O-CO2 at high undersaturation

Dissolution rates of limestone covered by a water film open to a CO₂-containing atmosphere are controlled by the chemical composition of the CaCO₃-H₂O-CO₂ solution at the water-mineral interface. This composition is determined by the Ca²⁺-concentration at this boundary, conversion of CO₂ into H⁺ and in the solution, and by diffusional mass transport of the dissolved species from and towards the water-limestone interface. A system of coupled diffusion-reaction equations for Ca²⁺, , and CO₂ is derived. The Ca²⁺ flux rates at the surface of the mineral are defined by the PWP-empirical rate law. These flux rates by the rules of stoichiometry must be equal to the flux rates of CO₂ across the air-water interface. In the solution, CO₂ is converted into H⁺ and . At low water-film thickness this reaction becomes rate limiting. The time dependent diffusion-reaction equations are solved for free drift dissolution by a finite-difference scheme, to obtain the dissolution rate of calcite as a function of the average calcium concentration in the water film. Dissolution rates are obtained for high undersaturation. The results reveal two regimes of linear dissolution kinetics, which can be described by a rate law F = α_i(m_ic_eq − c), where c is the calcium concentration in the water film, c_eq the equilibrium concentration with respect to calcite. For index i = 0, a fast rate law, which here is reported for the first time, is found with α₀ = 3 × 10⁻⁶ m s⁻¹ and m₀ = 0.3. For c > m₀c_eq, a slow rate law is valid with α₁ = 3 × 10⁻⁷ m  s⁻¹ and m₁ = 1, which confirms earlier work. The numbers given above are valid for film thickness of several tenths of a millimetre and at 20 °C. These rates are proven experimentally, using a flat inclined limestone plate covered by a laminar flowing water film injected at an input point with known flow rate Q and calcium concentration. From the concentration measured after flow distance x the dissolution rates are determined. These experiments have been performed at a carbon-dioxide pressure of 0.00035 atm and also of 0.01 atm. The results are in good agreement to the theoretical predictions.     

A thermodynamic model for the prediction of phase equilibria and speciation in the H2O-CO2-NaCl-CaCO3-CaSO4 system from 0 to 250 °C, 1 to 1000 bar with NaCl concentrations up to halite saturation

A thermodynamic model is developed for the calculation of both phase and speciation equilibrium in the H₂O-CO₂-NaCl-CaCO₃-CaSO₄ system from 0 to 250 °C, and from 1 to 1000 bar with NaCl concentrations up to the saturation of halite. The vapor-liquid-solid (calcite, gypsum, anhydrite and halite) equilibrium together with the chemical equilibrium of H⁺,Na⁺,Ca²⁺, , , and CaSO_4(aq) in the aqueous liquid phase as a function of temperature, pressure and salt concentrations can be calculated with accuracy close to the experimental results.Based on this model validated from experimental data, it can be seen that temperature, pressure and salinity all have significant effects on pH, alkalinity and speciations of aqueous solutions and on the solubility of calcite, halite, anhydrite and gypsum. The solubility of anhydrite and gypsum will decrease as temperature increases (e.g. the solubility will decrease by 90% from 360 K to 460 K). The increase of pressure may increase the solubility of sulphate minerals (e.g. gypsum solubility increases by about 20-40% from vapor pressure to 600 bar). Addition of NaCl to the solution may increase mineral solubility up to about 3 molality of NaCl, adding more NaCl beyond that may slightly decrease its solubility. Dissolved CO₂ in solution may decrease the solubility of minerals. The influence of dissolved calcite on the solubility of gypsum and anhydrite can be ignored, but dissolved gypsum or anhydrite has a big influence on the calcite solubility. Online calculation is made available on www.geochem-model.org/model.     

