ODM: an analytical solution-based tool for reacting oxygen diffusion modelling in mine spoils

A simple one-dimensional analytical solution is presented to model oxygen diffusion through the pore space of mine spoils containing pyrite. The model incorporates volumetric oxygen consumption terms due to pyrite oxidation, oxidation of Fe ²⁺ to Fe ³⁺ and bacterial activity. Based on this analytical solution, a graphical user interface (GUI) tool is programmed and designed in MATLAB software. This tool can be used to model transport of oxygen through the mine spoils either with or without a cap. Results of several simulation scenarios of sensitivity analysis showed a significant change in oxygen concentration with varying effective diffusion coefficient of oxygen transport model and simulation time. Efficiency and flexibility of the tool developed here is verified by modelling oxygen transport through the pore space of a coal waste pile (case A) and a copper mine tailings (case B). Maximum depth of oxygen diffusion is obtained approximately equal to 2 and 1.5 m through the cases A and B, respectively.     

Phenylalanine as a hydroxyl radical-specific probe in pyrite slurries

The abundant iron sulfide mineral pyrite has been shown to catalytically produce hydrogen peroxide (H₂O₂) and hydroxyl radical ( ^. OH) in slurries of oxygenated water. Understanding the formation and fate of these reactive oxygen species is important to biological and ecological systems as exposure can lead to deleterious health effects, but also environmental engineering during the optimization of remediation approaches for possible treatment of contaminated waste streams. This study presents the use of the amino acid phenylalanine (Phe) to monitor the kinetics of pyrite-induced ^. OH formation through rates of hydroxylation forming three isomers of tyrosine (Tyr) - ortho-, meta-, and para-Tyr. Results indicate that about 50% of the Phe loss results in Tyr formation, and that these products further react with ^. OH at rates comparable to Phe. The overall loss of Phe appeared to be pseudo first-order in [Phe] as a function of time, but for the first time it is shown that initial rates were much less than first-order as a function of initial substrate concentration, [Phe]_o. These results can be rationalized by considering that the effective concentration of ^. OH in solution is lower at a higher level of reactant and that an increasing fraction of ^. OH is consumed by Phe-degradation products as a function of time. A simplified first-order model was created to describe Phe loss in pyrite slurries which incorporates the [Phe]_o, a first-order dependence on pyrite surface area, the assumption that all Phe degradation products compete equally for the limited supply of highly reactive ^. OH, and a flux that is related to the release of H₂O₂ from the pyrite surface (a result of the incomplete reduction of oxygen at the pyrite surface). An empirically derived rate constant, K _pyr , was introduced to describe a variable ^. OH-reactivity for different batches of pyrite. Both the simplified first-order kinetic model, and a more detailed numerical simulation, yielded results that compare well to the observed kinetic data describing the effects of variations in concentrations of both initial Phe and pyrite. This work supports the use of Phe as a useful probe to assess the formation of ^. OH in the presence of pyrite, and its possible utility for similar applications with other minerals.     

Galena Non-oxidative Dissolution Kinetics in Seawater

The rate of non-oxidative galena dissolution in seawater compositions over the pH range of 2–4.5 was determined from batch reactor experiments. The derivative at zero time of a polynomial fit of the Pb concentration versus time data for the first 30 min was used to determine the rate. A plot of R_Gn (rate of galena dissolution) versus pH for data from six experiments is linear (R²?=?0.96), with a slope of 0.5. The rate equation describing the rate of galena dissolution as a function of hydrogen ion activity is

$$R_{\text{Gn}} = - \,10^{ - 10.72} \left( {a_{{{\text{H}}^{ + } }} } \right)^{0.50}$$

Varying the concentration of dissolved oxygen produced no significant effect on the measured rates. The activation energy, based on four experiments carried out over the temperature range of 7–30 °C, is 61.1 kJ/mol.

     

Solubility of gold in arsenian pyrite

Although Au and As can be enriched up to the weight percent level in arsenian pyrite, there is little knowledge of their limiting concentrations and nature of incorporation. This study reports SIMS and EMPA analyses showing that As and Au contents of arsenian pyrites plot in a wedge-shaped zone with an upper compositional limit defined by the line

CAu=0.02⋅CAs+4×10⁻⁵     

Retention of biosignatures and mass-independent fractionations in pyrite: Self-diffusion of sulfur

Self-diffusion of sulfur in pyrite (FeS₂) was characterized over the temperature range ∼500-725 °C (∼1 bar pressure) by immersing natural specimens in a bath of molten elemental ³⁴S and characterizing the resulting diffusive-exchange profiles by Rutherford backscattering spectroscopy (RBS). The temperature dependence of the sulfur diffusivity (D_S) conforms to D_S = D_o exp(−E_a/RT), where the pre-exponential constant (D_o) and the activation energy (E_a) are constrained as follows:

     

A mechanism for the production of hydroxyl radical at surface defect sites on pyrite

A previous contribution from our laboratory reported the formation of hydrogen peroxide (H₂O₂) upon addition of pyrite (FeS₂) to O₂-free water. It was hypothesized that a reaction between adsorbed H₂O and Fe(III), at a sulfur-deficient defect site, on the pyrite surface generates an adsorbed hydroxyl radical (OH^•).

     

Viscosity of andesitic melts—new experimental data and a revised calculation model

The viscosity of a synthetic andesite-like melt was measured in the low viscosity range (10¹-10⁶ Pa s) using the falling sphere(s) method and in the high viscosity range (10⁸-10¹³ Pa s) using parallel-plate viscometry. Falling sphere experiments with melts containing 2.3 and 5.6 wt.% H₂O were carried out in an internally heated gas pressure vessel (IHPV) at 500 MPa confining pressure. The sinking velocity of Pt and Pd spheres and in one case of a corundum sphere was used to measure the melt viscosity. In addition, a creep experiment was performed at ambient pressure using a glass containing 2.73 wt.% H₂O . A more water-rich glass (5.6 wt.% H₂O ) was investigated with a high pressure parallel-plate viscometer at 400 MPa confining pressure in an IPHV. By combining our new data with previous results for a similar melt composition we derived the following expression to describe the viscosity η (in Pa s) as a function of temperature T (in K) and water content w (in wt.%)

     

Water diffusion in Mount Changbai peralkaline rhyolitic melt

Diffusion couple experiments with wet half (up to 4.6 wt%) and dry half were carried out at 789–1,516 K and 0.47–1.42 GPa to investigate water diffusion in a peralkaline rhyolitic melt with major oxide concentrations matching Mount Changbai rhyolite. Combining data from this work and a related study, total water diffusivity in peralkaline rhyolitic melt can be expressed as:

$ D_{{{\text{H}}_{ 2} {\text{O}}_{\text{t}} }} = D_{{{\text{H}}_{ 2} {\text{O}}_{\text{m}} }} \left( {1 - \frac{0.5 - X}{{\sqrt {[4\exp (3110/T - 1.876) - 1](X - X^{2} ) + 0.25} }}} \right), $

$ {\text{with}}\;D_{{{\text{H}}_{ 2} {\text{O}}_{\text{m}} }} = \exp \left[ { - 1 2. 7 8 9- \frac{13939}{T} - 1229.6\frac{P}{T} + ( - 27.867 + \frac{60559}{T})X} \right], $

where D is in m² s^?1, T is the temperature in K, P is the pressure in GPa, and X is the mole fraction of water and calculated as X = (C/18.015)/(C/18.015 + (100 ? C)/33.14), where C is water content in wt%. We recommend this equation in modeling bubble growth and volcanic eruption dynamics in peralkaline rhyolitic eruptions, such as the ~1,000-ad eruption of Mount Changbai in North East China. Water diffusivities in peralkaline and metaluminous rhyolitic melts are comparable within a factor of 2, in contrast with the 1.0–2.6 orders of magnitude difference in viscosities. The decoupling of diffusivity of neutral molecular species from melt viscosity, i.e., the deviation from the inversely proportional relationship predicted by the Stokes–Einstein equation, might be attributed to the small size of H₂O molecules. With distinct viscosities but similar diffusivity, bubble growth controlled by diffusion in peralkaline and metaluminous rhyolitic melts follows similar parabolic curves. However, at low confining pressure or low water content, viscosity plays a larger role and bubble growth rate in peralkaline rhyolitic melt is much faster than that in metaluminous rhyolite.

     

Thermal equation of state of CaIrO<Subscript>3</Subscript> post-perovskite

The pressure–volume–temperature (P–V–T) relation of CaIrO₃ post-perovskite (ppv) was measured at pressures and temperatures up to 8.6 GPa and 1,273 K, respectively, with energy-dispersive synchrotron X-ray diffraction using a DIA-type, cubic-anvil apparatus (SAM85). Unit-cell dimensions were derived from the Le Bail full profile refinement technique, and the results were fitted using the third-order Birth-Murnaghan equation of state. The derived bulk modulus \( K_{T0} \) at ambient pressure and temperature is 168.3 ± 7.1 GPa with a pressure derivative \( K_{T0}^{\prime } \) = 5.4 ± 0.7. All of the high temperature data, combined with previous experimental data, are fitted using the high-temperature Birch-Murnaghan equation of state, the thermal pressure approach, and the Mie-Grüneisen-Debye formalism. The refined thermoelastic parameters for CaIrO₃ ppv are: temperature derivative of bulk modulus \( (\partial K_{T} /\partial T)_{P} \) = ?0.038 ± 0.011 GPa K^?1, \( \alpha K_{T} \) = 0.0039 ± 0.0001 GPa K^?1, \( \left( {\partial K_{T} /\partial T} \right)_{V} \) = ?0.012 ± 0.002 GPa K^?1, and \( \left( {\partial^{2} P/\partial T^{2} } \right)_{V} \) = 1.9 ± 0.3 × 10^?6 GPa² K^?2. Using the Mie-Grüneisen-Debye formalism, we obtain Grüneisen parameter \( \gamma_{0} \) = 0.92 ± 0.01 and its volume dependence q = 3.4 ± 0.6. The systematic variation of bulk moduli for several oxide post-perovskites can be described approximately by the relationship K _T0  = 5406.0/V(molar) + 5.9 GPa.     

Oxygen isotopic fractionation during inorganic calcite precipitation ― Effects of temperature, precipitation rate and pH

Stable oxygen isotopic fractionation during inorganic calcite precipitation was experimentally studied by spontaneous precipitation at various pH (8.3 < pH < 10.5), precipitation rates (1.8 < log R < 4.4 μmol m^− 2 h^− 1) and temperatures (5, 25, and 40 °C) using the CO₂ diffusion technique.The results show that the apparent stable oxygen isotopic fractionation factor between calcite and water (α_{calcite–water}) is affected by temperature, the pH of the solution, and the precipitation rate of calcite. Isotopic equilibrium is not maintained during spontaneous precipitation from the solution. Under isotopic non-equilibrium conditions, at a constant temperature and precipitation rate, apparent 1000lnα_{calcite–water} decreases with increasing pH of the solution. If the temperature and pH are held constant, apparent 1000lnα_{calcite–water} values decrease with elevated precipitation rates of calcite. At pH = 8.3, oxygen isotopic fractionation between inorganically precipitated calcite and water as a function of the precipitation rate (R) can be described by the expressions

at 5, 25, and 40 °C, respectively.The impact of precipitation rate on 1000lnα_{calcite–water} value in our experiments clearly indicates a kinetic effect on oxygen isotopic fractionation during calcite precipitation from aqueous solution, even if calcite precipitated slowly from aqueous solution at the given temperature range. Our results support Coplen's work [Coplen T. B. (2007) Calibration of the calcite–water oxygen isotope geothermometer at Devils Hole, Nevada, a natural laboratory. Geochim. Cosmochim. Acta 71, 3948–3957], which indicates that the equilibrium oxygen isotopic fractionation factor might be greater than the commonly accepted value.     

The P–V–T equation of state of CaPtO3 post-perovskite

Orthorhombic post-perovskite CaPtO₃ is isostructural with post-perovskite MgSiO₃, a deep-Earth phase stable only above 100 GPa. Energy-dispersive X-ray diffraction data (to 9.4 GPa and 1,024 K) for CaPtO₃ have been combined with published isothermal and isobaric measurements to determine its P–V–T equation of state (EoS). A third-order Birch–Murnaghan EoS was used, with the volumetric thermal expansion coefficient (at atmospheric pressure) represented by α(T) = α₀ + α₁(T). The fitted parameters had values: isothermal incompressibility, $ K_{{T_{0} }} $  = 168.4(3) GPa; $ K_{{T_{0} }}^{\prime } $  = 4.48(3) (both at 298 K); $ \partial K_{{T_{0} }} /\partial T $  = ?0.032(3) GPa K^?1; α₀ = 2.32(2) × 10^?5 K^?1; α₁ = 5.7(4) × 10^?9 K^?2. The volumetric isothermal Anderson–Grüneisen parameter, δ _T , is 7.6(7) at 298 K. $ \partial K_{{T_{0} }} /\partial T $ for CaPtO₃ is similar to that recently reported for CaIrO₃, differing significantly from values found at high pressure for MgSiO₃ post-perovskite (?0.0085(11) to ?0.024 GPa K^?1). We also report axial P–V–T EoS of similar form, the first for any post-perovskite. Fitted to the cubes of the axes, these gave $ \partial K_{{aT_{0} }} /\partial T $  = ?0.038(4) GPa K^?1; $ \partial K_{{bT_{0} }} /\partial T $  = ?0.021(2) GPa K^?1; $ \partial K_{{cT_{0} }} /\partial T $  = ?0.026(5) GPa K^?1, with δ _T  = 8.9(9), 7.4(7) and 4.6(9) for a, b and c, respectively. Although $ K_{{T_{0} }} $ is lowest for the b-axis, its incompressibility is the least temperature dependent.     

Nitrogen solubility in basaltic melt. Part I. Effect of oxygen fugacity

The role of the oxygen fugacity on the incorporation of nitrogen in basaltic magmas has been investigated using one atmosphere high temperature equilibration of tholeiitic-like compositions under controlled nitrogen and oxygen partial pressures in the [C-N-O] system. Nitrogen was extracted with a CO₂ laser under high vacuum and analyzed by static mass spectrometry. Over a redox range of 18 oxygen fugacity log units, this study shows that the incorporation of nitrogen in silicate melts follows two different behaviors. For log fO₂ values between −0.7 and −10.7 (the latter corresponding to IW − 1.3), nitrogen dissolves as a N₂ molecule into cavities of the silicate network (physical solubility). Nitrogen presents a constant solubility (Henry’s) coefficient of 2.21 ± 0.53 × 10⁻⁹ mol g⁻¹ atm⁻¹ at 1425°C, identical within uncertainties to the solubility of argon. Further decrease in the oxygen fugacity (log fO₂ between −10.7 and −18 corresponding to the range from IW − 1.3 to IW − 8.3) results in a drastic increase of the solubility of nitrogen by up to 5 orders of magnitude as nitrogen becomes chemically bounded with atoms of the silicate melt network (chemical solubility). The present results strongly suggest that under reducing conditions nitrogen dissolves in silicate melts as N³⁻ species rather than as CN⁻ cyanide radicals. The nitrogen content of a tholeiitic magma equilibrated with N₂ is computed from thermochemical processing of our data set as

     

CO<Subscript>2</Subscript> content of andesitic melts at graphite-saturated upper mantle conditions with implications for redox state of oceanic basalt source regions and remobilization of reduced carbon from subducted eclogite

We have performed experiments to determine the effects of pressure, temperature and oxygen fugacity on the CO₂ contents in nominally anhydrous andesitic melts at graphite saturation. The andesite composition was specifically chosen to match a low-degree partial melt composition that is generated from MORB-like eclogite in the convective, oceanic upper mantle. Experiments were performed at 1–3 GPa, 1375–1550?°C, and fO₂ of FMQ ?3.2 to FMQ ?2.3 and the resulting experimental glasses were analyzed for CO₂ and H₂O contents using FTIR and SIMS. Experimental results were used to develop a thermodynamic model to predict CO₂ content of nominally anhydrous andesitic melts at graphite saturation. Fitting of experimental data returned thermodynamic parameters for dissolution of CO₂ as molecular CO₂: ln(K ⁰) = ?21.79?±?0.04, ΔV ⁰?=?32.91?±?0.65 cm³mol^?1, ΔH ⁰?=?107?±?21 kJ mol^?1, and dissolution of CO₂ as CO₃ ^2?: ln(K ⁰ ) = ?21.38?±?0.08, ΔV ⁰?=?30.66?±?1.33 cm³ mol^?1, ΔH ⁰?=?42?±?37 kJ mol^?1, where K ⁰ is the equilibrium constant at some reference pressure and temperature, ΔV ⁰ is the volume change of reaction, and ΔH ⁰ is the enthalpy change of reaction. The thermodynamic model was used along with trace element partition coefficients to calculate the CO₂ contents and CO₂/Nb ratios resulting from the mixing of a depleted MORB and the partial melt of a graphite-saturated eclogite. Comparison with natural MORB and OIB data suggests that the CO₂ contents and CO₂/Nb ratios of CO₂-enriched oceanic basalts cannot be produced by mixing with partial melts of graphite-saturated eclogite. Instead, they must be produced by melting of a source containing carbonate. This result places a lower bound on the oxygen fugacity for the source region of these CO₂-enriched basalts, and suggests that fO₂ measurements made on cratonic xenoliths may not be applicable to the convecting upper mantle. CO₂-depleted basalts, on the other hand, are consistent with mixing between depleted MORB and partial melts of a graphite-saturated eclogite. Furthermore, calculations suggest that eclogite can remain saturated in graphite in the convecting upper mantle, acting as a reservoir for C.     

Experimental studies of oxygen isotope fractionation between rhodochrosite (MnCO3) and water at low temperatures

Rhodochrosite crystals were precipitated from Na-Mn-Cl-HCO₃ parent solutions following passive, forced and combined passive-to-forced CO₂ degassing methods. Forced and combined passive-to-forced CO₂ degassing produced rhodochrosite crystals with a small non-equilibrium oxygen isotope effect whereas passive CO₂ degassing protocols yielded rhodochrosite in apparent isotopic equilibrium with water. On the basis of the apparent equilibrium isotopic data, a new temperature-dependent relation is proposed for the oxygen isotope fractionation between rhodochrosite and water between 10 and 40 °C:

1000lnα_{rhodochrosite-water}=17.84±0.18(10³/T)-30.24±0.62     

Reconstruction of seasonal temperature variability in the tropical Pacific Ocean from the shell of the scallop, Comptopallium radula

We investigated the oxygen isotope composition (δ¹⁸O) of shell striae from juvenile Comptopallium radula (Mollusca; Pectinidae) specimens collected live in New Caledonia. Bottom-water temperature and salinity were monitored in-situ throughout the study period. External shell striae form with a 2-day periodicity in this scallop, making it possible to estimate the date of precipitation for each calcite sample collected along a growth transect. The oxygen isotope composition of shell calcite (δ¹⁸O_{shell calcite}) measured at almost weekly resolution on calcite accreted between August 2002 and July 2003 accurately tracks bottom-water temperatures. A new empirical paleotemperature equation for this scallop species relates temperature and δ¹⁸O_{shell calcite}:

t(°C)=20.00(±0.61)-3.66(±0.39)×(δ¹⁸O_{shell calcite VPDB}-δ¹⁸O_{water VSMOW})     

Influence of oxygen fugacity and temperature on the redox state of iron in natural silicic aluminosilicate melts

The influence of oxygen fugacity (fO₂) and temperature on the valence and structural state of iron was experimentally studied in glasses quenched from natural aluminosilicate melts of granite and pantellerite compositions exposed to various T-fO₂ conditions (1100–1420°C and 10^?12–10^?0.68 bar) at a total pressure of 1 atm. The quenched glasses were investigated by Mössbauer spectroscopy. It was shown that the effect of oxygen fugacity on the redox state of iron at 1320–1420°C can be described by the equation log(Fe³⁺/Fe²⁺) = k log(fO₂) + q, where k and q are constants depending on melt composition and temperature. The Fe³⁺/Fe²⁺ ratio decreases with decreasing fO₂ (T = const) and increasing temperature (fO₂ = const). The structural state of Fe³⁺ depends on the degree of iron oxidation. With increasing Fe³⁺/Fe²⁺ ≥ 1, the dominant coordination of Fe³⁺ changes from octahedral to tetrahedral. Ferrous iron ions occur in octahedral (and/or five-coordinated) sites independent of Fe³⁺/Fe²⁺.     

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)     

Separation of chemical elements in the atmospheres of CP stars under the action of light induced drift

A mechanism for the separation of chemical elements and isotopes in the atmospheres of chemically peculiar (CP) stars due to light-induced drift (LID) of ions is discussed. The efficiency of separation due to LID is proportional to the relative difference of the transport frequencies for collisions of ions of heavy elements located in the excited state (collision frequency ν _e ) and ground state (collision frequency ν _g ) with neutral buffer particles (hydrogen and helium), (ν _e ? ν _g )/ν _g . The known interaction potentials are used to numerically compute the relative difference (ν _e ^H ? ν _g ^H )/νg H for collisions between the ions Be⁺, Mg⁺, Ca⁺, Sr⁺, Cd⁺, Ba⁺, Al⁺, and C⁺ and hydrogen atoms. These computations show that, at the temperatures characteristic of the atmospheres of CP stars, T = 7000?20 000 K, values of |ν _e ^H ?ν _g ^H |/ν _g ^H ≈ 0.1?0.4 are obtained. With such relative differences in the transport collision frequencies, the LID rate of ions in the atmospheres of coolCP stars (T < 10000 K) can reach ~0.1 cm/s,which exceeds the drift rate due to light pressure by an order of magnitude. This means that, under these conditions, the separation of chemical elements under the action of LID of ions could be an order of magnitude more efficient than separation due to light pressure. Roughly the same manifestations of LID and light pressure are also expected in the atmospheres of hotter stars (20 000 > T > 10 000 K). LID of heavy ions is manifest only weakly in very hot stars (T > 20 000 K).     

Equilibrium mass-dependent fractionation relationships for triple oxygen isotopes

With a growing interest in small ¹⁷O-anomaly, there is a pressing need for the precise ratio, ln¹⁷α/ln¹⁸α, for a particular mass-dependent fractionation process (MDFP) (e.g., for an equilibrium isotope exchange reaction). This ratio (also denoted as “θ”) can be determined experimentally, however, such efforts suffer from the demand of well-defined process or a set of processes in addition to high precision analytical capabilities. Here, we present a theoretical approach from which high-precision ratios for MDFPs can be obtained. This approach will complement and serve as a benchmark for experimental studies. We use oxygen isotope exchanges in equilibrium processes as an example.We propose that the ratio at equilibrium, θ^E ≡ ln¹⁷α/ln¹⁸α, can be calculated through the equation below: