Oxygen and sulfur isotope equilibria in the BaSO4HSO4−H2O system from 110 to 350°C and applications

Oxygen isotope exchange between BaSO₄ and H₂O from 110 to 350°C was studied using 1 m H₂SO₄-1 m NaCl and 1 m NaCl solutions to recrystallize the barite. The slow exchange rate (only 7% exchange after 1 yr at 110°C and 91% exchange after 22 days at 350°C in 1 m NaCl solution) prompted the use of the partial equilibrium technique. However, runs at 300 and 350°C were checked by complete exchange experiments. The temperature calibration curve for the isotope exchange is calculated giving most weight to the high temperature runs where the partial equilibrium technique can be tested. Oxygen isotope fractionation factors (α) in 1 m NaCl solution (110–350°C), assuming a value of 1.0407 for α_CO2H2O at 25°C, are:

.These data, when corrected for ion hydration effects in solution (Truesdell, 1974), give the fractionation factors in pure water:

.In the 1 m H₂SO₄-1 m NaCl runs, sulfur isotope fractionation between HSO^?₄ and BaSO₄ is less than the detection limit of 0.4%. A barite-sulfide geothermometer is obtained by combining HSO^?₄H₂S and sulfide-H₂S calibration data.Barite in the Derbyshire ore field, U.K., appears to have precipitated in isotopic equilibrium with water and sulfur in the ore fluid at temperatures less than 150°C. At the Tui Mine, New Zealand, the barite-water geothermometer indicates temperatures of late stage mineralization in the range 100–200°C. A temperature of 350 ± 20°C is obtained from the barite-pyrite geothermometer at the Yauricocha copper deposit, Peru, and oxygen isotope analyses of the barite are consistent with a magmatic origin for the ore fluids.     

Light hydrocarbons in Red Sea brines and sediments

Light hydrocarbon (C₁-C₃) concentrations in the water from four Red Sea brine basins (Atlantis II, Suakin, Nereus and Valdivia Deeps) and in sediment pore waters from two of these areas (Atlantis II and Suakin Deeps) are reported. The hydrocarbon gases in the Suakin Deep brine (T = ~ 25°C, $\text{Cl}{}^{\mathrm{?}}\text{= ~ 85‰}$, $\text{CH}{}_4\text{=}\text{~ 71}\text{1}$) are apparently of biogenic origin as evidenced by $\text{C}{}_1\text{(C}{}_2\text{+ C}{}_3$) ratios of ~ 1000. Methane concentrations (6–8 μl/l) in Suakin Deep sediments are nearly equal to those in the brine, suggesting sedimentary interstitial waters may be the source of the brine and associated methane.The Atlantis II Deep has two brine layers with significantly different light hydrocarbon concentrations indicating separate sources. The upper brine (T = ~ 50°C, $\text{Cl}{}^{\mathrm{?}}\text{= ~ 73‰}$, CH₄ = ~ 155 μl/l) gas seems to be of biogenic origin [$\text{C}{}_1\text{(C}{}_2\text{+ C}{}_3\text{) = ~1100}$], whereas the lower brine (T = ~ 61°C, $\text{Cl}{}^{\mathrm{?}}\text{= ~ 155‰}$, CH₄ = ~ 120μl/l) gas is apparently of thermogenic origin [$\text{C}{}_1\text{(C}{}_2\text{+ C}{}_3\text{) = ~ 50}$]. The thermogenic gas resulting from thermal cracking of organic matter in the sedimentary column apparently migrates into the basin with the brine, whereas the biogenic gas is produced in situ or at the seawater-brine interface. Methane concentrations in Atlantis II interstitial waters underlying the lower brine are about one half brine concentrations; this difference possibly reflects the known temporal variations of hydrothermal activity in the basin.     

Complex formation in the ternary system Mg(II)-CO2-H2O

The carbonato and hydrogencarbonato complexes of Mg²⁺ were investigated at 25 and 50° in solutions of the constant ClO₄^? molality (3 M) consisting preponderantly of NaClO₄. The experimental data could be explained assuming the following equilibria: Mg²⁺ + CO_2B + H₂O ag MgHCO⁺₃ + H⁺, $\text{log}{}^1\text{β}{}_1\text{= ?7.64}{}_4\text{± 0.01}{}_7\text{(25°), ?7.46}{}_2\text{± 0.01}{}_1\text{(50°)}$, Mg²⁺ + 2 CO_2g + 2 H₂Oag Mg(HCO₃)⁰₂ ± 2 H⁺, $\text{log}{}^1\text{β}{}_2\text{= ?15.00 ± 0.14 (25°)}$, ?15.37 ± 0.39 (50°), Mg²⁺ + CO_2g + H₂Oag MgCO⁰₃ + 2 H⁺, $\text{log}{}^1\text{k}{}_1\text{= ?15.64 ± 0.06 (25°)}$,?15.23 ± 0.02 (50°), with the assumption γ_MgCO3⁰ = γ_Mg(HCO3)⁰2, ΔG⁰(I = 0) for the reaction MgCO⁰₃ + CO_2g + H₂O = Mg(HCO₃)⁰₂ was estimated to be ?3.9₁ ± 0.8₆ and 0.6 ± 2.4 kJ/mol at 25 and 50°C, respectively. The abundance of carbonate linked Mg(II) species in fresh water systems is discussed.     

Growth pattern and 13C12C isotope fractionation of Cyanidium caldarium and hot spring algal mats

A study was undertaken with the thermophilic green alga Cyanidium caldarium which grows optimally at low pH and high concentrations of CO₂. Carbon-isotope fractionation was not found to be a simple linear function of temperature. Maximum enrichment of ¹²C in cellular material occurred under optimum growth conditions (at approximately pH 2 and at temperatures between 40–50°C in a CO₂ atmosphere). A maximum measured fractionation of ?24‰ may account for low values ($\text{δ}{}^{13}\text{C}\text{< ?30‰}{}_{\mathrm{<mtext>PDB</mtext>}}$) in Precambrian kerogen presumably derived from algal mats.     

Carbonate equilibria in hydrothermal systems: First ionization of carbonic acid in NaCl media to 300°C

The ionization quotients of aqueous carbon dioxide (carbonic acid) have been precisely determined in NaCl media to 5 m and from 50° to 300°C using potentiometric apparatus previously developed at Oak Ridge National Laboratory. The pressure coefficient was also determined to 250°C in the same media. These results have been combined with selected information in the literature and modeled in two ways to arrive at the best fits and to derive the thermodynamic parameters for the ionization reaction, including the equilibrium constant, activity coefficient quotients, and pressure coefficients. The variation with temperature of the two fundamental quantities $\text{Δ}\text{V}\text{?}{}^{\mathrm{o}}$ and $\text{Δ}\text{C}\text{?}{}^{\mathrm{o}}{}_{\mathrm{p}}$ were examined along the saturation vapor pressure curve and at constant density. The results demonstrated again that for reactions with minimal electrostriction changes the magnitudes and variations of $\text{Δ}\text{C}\text{?}{}^{\mathrm{o}}{}_{\mathrm{p}}$ and $\text{Δ}\text{V}\text{?}{}^{\mathrm{o}}$ with temperature are small and, in addition, $\text{Δ}\text{C}\text{?}{}_{\mathrm{p}}$ and $\text{Δ}\text{V}\text{?}$ are approximately independent of salt concentration.The results have also been applied to an examination of the solubility of calcite as a function of pH (in a given NaCl medium) for the neutral to acidic region both for systems with fixed CO₂ pressure and systems where the calcium ion concentration equals the concentration of carbon. The pH of saturated solutions of calcite with P_CO2 of 12 bars increases from 5.1 to 5.5 between 100° and 300°C.     

Experimental hydrogen isotope studies—I. Systematics of hydrogen isotope fractionation in the systems epidote-H2O,zoisite-H2O and AlO(OH)-H2O

$\text{H}\text{D}$ Fractionation factors between epidote minerals and water, and between the AlO(OH) dimorphs boehmite and diaspore and water, have been determined between 150 and 650°C. Small water mineral ratios were used to minimise the effect of incongruent dissolution of epidote minerals. Waters were extracted and analysed directly by puncturing capsules under vacuum. Hydrogen diffusion effects were eliminated by using thick-walled capsules.$\text{H}\text{D}$ Exchange rates are very fast between epidote and water (and between boehmite and water), complete exchange taking only minutes above 450°C but several months at 250°C. Exchange between zoisite and water (and between diaspore and water) is very much slower, and an interpolation method was necessary to determine fractionation factors at 450 and below.For the temperature range 300–650°C, the $\text{H}\text{D}$ equilibrium fractionation factor (α^e) between epidote and water is independent of temperature and Fe content of the epidote, and is given by 1000 In α_epidote-H2O^e = ?35.9 ± 2.5, while below 300°C 1000 In , with a ‘cross-over’ estimated to occur at around 185°C. By contrast, zoisite-water fractionations fit the relationship 1000 In .All studied minerals have hydrogen bonding. Fractionations are consistent with the general relationship: the shorter the O-H -- O bridge, the more depleted is the mineral in D.On account of rapid exchange rates, natural epidotes probably acquired their H-isotope compositions at or below 200°C, where fractionations are near or above 0%.; this is in accord with the observation that natural epidotes tend to concentrate D relative to other coexisting hydrous minerals.     

Sodium and potassium coprecipitation in aragonite

The coprecipitation of Na and K was experimentally investigated in aragonite. The distribution functions were determined at pH 6.8 and 8.8 over aqueous Na and K concentrations of between 5 × 10^?4and 2.0 M and temperatures of between 25 and 75°C.The mole fractions of Na and K in aragonite are related to the aqueous ratios of Na and Ca by a function of the form

$\text{log}\text{X}{}_{\mathrm{Na}}{}_2\text{CO}{}_3\text{,K}{}_2\text{CO}{}_3\text{= C}{}_0\text{+ C}{}_1\text{log}\text{a}{}^2{}_{\mathrm{Na}}\text{? ,K}{}^{\mathrm{?}}\text{a}{}_{\mathrm{Ca}}{}^{\mathrm{2+}}$

where C₀ and C₁ are constants at a given temperature. This equation was derived by a statistical model assuming a heterogeneous energy distribution for the sites of incorporation. The independence of the coprecipitation process from aqueous anion activities suggests that carbonate is the only anionic species in the solid solution.     

The solubilities of calcite,aragonite and vaterite in CO2-H2O solutions between 0 and 90°C,and an evaluation of the aqueous model for the system CaCO3-CO2-H2O

Calculations based on approximately 350 new measurements (Ca_T-PCO₂) of the solubilities of calcite, aragonite and vaterite in CO₂-H₂O solutions between 0 and 90°C indicate the following values for the log of the equilibrium constants K_C, K_A, and K_V respectively, for the reaction CaCO₃(s) = Ca²⁺ + CO^2?₃: $\text{Log K}{}_{\mathrm{C}}\text{= ?171.9065 ? 0.077993T +}\text{2839.319}\text{T}\text{+ 71.595 log T}$$\text{Log K}{}_{\mathrm{A}}\text{= ?171.9773 ? 0.077993T +}\text{2903.293}\text{T}\text{+71.595 log T}$$\text{Log K}{}_{\mathrm{V}}\text{= ?172.1295 ? 0.077993T +}\text{3074.688}\text{T}\text{+ 71.595 log T}$ where T is in ^oK. At 25°C the logarithms of the equilibrium constants are ?8.480 ± 0.020, ?8.336 ± 0.020 and ?7.913 ± 0.020 for calcite, aragonite and vaterite, respectively.The equilibrium constants are internally consistent with an aqueous model that includes the CaHCO⁺₃ and CaCO⁰₃ ion pairs, revised analytical expressions for CO₂-H₂O equilibria, and extended Debye-Hückel individual ion activity coefficients. Using this aqueous model, the equilibrium constant of aragonite shows no PCO₂-dependence if the CaHCO⁺₃ association constant is between 0 and 90°C, corresponding to the value logK_Cahco⁺3 = 1.11 ± 0.07 at 25°C. The CaCO⁰₃ association constant was measured potentiometrically to be between 5 and 80°C, yielding logK_CaCO⁰3 = 3.22 ± 0.14 at 25°C.The CO₂-H₂O equilibria have been critically evaluated and new empirical expressions for the temperature dependence of K_H, K₁ and K₂ are $\text{log K}{}_{\mathrm{H}}\text{= 108.3865 + 0.01985076T ?}\text{6919.53}\text{T}\text{? 40.45154 log T +}\text{669365.}\text{T}{}^2$, $\text{log K}{}_1\text{= ?356.3094 ? 0.06091964T +}\text{21834.37}\text{T}\text{+ 126.8339 log T —}\text{1684915.}\text{T}{}^2$ and logK₂ = ?107.8871 ? 0.03252849T + 5151.79/T + 38.92561 logT ? 563713.9/T² which may be used to at least 250°C. These expressions hold for 1 atm. total pressure between 0 and 100°C and follow the vapor pressure curve of water at higher temperatures.Extensive measurements of the pH of Ca-HCO₃ solutions at 25°C and 0.956 atm PCO₂ using different compositions of the reference electrode filling solution show that measured differences in pH are closely approximated by differences in liquid-junction potential as calculated by the Henderson equation. Liquid-junction corrected pH measurements agree with the calculated pH within 0.003-0.011 pH.Earlier arguments suggesting that the CaHCO⁺₃ ion pair should not be included in the CaCO₃-CO₂-H₂O aqueous model were based on less accurate calcite solubility data. The CaHCO⁺₃ ion pair must be included in the aqueous model to account for the observed PCO₂-dependence of aragonite solubility between 317 ppm CO₂ and 100% CO₂.Previous literature on the solubility of CaCO₃ polymorphs have been critically evaluated using the aqueous model and the results are compared.     

Rutile solubility and titanium coordination in silicate melts

The solubility of rutile has been determined in a series of compositions in the K₂O-Al₂O₃-SiO₂ system ($\text{K}{}^1\text{=}\text{K}{}_2\text{O}\text{(K}{}_2\text{O}$ + Al₂O₃) = 0.38–0.90), and the CaO-Al₂O₃-SiO₂ system ($\text{C}{}^1\text{=}\text{CaO}\text{(CaO + Al}{}_2\text{O}{}_3\text{)}\text{= 0.47–0.59}$). Isothermal results in the KAS system at 1325°C, 1400°C, and 1475°C show rutile solubility to be a strong function of the $\text{K}{}^1$ ratio. For example, at 1475°C the amount of TiO₂ required for rutile saturation varies from 9.5 wt% ($\text{K}{}^1\text{= 0.38}$) to 11.5 wt% ($\text{K}{}^1\text{= 0.48}$) to 41.2 wt% ($\text{K}{}^1\text{= 0.90}$). In the CAS system at 1475°C, rutile solubility is not a strong function of $\text{C}{}^1$. The amount of TiO₂ required for saturation varies from 14 wt% ($\text{C}{}^1\text{= 0.48}$) to 16.2 wt% ($\text{C}{}^1\text{= 0.59}$).The solubility changes in KAS melts are interpreted to be due to the formation of strong complexes between Ti and K⁺ in excess of that needed to charge balance Al³⁺. The suggested stoichiometry of this complex is K₂Ti₂O₅ or K₂Ti₃O₇. In CAS melts, the data suggest that Ca²⁺ in excess of A1³⁺ is not as effective at complexing with Ti as is K⁺. The greater solubility of rutile in CAS melts when $\text{C}{}^1$ is less than 0.54 compared to KAS melts of equal $\text{K}{}^1$ ratio results primarily from competition between Ti and Al for complexing cations (Ca vs. K).TiK_β x-ray emission spectra of KAS glasses ($\text{K}{}^1\text{= 0.43–0.60}$) with 7 mole% added TiO₂, rutile, and Ba₂TiO₄, demonstrate that the average Ti-O bond length in these glasses is equal to that of rutile rather than Ba₂TiO₄, implying that Ti in these compositions is 6-fold rather than 4-fold coordinated. Re-examination of published spectroscopic data in light of these results and the solubility data, suggests that the 6-fold coordination polyhedron of Ti is highly distorted, with at least one Ti-O bond grossly undersatisfied in terms of Pauling's rules.     

Dissociation constants for alkali earth and sodium borate ion pairs from 10 to 50°C

Potentiometric measurements in dilute sodium borate solutions with added alkali earth chlordie salts yield the following expressions for the dissociation constants of alkali earth borate ion pairs from 10 to 50°C:

$\text{p}\text{K(}\text{MgH}{}_2\text{BO}{}_3{}^{\mathrm{+}}\text{=1.266+0.001204 T}$

$\text{p}\text{K(}\text{CaH}{}_2\text{BO}{}_3{}^{\mathrm{+}}\text{=1.154+0.002170 T}$

$\text{p}\text{K(}\text{SrH}{}_2\text{BO}{}_3{}^{\mathrm{+}}\text{=1.033+0.001738 T}$

$\text{p}\text{K(}\text{BaH}{}_2\text{BO}{}_3{}^{\mathrm{+}}\text{=1.942+0.001850 T}$

where T is in °K. Enthalpies for the dissociation reactions at 25°C are less than 1 kcal./mole for all the alkali earth borate ion pairs.Values for pK(NaH₂BO₃°) from 5 to 55°C computed from the experimental data of Owen and King are in good agreement with those determined potentiometrically. The average value from both methods is 0.22 ± 0.1 at 25°C.Application to seawater of computed pK's for MgH₂BO₃⁺, CaH₂BO₃⁺ and NaH₂BO₃⁰ yields an apparent dissociation constant for boric acid of 8.73 vs. 8.70 measured by Lyman, 8.68 by Buch and 8.73 by Byrne and Kester.     

Revised values for the Gibbs free energy of formation of [Al(OH)4 aq−], diaspore,boehmite and bayerite at 298.15 K and 1 bar,the thermodynamic properties of kaolinite to 800 K and 1 bar,and the heats of solution of several gibbsite samples

Solution calorimetric measurements compared with solubility determinations from the literature for the same samples of gibbsite have provided a direct thermochemical cycle through which the Gibbs free energy of formation of [Al(OH)_{4 aq}^?] can be determined. The Gibbs free energy of formation of [Al(OH)_{4 aq}^?] at 298.15 K is ?1305 ± 1 kJ/mol. These heat-of-solution results show no significant difference in the thermodynamic properties of gibbsite particles in the range from 50 to 0.05 μm.The Gibbs free energies of formation at 298.15 K and 1 bar pressure of diaspore, boehmite and bayerite are ?9210 ± 5.0, ?918.4 ± 2.1 and ?1153 ± 2 kJ/mol based upon the Gibbs free energy of [A1(OH)_{4 aq}^?] calculated in this paper and the acceptance of ?1582.2 ± 1.3 and ?1154.9 ± 1.2 kJ/mol for the Gibbs free energy of formation of corundum and gibbsite, respectively.Values for the Gibbs free energy formation of [Al(OH)_{2 aq}⁺] and [AlO_{2 aq}^?] were also calculated as ?914.2 ± 2.1 and ?830.9 ± 2.1 kJ/mol, respectively. The use of [AlC_{2 aq}^?] as a chemical species is discouraged.A revised Gibbs free energy of formation for [H₄SiO_4aq⁰] was recalculated from calorimetric data yielding a value of ?1307.5 ± 1.7 kJ/mol which is in good agreement with the results obtained from several solubility studies.Smoothed values for the thermodynamic functions C_P⁰, ($\text{H}{}_{\mathrm{T}}{}^0\text{- H}{}_{298}{}^0\text{)}\text{T}$, $\text{(}\text{G}{}_{\mathrm{T}}{}^0\text{- H}{}_{298}{}^0\text{)}\text{T}$, S_T⁰ - S₀⁰, $\text{ΔH}{}_{\mathrm{?,298}}{}^0$ kaolinite are listed at integral temperatures between 298.15 and 800 K. The heat capacity of kaolinite at temperatures between 250 and 800 K may be calculated from the following equation: C_P⁰ = 1430.26 ? 0.78850 T + 3.0340 × 10^?4T² ?1.85158 × 10^?4T²$\text{1}\text{2}\text{+ 8.3341 × 10}{}^6\text{T}{}^{\mathrm{?2}}$.The thermodynamic properties of most of the geologically important Al-bearing phases have been referenced to the same reference state for Al, namely gibbsite.     

Etude des inclusions fluides du système N2-CO2 de dolomites et de quartz de Tunisie septentrionale. Données de la microcryoscopie et de l'analyse à la microsonde à effet Raman

Optical and analytical studies were performed on 400 N₂ + CO₂ gas bearing inclusions in dolomites and quartz from Triassic outcrops in northern Tunisia. Other fluids present include brines (NaCl and KCl bearing inclusions) and rare liquid hydrocarbons. At the time of trapping, such fluids were heterogeneous gas + brine mixtures. In hydrocarbon free inclusions the $\text{N}{}_2\text{(N}{}_2\text{+ CO}{}_2\text{)}$ mole ratio was determined using two different non-destructive and punctual techniques: Raman microprobe analysis, and optical estimation of the volume ratios of the different phases selected at low temperatures. In the observed range of compositions, the two methods agree reasonably well.The N₂ + CO₂ inclusions are divided into three classes of composition: (a) $\text{N}{}_2\text{(N}{}_2\text{+ CO}{}_2\text{)}\text{> 0,57}$: Liquid nitrogen is always visible at very low temperature and homogenisation occurs in the range ?151°C to ? 147°C (nitrogen critical temperature) dry ice (solid CO₂) sublimates between ?75°C and ?60°C; (b) $\text{0,20 <}\text{N}{}_2\text{(N}{}_2\text{+ CO}{}_2\text{)}\text{? 0,57}$: liquid nitrogen is visible at very low temperature but dry ice melts on heating; liquid and gas CO₂ homogenise to liquid phase between ?51°C to ?22°C; (c) $\text{N}{}_2\text{(N}{}_2\text{+ CO}{}_2\text{)}\text{? 0,20}$: liquid nitrogen is not visible even at very low temperature (?195°C) and liquid and gas CO₂ homogenise to liquid phase between ?22°C and ?15°C. The observed phases changes are used to propose a preliminary phase diagram for the system CO₂-N₂ at low temperatures.Assuming additivity of partial pressures, isochores for the CO₂-N₂ inclusions have been computed. The intersection of these isochores with those for brine inclusions in the same samples may give the P and T of trapping of the fluids.     

The system pargasite-H2O-CO2: a model for melting of a hydrous mineral with a mixed-volatile fluid—I. Experimental results to 8 kbar

The stability of the amphibole pargasite [NaCa₂Mg₄Al(Al₂Si₆))O₂₂(OH)₂] in the melting range has been determined at total pressures (P) of 1.2 to 8 kbar. The activity of H₂O was controlled independently of P by using mixtures of H₂O + CO₂ in the fluid phase. The mole fraction of H₂O in the fluid ($\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}$) ranged from 1.0 to 0.2.At P < 4 kbar the stability temperature (T) of pargasite decreases with decreasing $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}$ at constant P. Above P ? 4 kbar stability T increases as $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}$ is decreased below one, passes through a T maximum and then decreases with a further decrease in $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}$. This behavior is due to a decrease in the H₂O content of the silicate liquid as $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}$ decreases. The magnitude of the T maximum increases from about 10°C (relative to the stability T for $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}\text{= 1}$) at P = 5 kbar to about 30°C at P = 8 kbar, and the position of the maximum shifts from $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}\text{? 0.6}$ at P = 5 kbar to $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}\text{? 0.4}$ at P = 8 kbar.The H₂O content of liquid coexisting with pargasite has been estimated as a function of at 5 and 8 kbar P, and can be used to estimate the H₂O content of magmas. Because pargasite is stable at low values of $\text{X}{}_{\mathrm{H}}{}_2\text{O}{}^{\mathrm{1fl}}$ at high P and T, hornblende can be an important phase in igneous processes even at relatively low H₂O fugacities.     

Pressure sensitive “silica geothermometer” determined from quartz solubility experiments at 250 °C

Experimentally reversed quartz solubilities at 250°C and at 250, 500 and 1000 bars yield values of the logarithm of the molality of aqueous silica of ?2.126, ?2.087 and ?2.038, respectively. Extrapolation of quartz solubility to the saturation pressure of water at 250°C results in a log molality of aqueous silica of-2.168. These solubility determinations and analyses of fluid pressures in geothermal systems indicate that pressure is significant when calculating quartz equilibrium temperatures from silica concentrations in waters of deep thermal reservoirs.The results of this investigation, combined with other reported quartz solubility measurements, yielded a pressure-sensitive “silica geothermometer” for fluids that have undergone adiabatic steam loss of where P is the fluid pressure in bars and m_Si(OH)₄ · 2H₂O represents the molality of aqueous silica measured in surface samples. The geothermometer is applicable to solutions in equilibrium with quartz from 180°C to 340°C and fluid pressures from H₂O saturation to 500 bars.     

Thermodynamic stability studies of the basic copper carbonate mineral,malachite

Natural malachite is a well defined solid demonstrating reproducible solubility behavior over a wide range of pH. The following equilibrium constants associated with the malachite dissolution equilibrium at 25°C, 1 atm were determined:

(infinite dilution)

$\text{K}{}^1{}_{\mathrm{sp}}\text{=}\text{[Cu}{}^{\mathrm{2+}}\text{]}{}^2\text{[CO}{}^{\mathrm{2?}}{}_3\text{]K}{}^2{}_{\mathrm{w}}\text{a}{}^2{}_{\mathrm{H+}}\text{= 10. ± 0.2 × 10}{}^{\mathrm{?32}}$

(0.72 ionic strength)

(36.9‰ salinity seawater). The temperature dependence of a “mixed” equilibrium constant, K_sp⁺, of the form:

has been measured at I = 0.72, yielding the relationship:

$\text{log K}{}^2{}_{\mathrm{sp}}\text{= (? 9.8 ± 0.03) × 10}{}^4\text{(}\text{1}\text{T°K}\text{) + (1.52 ± 0.09)}$

within a 5–25°C temperature range. The effect of pressure on the solubility of malachite in water and seawater was estimated from partial molar volume and compressibility data. For 25 °C at infinite dilution $\text{K'}{}_{\mathrm{sp}}\text{(1000 bar)}\text{K'}{}_{\mathrm{sp}}\text{(0)}\text{= 240}$ and in seawater $\text{K′}{}_{\mathrm{sp}}\text{(1000)}\text{K'}{}_{\mathrm{sp}}\text{(0)}\text{= 44}$.Comparison of stoichiometric and apparent malachite equilibrium constants has been used to estimate the extent of copper(II) ion interaction at the ionic strength of seawater. In dilute carbonate medium (total alkalinity, TA = 2.4 meq/kg H₂O, pH 8.3), 2.9% of total dissolved copper exists as the free copper(II) ion and in seawater (S = 36.9%., TA = 2.3 meq/kg H₂O, pH = 8.1), $\text{[Cu}{}^{\mathrm{2+}}\text{]}\text{T(Cu)}$ is 3.1%.Total dissolved copper levels of approximately 450–750 nMol/Kg are necessary to attain malachite saturation conditions in the open ocean. Observations of malachite particles suspended in seawater must be explained by precipitation or solid phase substitution reactions from localized environments rather than by direct precipitation from bulk seawater.     

An empirical solvus for CO2-H2O-2.6 wt% salt

The solvus in the system CO₂-H₂O-2.6 wt% NaCl-equivalent was determined by measuring temperature of homogenization in fluid inclusions which contained variable $\text{CO}{}_2\text{H}{}_2\text{O}$ but the same amount of salt dissolved in the aqueous phase at room temperature. The critical point of the solvus is at 340 ± 5°C, at pressures between 1 and 2 kbar; this is about 65°C higher than for the pure CO₂-H₂O system. The solvus is assymetrical, with a steeper H₂O-rich limb and with the critical point at mole fraction of water between 0.65 and 0.8.     

Fractionation of stable carbon isotopes by marine phytoplankton

Stable carbon isotope fractionation by seventeen species of marine phytoplankton, representing the classes of Bacillariophyceae, Chlorophyceae, Prasinophyceae, Chrysophyceae, Haptophyceae and Dinophyceae have been determined in laboratory culture experiments using bicarbonate enriched artificial sea water. The values (Δ_HCO3^? = δ¹³C of algae vs HCO₃^?) range from ?22.1 to ?35.5%. Nitzschia closterium shows the smallest fractionation of ? 22.1% and Isochrysis galbana, the greatest of ?35.5%,. Since these algae were cultured under identical laboratory conditions, the wide range of Δ_HCO3^? values is seemingly due to the presence of different metabolic pathways within these organisms.A temperature dependent fractionation of 0.36% per °C with decreasing temperatures was measured for Skeletonema costatum whereas, smaller temperature dependencies of ?0.13, +0.15 and ?0.07%. per °C were observed for Dunaliella sp., Monochrysis lutheri and Glenodinium foliaceum, respectively.The consistency of values of Skeletonema costatum, Dunaliella sp. and Monochrysis lutheri grown at salinities of 22, 26, 32 and 36% indicates that natural salinity variations have negligible effects on the isotopic composition of marine phytoplankton.     

Determination of the 87Rb decay constant

The decay constant ⁸⁷Rb has been redetermined by measuring the amount of radiogenic ⁸⁷Sr produced over a period of 19 years, in 20 g samples of purified RbClO₄, using isotope dilution techniques. The rubidium sample was spiked with ⁸⁴Sr and the nanogram quantities of strontium separated by coprecipitation with Ba(NO₃)₂. Analyses were carried out on a 25cm, 90° sector mass spectrometer equipped with a Spiraltron electron multiplier. Measurement of three independent ratios permitted continuous monitoring of the ion beam fractionation. The average of nine determinations gives a value for the decay constant of $\text{1.419(±0.012) × 10}{}^{\mathrm{?11}}\text{yr}{}^{\mathrm{?1}}\text{(2σ)}$. [$\text{τ}{}_{\mathrm{<mtext>1</mtext><mtext>2</mtext>}}\text{= 4.89(±0.04) × 10}{}^{10}\text{yr}$.]     

Transfer and exchange equilibria in a portion of the pyroxene quadrilateral as deduced from natural and experimental data

Differences in the chemical composition of metamorphic and igneous pyroxene minerals may be attributed to a transfer reaction, which determines the Ca content of the minerals, and an exchange reaction, which determines the relative Mg:Fe²⁺ ratios. Natural data for associated Ca pyroxene (Cpx) and orthopyroxene (Opx) or pigeonite are combined with experimental data for Fe-free pyroxenes, to produce the following equations for the Cpx slope of the solvus surface: $\text{> 1080°C: T =}\text{1000}\text{(0.468 + 0.246X}{}^{\mathrm{Cpx}}\text{? 0.123 ln (1–2 [Ca]))}$$\text{< 1080°C: T =}\text{1000}\text{(0.054 + 0.608X}{}^{\mathrm{Cpx}}\text{? 0.304 ln (1–2 [Ca]))}$, and the following equation for the temperature-dependence of the Mg-Fe distribution coefficient: $\text{T =}\text{1130}\text{(}\text{ln}\text{K}{}_{\mathrm{p}}\text{+ 0.505)}$, where T is absolute temperature, X is $\text{Fe}{}^{\mathrm{2+}}\text{(Mg + Fe}{}^{\mathrm{2+}}\text{))}$, [Ca] is $\text{Ca}\text{(Ca + Mg + Fe}{}^{\mathrm{2+}}\text{)}$ in Cpx, and K_D is the distribution coefficient, defined as $\text{X}{}^{\mathrm{Opx}}\text{/(1 ? X}{}^{\mathrm{Opx}}\text{) ÷ X}{}_{\mathrm{Cpx}}\text{/(1 ?}{}^{\mathrm{Cpx}}\text{)}$.The transfer and exchange equations form useful temperature indicators, and when applied to 9 sets of well-studied rocks, yield pairs of temperatures that are in good agreement. For example, temperatures obtained for the Bushveld Complex are 1020°C (solvus equation) and 980°C (exchange equation), based on 7 specimens. The uncertainty in these numbers, due to precision and accuracy errors, is estimated to be ±60°.