The solution behavior of H2O in peralkaline aluminosilicate melts at high pressure with implications for properties of hydrous melts

Solubility and solution mechanisms of H₂O in depolymerized melts in the system Na₂O-Al₂O₃-SiO₂ were deduced from spectroscopic data of glasses quenched from melts at 1100 °C at 0.8-2.0 GPa. Data were obtained along a join with fixed nominal NBO/T = 0.5 of the anhydrous materials [Na₂Si₄O₉-Na₂(NaAl)₄O₉] with Al/(Al+Si) = 0.00-0.25. The H₂O solubility was fitted to the expression, X_H2O=0.20+0.0020f_H2O-0.7X_Al+0.9(X_Al)², where X_H2O is the mole fraction of H₂O (calculated with O = 1), f_H2O the fugacity of H₂O, and X_Al = Al/(Al+Si). Partial molar volume of H₂O in the melts, , calculated from the H₂O-solulbility data assuming ideal mixing of melt-H₂O solutions, is 12.5 cm³/mol for Al-free melts and decreases linearly to 8.9 cm³/mol for melts with Al/(Al+Si) ∼ 0.25. However, if recent suggestion that is composition-independent is applied to constrain activity-composition relations of the hydrous melts, the activity coefficient of H₂O, , increases with Al/(Al+Si).Solution mechanisms of H₂O were obtained by combining Raman and ²⁹Si NMR spectroscopic data. Degree of melt depolymerization, NBO/T, increases with H₂O content. The rate of NBO/T-change with H₂O is negatively correlated with H₂O and positively correlated with Al/(Al+Si). The main depolymerization reaction involves breakage of oxygen bridges in Q⁴-species to form Q² species. Steric hindrance appears to restrict bonding of H⁺ with nonbridging oxygen in Q³ species. The presence of Al³⁺ does not affect the water solution mechanisms significantly.     

Thermodynamic properties of anhydrous smectite MX-80, illite IMt-2 and mixed-layer illite-smectite ISCz-1 as determined by calorimetric methods. Part I: Heat capacities, heat contents and entropies

The heat capacities of the anhydrous international reference clay minerals, smectite MX-80, illite IMt-2 and mixed-layer illite-smectite ISCz-1, were measured by low temperature adiabatic calorimetry and differential scanning calorimetry, from 6 to 520 K (at 1 bar). The samples were chemically purified and Na-saturated. Dehydrated clay fractions <2 μm were studied. The structural formulae of the corresponding clay minerals, obtained after subtracting the remaining impurities, are K_0.026Na_0.435Ca_0.010(Si_3.612Al_0.388) (Al_1.593Mg_0.228Ti_0.011)O₁₀(OH)₂ for smectite MX-80, K_0.762Na_0.044(Si_3.387Al_0.613) (Al_1.427Mg_0.241O₁₀(OH)₂ for illite IMt-2 and K_0.530Na_0.135(Si_3.565Al_0.435)(Al_1.709Mg_0.218Ti_0.005)O₁₀(OH)₂for mixed-layer ISCz-1. From the heat capacity values, we determined the molar entropies, standard entropies of formation and heat contents of these minerals. The following values were obtained at 298.15 K and 1 bar:

	(J mol−1 K−1)	S0 (J mol−1 K−1)
Smectite MX-80	326.13 ± 0.10	280.56 ± 0.16
Illite IMt-2	328.21 ± 0.10	295.05 ± 0.17
Mixed-layer ISCz-1	320.79 ± 0.10	281.62 ± 0.15

Full-size table

     

Molecular H2O in armenite, BaCa2Al6Si9O30·2H2O, and epididymite, Na2Be2Si6O15·H2O: Heat capacity, entropy and local-bonding behavior of confined H2O in microporous silicates

Armenite, ideal formula BaCa₂Al₆Si₉O₃₀·2H₂O, and its dehydrated analog BaCa₂Al₆Si₉O₃₀ and epididymite, ideal formula Na₂Be₂Si₆O₁₅·H₂O, and its dehydrated analog Na₂Be₂Si₆O₁₅ were studied by low-temperature relaxation calorimetry between 5 and 300 K to determine the heat capacity, C_p, behavior of their confined H₂O. Differential thermal analysis and thermogravimetry measurements, FTIR spectroscopy, electron microprobe analysis and powder Rietveld refinements were undertaken to characterize the phases and the local environment around the H₂O molecule.The determined structural formula for armenite is Ba_0.88(0.01)Ca_1.99(0.02)Na_0.04(0.01)Al_5.89(0.03)Si_9.12(0.02)O₃₀·2H₂O and for epididymite Na_1.88(0.03)K_0.05(0.004)Na_0.01(0.004)Be_2.02(0.008)Si_6.00(0.01)O₁₅·H₂O. The infrared (IR) spectra give information on the nature of the H₂O molecules in the natural phases via their H₂O stretching and bending vibrations, which in the case of epididymite only could be assigned. The powder X-ray diffraction data show that armenite and its dehydrated analog have similar structures, whereas in the case of epididymite there are structural differences between the natural and dehydrated phases. This is also reflected in the lattice IR mode behavior, as observed for the natural phases and the H₂O-free phases. The standard entropy at 298 K for armenite is S° = 795.7 ± 6.2 J/mol K and its dehydrated analog is S° = 737.0 ± 6.2 J/mol K. For epididymite S° = 425.7 ± 4.1 J/mol K was obtained and its dehydrated analog has S° = 372.5 ± 5.0 J/mol K. The heat capacity and entropy of dehydration at 298 K are Δ = 3.4 J/mol K and ΔS^rxn = 319.1 J/mol K and Δ = −14.3 J/mol K and ΔS^rxn = 135.7 J/mol K for armenite and epididymite, respectively. The H₂O molecules in both phases appear to be ordered. They are held in place via an ion-dipole interaction between the H₂O molecule and a Ca cation in the case of armenite and a Na cation in epididymite and through hydrogen-bonding between the H₂O molecule and oxygen atoms of the respective silicate frameworks. Of the three different H₂O phases ice, liquid water and steam, the C_p behavior of confined H₂O in both armenite and epididymite is most similar to that of ice, but there are differences between the two silicates and from the C_p behavior of ice. Hydrogen-bonding behavior and its relation to the entropy of confined H₂O at 298 K is analyzed for various microporous silicates.The entropy of confined H₂O at 298 K in various silicates increases approximately linearly with increasing average wavenumber of the OH-stretching vibrations. The interpretation is that decreased hydrogen-bonding strength between a H₂O molecule and the silicate framework, as well as weak ion-dipole interactions, results in increased entropy of H₂O. This results in increased amplitudes of external H₂O vibrations, especially translations of the molecule, and they contribute strongly to the entropy of confined H₂O at T < 298 K.     

Comparative carbonate and hydroxide complexation of cations in seawater

Comparative concentrations of carbonate and hydroxide complexes in natural solutions can be expressed in terms of reactions with bicarbonate that have no explicit pH dependence (). Stability constants for this reaction with n = 1 were determined using conventional formation constant data expressed in terms of hydroxide and carbonate. Available data indicate that stability constants appropriate to seawater at 25 °C expressed in the form are on the order of 10^4.2 for a wide range of cations (M^z+) with z = +1, +2 and +3. Φ₁ is sufficiently large that species appear to substantially dominate MOH^z−1 species in seawater. Evaluations of comparative stepwise carbonate and hydroxide stability constant behavior leading to the formation of n = 2 and n = 3 complexes suggest that carbonate complexes generally dominate hydroxide complexes in seawater, even for cations whose inorganic speciation schemes in seawater are currently presumed to be strongly dominated by hydrolyzed forms (). Calculated stability constants, and , indicate that the importance of carbonate complexation is sufficiently large that carbonate and hydroxide complexes would be generally comparable even if calculated Φ₂ and Φ₃ values are overestimated by two or more orders of magnitude. Inclusion of mixed ligand species in carbonate-hydroxide speciation models allows cation complexation intensities (M_T/[M^z+]) to be expressed in the following form:

     

Trace element partitioning between ilmenite, armalcolite and anhydrous silicate melt: Implications for the formation of lunar high-Ti mare basalts

We performed a series of experiments at high pressures and temperatures to determine the partitioning of a wide range of trace elements between ilmenite (Ilm), armalcolite (Arm) and anhydrous lunar silicate melt, to constrain geochemical models of the formation of titanium-rich melts in the Moon. Experiments were performed in graphite-lined platinum capsules at pressures and temperatures ranging from 1.1 to 2.3 GPa and 1300-1400 °C using a synthetic Ti-enriched Apollo ‘black glass’ composition in the CaO-FeO-MgO-Al₂O₃-TiO₂-SiO₂ system. Ilmenite-melt and armalcolite-melt partition coefficients (D) show highly incompatible values for the rare earth elements (REE) with the light REE more incompatible compared to the heavy REE ( 0.0020 ± 0.0010 to 0.069 ± 0.010 for ilmenite; 0.0048 ± 0.0023 to 0.041 ± 0.008 for armalcolite). D values for the high field strength elements vary from highly incompatible for Th, U and to a lesser extent W (for ilmenite: 0.0013 ± 0.0008, 0.0035 ± 0.0015 and 0.039 ± 0.005, and for armalcolite 0.008 ± 0.003, 0.0048 ± 0.0022 and 0.062 ± 0.03), to mildly incompatible for Nb, Ta, Zr, and Hf (e.g. 0.28 ± 0.05 and : 0.76 ± 0.07). Both minerals fractionate the high field strength elements with D_Ta/D_Nb and D_Hf/D_Zr between 1.3 and 1.6 for ilmenite and 1.3 and 1.4 for armalcolite. Armalcolite is slightly more efficient at fractionating Hf from W during lunar magma ocean crystallisation, with D_Hf/D_W = 12-13 compared to 6.7-7.5 for ilmenite. The transition metals vary from mildly incompatible to compatible, with the highest compatibilities for Cr in ilmenite (D ∼ 7.5) and V in armalcolite (D ∼ 8.1). D values show no clear variation with pressure in the small range covered.Crystal lattice strain modelling of D values for di-, tri- and tetravalent trace elements shows that in ilmenite, divalent elements prefer to substitute for Fe while armalcolite data suggest REE replacing Mg. Tetravalent cations appear to preferentially substitute for Ti in both minerals, with the exception of Th and U that likely substitute for the larger Fe or Mg cations. Crystal lattice strain modelling is also used to identify and correct for very small (∼0.3 wt.%) melt contamination of trace element concentration determinations in crystals.Our results are used to model the Lu-Hf-Ti concentrations of lunar high-Ti mare basalts. The combination of their subchondritic Lu/Hf ratios and high TiO₂ contents requires preferential dissolution of ilmenite or armalcolite from late-stage, lunar magma ocean cumulates into low-Ti partial melts of deeper pyroxene-rich cumulates.     

Modeling the effects of bond environment on equilibrium iron isotope fractionation in ferric aquo-chloro complexes

Four or five sets of ab initio models, including Unrestricted Hartree Fock (UHF) and hybrid Density Functional Theory (DFT) are calculated for each species in a series of aqueous ferric aquo-chloro complexes: , , , FeCl₃(H₂O)₃, FeCl₃(H₂O)₂, , FeCl₅H₂O²⁻, , ) in order to determine the relative isotopic fractionation among the complexes, to compare the results of different models for the same complexes, to examine factors that influence the magnitude of the isotopic fractionation, and to compare bond-partner-driven fractionation with redox-driven fractionation.Relative to , all models show a nearly linear decrease in ⁵⁶Fe/⁵⁴Fe as the number of Cl⁻ ions per Fe³⁺ ion increases, with slopes of −0.8‰ to −1.0‰ per Cl⁻ at 20 °C. At 20 °C, 1000 ln β (β = ⁵⁶Fe/⁵⁴Fe reduced partition function ratio relative to a dissociated Fe atom) values range from 8.93‰ to 9.73‰ for , 8.04-9.12‰ for , 7.61-8.73‰ for , 7.14-8.25‰ for , and 3.09-4.41‰ for . The fractionation between and ranges from 1.5‰ to 2.6‰, depending on the model; this is comparable in magnitude to fractionation effects due to Fe³⁺/Fe²⁺ redox reactions. β values from the UHF models are consistently higher than those from the hybrid DFT models.Isotopic fractionation is shown to be sensitive to differences in ligand bond stiffness (above), coordination number, bond length, and the frequency of the asymmetric Fe-X stretching vibrational mode, as predicted by previous theoretical studies. Complexes with smaller coordination numbers have higher 1000 ln β (7.46‰, 5.25‰, and 3.48‰ for , ,, respectively, from the B3LYP/6-31G(d) model). Species with the same number of chlorides but fewer waters also show the effect of coordination number on 1000 ln β: (7.46‰ vs. 7.05‰ for FeCl₃(H₂O)₂ vs. FeCl₃(H₂O)₃ and 5.25‰ vs. 4.94‰ for vs. FeCl₅H₂O²⁻ with the B3LYP/6-31G(d) model). As more Fe-Cl bonds substitute for Fe-OH₂ bonds (with a resulting decrease in β), the lengths of the Fe-Cl bonds and the Fe-O bonds increase.Preliminary modeling of shows an Fe³⁺/Fe²⁺ fractionation of 3.2‰ for the B3LYP/6-31G(d) model, in agreement with previous studies. The addition of an explicit outer hydration sphere of 12 H₂O molecules to models of improves agreement with measured vibrational frequencies and bond lengths; 1000 ln β increases by 0.8-1.0‰. An additional hydration sphere around increases 1000 ln β by only 0.1‰.Isotopic fractionations predicted for this simple system imply that ligands present in an aqueous iron environment are potentially important drivers of fractionation, and suggest that significant fractionation effects are likely in other aqueous systems containing sulfides or organic ligands. Fractionation effects due to both speciation and redox must be considered when interpreting iron isotope fractionations in the geological record.     

The activity coefficients of Fe(III) hydroxide complexes in NaCl and NaClO4 solutions

The osmotic coefficients of FeCl₃ at 25 °C from 0.15 to 1.7 m [Rumyantsev et al., Z. Phys. Chem., 218, 1089-1127, 2004] have been used to determine the Pitzer parameters (β⁽⁰⁾, β⁽¹⁾ and C^?) for FeCl₃. Since the differences in the Pitzer coefficients of rare earths in NaCl and NaClO₄ are small, the values of Fe(ClO₄)₃ have been estimated using the differences between La(ClO₄)₃ and LaCl₃. The Pitzer coefficients for FeCl₃ combined with enthalpy and heat capacity data for the rare earths can be used to estimate the activity coefficients of Fe³⁺ in NaCl over a wide range of temperatures (0 to 50 °C) and ionic strength (0 to 6 m).The activity coefficients of Fe³⁺ in NaCl and NaClO₄ solutions have been used to determine the activity coefficients of Fe(OH)²⁺ in these solutions from the measured first hydrolysis constants of Fe³⁺ [Byrne et al., Mar. Chem., 97, 34-48, 2005]. The activity coefficients of , Fe(OH)₃ and from 0 to 50 °C have also been determined from the solubility measurements of Fe(III) in NaCl solutions [Liu and Millero, Geochim. Cosmochim Acta, 63, 3487-3497, 1999]. These activity coefficients have been fitted to the Pitzer equations. These results can be used to estimate the speciation of Fe(III) with OH⁻ in natural waters with high concentrations of NaCl from 0 to 50 °C.     

A general algorithm for the O abundance correction to C/C determinations from CO2 isotopologue measurements, including CO2 characterised by ‘mass-independent’ oxygen isotope distributions

It is widely recognised that a significant limitation to the ultimate precision of carbon stable isotope ratio measurements, as obtained from dual-inlet mass spectrometric measurements of CO₂ isotopologue ion abundances at m/z 44, 45, and 46, is the correction for interference from ¹⁷O-bearing molecular ions. Two long-established, alternative procedures for determining the magnitude of this correction are in widespread use (although only one has IAEA approval); their differences lead to small but potentially significant discrepancies in the magnitude of the resulting correction. Furthermore, neither approach was designed to accommodate oxygen three-isotope distributions which do not conform to terrestrial mass-dependent behaviour. Stratospheric CO₂, for example, contains a strongly ‘mass-independent’ oxygen isotope composition. A new strategy for determining the ¹⁷O-bearing ion correction is presented, for application where the oxygen three-isotope characteristics of the analyte CO₂ are accurately known (or assigned) in terms of the slope λ of the three-isotope fractionation line and the ordinate axis intercept 10³ ln(1 + k) on a 10³ ln(1 + δ¹⁷O) versus 10³ ln(1 + δ¹⁸O) plot. At the heart of the approach is the relationship between ¹⁷R, which is the ¹⁷O/¹⁶O ratio of the sample CO₂, and other assigned or empirically determined parameters needed for the δ¹³C evaluation: