Dissolution rates of metals in Fe oxides: implications for sampling ferruginous materials with significant relict Fe oxides

A study of the pattern of dissolution of synthetic and natural Fe oxides in 6 M HCl indicates that the rate of element release from synthetic Fe oxides is strongly influenced by mineralogy and the level of element incorporation. Synthetic maghemite (γ-Fe₂O₃) samples are subject to much more rapid dissolution than goethite (FeOOH) and hematite (α-Fe₂O₃). In samples dominated by hematite and maghemite, Cu, Zn and particularly Pb, in comparison to Fe, are preferentially released during the early stages of dissolution. Similar patterns are apparent from the dissolution of hematite- and maghemite-dominated samples derived from natural gossan. Comparison of XRD scans with data from the dissolution of natural gossan samples transformed by incremental heating to hematite- and maghemite-dominated assemblages suggests that the degree of crystallinity may also be a significant factor in the release of elements incorporated in the Fe oxides. Ferruginous materials made up of varying proportions of goethite, hematite, maghemite, kaolinite and quartz are important sampling materials in a range of regolith environments. These are products of complex chemical and mechanical mobilization over long periods of geological time. If the patterns of Fe oxide dissolution in 6 M HCl and the release of incorporated metals reflect stability in such weathering regimes, knowledge of the retention characteristics of incorporated metals in different Fe oxide phases, as indicated by this study, will be useful in the planning and interpretation of geochemical surveys in such regions.     

Inner-Helmholtz potential development at the hematite (α-Fe2O3) (0 0 1) surface

Electric potentials of the (0 0 1) surface of hematite were measured as a function of pH and ionic strength in solutions of sodium nitrate and oxalic acid using the single-crystal electrode approach. The surface is predominantly charge-neutral in the pH 4-14 range, and develops a positive surface potential below pH 4 due to protonation of μ-OH⁰ sites (pK_1,1,0,int = −1.32). This site is resilient to deprotonation up to at least pH 14 (−pK_−1,1,0,int ? 19). The associated Stern layer capacitance of 0.31-0.73 F/m² is smaller than typical values of powders, and possibly arises from a lower degree of surface solvation. Acid-promoted dissolution under elevated concentrations of HNO₃ etches the (0 0 1) surface, yielding a convoluted surface populated by sites. The resulting surface potential was therefore larger under these conditions than in the absence of dissolution. Oxalate ions also promoted (0 0 1) dissolution. Associated electric potentials were strongly negative, with values as large as −0.5 V, possibly from metal-bonded interactions with oxalate. The hematite surface can also acquire negative potentials in the pH 7-11 range due to surface complexation and/or precipitation of iron species (0.0038 Fe/nm²) produced from acidic conditions. Oxalate-bearing systems also result in negative potentials in the same pH range, and may include ferric-oxalate surface complexes and/or surface precipitates. All measurements can be modeled by a thermodynamic model that can be used to predict inner-Helmholtz potentials of hematite surfaces.     

First experimental determination of iron isotope fractionation between hematite and aqueous solution at hydrothermal conditions

Although iron isotopes provide a new powerful tool for tracing a variety of geochemical processes, the unambiguous interpretation of iron isotope ratios in natural systems and the development of predictive theoretical models require accurate data on equilibrium isotope fractionation between fluids and minerals. We investigated Fe isotope fractionation between hematite (Fe₂O₃) and aqueous acidic NaCl fluids via hematite dissolution and precipitation experiments at temperatures from 200 to 450 °C and pressures from saturated vapor pressure (P_sat) to 600 bar. Precipitation experiments at 200 °C and P_sat from aqueous solution, in which Fe aqueous speciation is dominated by ferric iron (Fe^III) chloride complexes, show no detectable Fe isotope fractionation between hematite and fluid, Δ⁵⁷Fe_{fluid-hematite} = δ⁵⁷Fe_fluid − δ⁵⁷Fe_hematite = 0.01 ± 0.08‰ (2 × standard error, 2SE). In contrast, experiments at 300 °C and P_sat, where ferrous iron chloride species (FeCl₂ and FeCl⁺) dominate in the fluid, yield significant fluid enrichment in the light isotope, with identical values of Δ⁵⁷Fe_{fluid-hematite} = −0.54 ± 0.15‰ (2SE) both for dissolution and precipitation runs. Hematite dissolution experiments at 450 °C and 600 bar, in which Fe speciation is also dominated by ferrous chloride species, yield Δ⁵⁷Fe_{fluid-hematite} values close to zero within errors, 0.15 ± 0.17‰ (2SE). In most experiments, chemical, redox, and isotopic equilibrium was attained, as shown by constancy over time of total dissolved Fe concentrations, aqueous Fe^II and Fe^III fractions, and Fe isotope ratios in solution, and identical Δ⁵⁷Fe values from dissolution and precipitation runs. Our measured equilibrium Δ⁵⁷Fe_{fluid-hematite} values at different temperatures, fluid compositions and iron redox state are within the range of fractionations in the system fluid-hematite estimated using reported theoretical β-factors for hematite and aqueous Fe species and the distribution of Fe aqueous complexes in solution. These theoretical predictions are however affected by large discrepancies among different studies, typically ±1‰ for the Δ⁵⁷Fe _{Fe(aq)-hematite} value at 200 °C. Our data may thus help to refine theoretical models for β-factors of aqueous iron species. This study provides the first experimental calibration of Fe isotope fractionation in the system hematite-saline aqueous fluid at elevated temperatures; it demonstrates the importance of redox control on Fe isotope fractionation at hydrothermal conditions.     

Kinetics of biotite dissolution and Fe behavior under low O2 conditions and their implications for Precambrian weathering

Biotite dissolution experiments were carried out to better understand the dissolution kinetics and Fe behavior under low O₂ conditions, and to give an insight into the Precambrian weathering. Mineral dissolution with a continuous flow-through reactor was employed at 25 °C for up to 65 days varying partial pressure of atmospheric oxygen (PO₂), pH (6.86 and 3.01) and Fe content in mineral (1.06 and 0.11 mol of Fe per O₁₀(OH,F)₂ for biotite and phlogopite, respectively) independently for the examination of their effects on biotite dissolution. Low PO₂ conditions were achieved in a newly developed glove box (PO₂ ? 6 × 10⁻⁴ atm; referred to as anoxic conditions), which was compared to the present, ambient air conditions (0.2 atm of PO₂; oxic conditions). The biotite dissolution rate was slightly faster under anoxic conditions at pH 6.86 while it was not affected by PO₂ at pH 3.01. There was no direct effect of Fe content on dissolution rate at pH 6.86 while there was a small difference in dissolution rate between biotite and phlogopite at pH 3.01. The 1.5 order-of-magnitude faster release rate of Fe under anoxic conditions for biotite dissolution at pH 6.86 resulted from the difference in ratio of Fe³⁺ precipitates remaining in the reactor to Fe dissolved (about 60% and 100% under anoxic and oxic conditions, respectively), which is caused mainly by the difference in PO₂. The results infer that the Fe²⁺ and Fe³⁺ contents in the Paleoproterozoic paleosols, fossil weathering profiles, are reflected by atmospheric oxygen levels at the time of weathering.     

Comparative dissolution kinetics of biogenic and chemogenic uraninite under oxidizing conditions in the presence of carbonate

The long-term stability of biogenic uraninite with respect to oxidative dissolution is pivotal to the success of in situ bioreduction strategies for the subsurface remediation of uranium legacies. Batch and flow-through dissolution experiments were conducted along with spectroscopic analyses to compare biogenic uraninite nanoparticles obtained from Shewanella oneidensis MR-1 and chemogenic UO_2.00 with respect to their equilibrium solubility, dissolution mechanisms, and dissolution kinetics in water of varied oxygen and carbonate concentrations. Both materials exhibited a similar intrinsic solubility of ∼10⁻⁸ M under reducing conditions. The two materials had comparable dissolution rates under anoxic as well as oxidizing conditions, consistent with structural bulk homology of biogenic and stoichiometric uraninite. Carbonate reversibly promoted uraninite dissolution under both moderately oxidizing and reducing conditions, and the biogenic material yielded higher surface area-normalized dissolution rates than the chemogenic. This difference is in accordance with the higher proportion of U(V) detected on the biogenic uraninite surface by means of X-ray photoelectron spectroscopy. Reasonable sources of a stable U(V)-bearing intermediate phase are discussed. The observed increase of the dissolution rates can be explained by carbonate complexation of U(V) facilitating the detachment of U(V) from the uraninite surface. The fraction of surface-associated U(VI) increased with dissolved oxygen concentration. Simultaneously, X-ray absorption spectra showed conversion of the bulk from UO_2.0 to UO_2+x. In equilibrium with air, combined spectroscopic results support the formation of a near-surface layer of approximate composition UO_2.25 (U₄O₉) coated by an outer layer of U(VI). This result is in accordance with flow-through dissolution experiments that indicate control of the dissolution rate of surface-oxidized uraninite by the solubility of metaschoepite under the tested conditions. Although U(V) has been observed in electrochemical studies on the dissolution of spent nuclear fuel, this is the first investigation that demonstrates the formation of a stable U(V) intermediate phase on the surface of submicron-sized uraninite particles suspended in aqueous solutions.     

Reaction-based modeling of quinone-mediated bacterial iron(III) reduction

This paper presents and validates a new paradigm for modeling complex biogeochemical systems using a diagonalized reaction-based approach. The bioreduction kinetics of hematite (α-Fe₂O₃) by the dissimilatory metal-reducing bacterium (DMRB) Shewanella putrefaciens strain CN32 in the presence of the soluble electron shuttling compound anthraquinone-2,6-disulfonate (AQDS) is used for presentation/validation purposes. Experiments were conducted under nongrowth conditions with H₂ as the electron donor. In the presence of AQDS, both direct biological reduction and indirect chemical reduction of hematite by bioreduced anthrahydroquinone-2,6-disulfonate (AH₂DS) can produce Fe(II). Separate experiments were performed to describe the bioreduction of hematite, bioreduction of AQDS, chemical reduction of hematite by AH₂DS, Fe(II) sorption to hematite, and Fe(II) biosorption to DMRB. The independently determined rate parameters and equilibrium constants were then used to simulate the parallel kinetic reactions of Fe(II) production in the hematite-with-AQDS experiments. Previously determined rate formulations/parameters for the bioreduction of hematite and Fe(II) sorption to hematite were systematically tested by conducting experiments with different initial conditions. As a result, the rate formulation/parameter for hematite bioreduction was not modified, but the rate parameters for Fe(II) sorption to hematite were modified slightly. The hematite bioreduction rate formulation was first-order with respect to hematite ”free“ surface sites and zero-order with respect to DMRB based on experiments conducted with variable concentrations of hematite and DMRB. The AQDS bioreduction rate formulation was first-order with respect to AQDS and first-order with respect to DMRB based on experiments conducted with variable concentrations of AQDS and DMRB. The chemical reduction of hematite by AH₂DS was fast and considered to be an equilibrium reaction. The simulations of hematite-with-AQDS experiments were very sensitive to the equilibrium constant for the hematite-AH₂DS reaction. The model simulated the hematite-with-AQDS experiments well if it was assumed that the ferric oxide “surface” phase was more disordered than pure hematite. This is the first reported study where a diagonalized reaction-based model was used to simulate parallel kinetic reactions based on rate formulations/parameters independently obtained from segregated experiments.     

Mechanisms of iron oxide transformations in hydrothermal systems

Coexistence of magnetite and hematite in hydrothermal systems has often been used to constrain the redox potential of fluids, assuming that the redox equilibrium is attained among all minerals and aqueous species. However, as temperature decreases, disequilibrium mineral assemblages may occur due to the slow kinetics of reaction involving the minerals and fluids. In this study, we conducted a series of experiments in which hematite or magnetite was reacted with an acidic solution under H₂-rich hydrothermal conditions (T = 100-250 °C, ) to investigate the kinetics of redox and non-redox transformations between hematite and magnetite, and the mechanisms of iron oxide transformation under hydrothermal conditions. The formation of euhedral crystals of hematite in 150 and 200 °C experiments, in which magnetite was used as the starting material, indicates that non-redox transformation of magnetite to hematite occurred within 24 h. The chemical composition of the experimental solutions was controlled by the non-redox transformation between magnetite and hematite throughout the experiments. While solution compositions were controlled by the non-redox transformation in the first 3 days in a 250 °C experiment, reductive dissolution of magnetite became important after 5 days and affected the solution chemistry. At 100 °C, the presence of maghemite was indicated in the first 7 days. Based on these results, equilibrium constants of non-redox transformation between magnetite and hematite and those of non-redox transformation between magnetite and maghemite were calculated. Our results suggest that the redox transformation of hematite to magnetite occurs in the following steps: (1) reductive dissolution of hematite to and (2) non-redox transformation of hematite and to magnetite.     

BIF-hosted iron mineral system: A review

The BIF-hosted iron ore system represents the world's largest and highest grade iron ore districts and deposits. BIF, the precursor to low- and high-grade BIF hosted iron ore, consists of Archean and Paleoproterozoic Algoma-type BIF (e.g., Serra Norte iron ore district in the Carajás Mineral Province), Proterozoic Lake Superior-type BIF (e.g., deposits in the Hamersley Province and craton), and Neoproterozoic Rapitan-type BIF (e.g., the Urucum iron ore district).The BIF-hosted iron ore system is structurally controlled, mostly via km-scale normal and strike-slips fault systems, which allow large volumes of ascending and descending hydrothermal fluids to circulate during Archean or Proterozoic deformation or early extensional events. Structures are also (passively) accessed via downward flowing supergene fluids during Cenozoic times.At the depositional site the transformation of BIF to low- and high-grade iron ore is controlled by: (1) structural permeability, (2) hypogene alteration caused by ascending deep fluids (largely magmatic or basinal brines), and descending ancient meteoric water, and (3) supergene enrichment via weathering processes. Hematite- and magnetite-based iron ores include a combination of microplaty hematite–martite, microplaty hematite with little or no goethite, martite–goethite, granoblastic hematite, specular hematite and magnetite, magnetite–martite, magnetite-specular hematite and magnetite–amphibole, respectively. Goethite ores with variable amounts of hematite and magnetite are mainly encountered in the weathering zone.In most large deposits, three major hypogene and one supergene ore stages are observed: (1) silica leaching and formation of magnetite and locally carbonate, (2) oxidation of magnetite to hematite (martitisation), further dissolution of quartz and formation of carbonate, (3) further martitisation, replacement of Fe silicates by hematite, new microplaty hematite and specular hematite formation and dissolution of carbonates, and (4) replacement of magnetite and any remaining carbonate by goethite and magnetite and formation of fibrous quartz and clay minerals.Hypogene alteration of BIF and surrounding country rocks is characterised by: (1) changes in the oxide mineralogy and textures, (2) development of distinct vertical and lateral distal, intermediate and proximal alteration zones defined by distinct oxide–silicate–carbonate assemblages, and (3) mass negative reactions such as de-silicification and de-carbonatisation, which significantly increase the porosity of high-grade iron ore, or lead to volume reduction by textural collapse or layer-compaction. Supergene alteration, up to depths of 200 m, is characterised by leaching of hypogene silica and carbonates, and dissolution precipitation of the iron oxyhydroxides.Carbonates in ore stages 2 and 3 are sourced from external fluids with respect to BIF. In the case of basin-related deposits, carbon is interpreted to be derived from deposits underlying carbonate sequences, whereas in the case of greenstone belt deposits carbonate is interpreted to be of magmatic origin. There is only limited mass balance analyses conducted, but those provide evidence for variable mobilization of Fe and depletion of SiO₂. In the high-grade ore zone a volume reduction of up to 25% is observed.Mass balance calculations for proximal alteration zones in mafic wall rocks relative to least altered examples at Beebyn display enrichment in LOI, F, MgO, Ni, Fe₂O_3total, C, Zn, Cr and P₂O₅ and depletions of CaO, S, K₂O, Rb, Ba, Sr and Na₂O. The Y/Ho and Sm/Yb ratios of mineralised BIF at Windarling and Koolyanobbing reflect distinct carbonate generations derived from substantial fluid–rock reactions between hydrothermal fluids and igneous country rocks, and a chemical carbonate-inheritance preserved in supergene goethite.Hypogene and supergene fluids are paramount for the formation of high-grade BIF-hosted iron ore because of the enormous amount of: (1) warm (100–200 °C) silica-undersaturated alkaline fluids necessary to dissolve quartz in BIF, (2) oxidized fluids that cause the oxidation of magnetite to hematite, (3) weakly acid (with moderate CO₂ content) to alkaline fluids that are necessary to form widespread metasomatic carbonate, (4) carbonate-undersaturated fluids that dissolve the diagenetic and metasomatic carbonates, and (5) oxidized fluids to form hematite species in the hypogene- and supergene-enriched zone and hydroxides in the supergene zone.Four discrete end-member models for Archean and Proterozoic hypogene and supergene-only BIF hosted iron ore are proposed: (1) granite–greenstone belt hosted, strike-slip fault zone controlled Carajás-type model, sourced by early magmatic (± metamorphic) fluids and ancient “warm” meteoric water; (2) sedimentary basin, normal fault zone controlled Hamersley-type model, sourced by early basinal (± evaporitic) brines and ancient “warm” meteoric water. A variation of the latter is the metamorphosed basin model, where BIF (ore) is significantly metamorphosed and deformed during distinct orogenic events (e.g., deposits in the Quadrilátero Ferrífero and Simandou Range). It is during the orogenic event that the upgrade of BIF to medium- and high-grade hypogene iron took place; (3) sedimentary basin hosted, early graben structure controlled Urucum-type model, where glaciomarine BIF and subsequent diagenesis to very low-grade metamorphism is responsible for variable gangue leaching and hematite mineralisation. All of these hypogene iron ore models do not preclude a stage of supergene modification, including iron hydroxide mineralisation, phosphorous, and additional gangue leaching during substantial weathering in ancient or Recent times; and (4) supergene enriched BIF Capanema-type model, which comprises goethitic iron ore deposits with no evidence for deep hypogene roots. A variation of this model is ancient supergene iron ores of the Sishen-type, where blocks of BIF slumped into underlying karstic carbonate units and subsequently experienced Fe upgrade during deep lateritic weathering.     

Kinetic and equilibrium Fe isotope fractionation between aqueous Fe(III) and hematite

Application of the Fe isotope system to studies of natural rocks and fluids requires precise knowledge of equilibrium Fe isotope fractionation factors among various aqueous Fe species and minerals. These are difficult to obtain at the low temperatures at which Fe isotope fractionation is expected to be largest and requires careful distinction between kinetic and equilibrium isotope effects. A detailed investigation of Fe isotope fractionation between [Fe^III(H₂O)₆]³⁺ and hematite at 98°C allows the equilibrium ⁵⁶Fe/⁵⁴Fe fractionation to be inferred, which we estimate at 10³lnα_{Fe(III)-hematite} = −0.10 ± 0.20‰. We also infer that the slope of Fe(III)-hematite fractionation is modest relative to 10⁶/T², which would imply that this fractionation remains close to zero at lower temperatures. These results indicate that Fe isotope compositions of hematite may closely approximate those of the fluids from which they precipitated if equilibrium isotopic fractionation is assumed, allowing inference of δ⁵⁶Fe values of ancient fluids from the rock record. The equilibrium Fe(III)-hematite fractionation factor determined in this study is significantly smaller than that obtained from the reduced partition function ratios calculated for [Fe^III(H₂O)₆]³⁺ and hematite based on vibrational frequencies and Mössbauer shifts by [Polyakov 1997] and [Polyakov and Mineev 2000], and Schauble et al. (2001), highlighting the importance of experimental calibration of Fe isotope fractionation factors. In contrast to the long-term (up to 203 d) experiments, short-term experiments indicate that kinetic isotope effects dominate during rapid precipitation of ferric oxides. Precipitation of hematite over ∼12 h produces a kinetic isotope fractionation where 10³lnα_{Fe(III)-hematite} = +1.32 ± 0.12‰. Precipitation under nonequilibrium conditions, however, can be recognized through stepwise dissolution in concentrated acids. As expected, our results demonstrate that dissolution by itself does not measurably fractionate Fe isotopes.     

Olivine dissolution in basaltic melt

The main purpose of this work is to understand and quantify diffusive and convective olivine dissolution in basaltic melt. Crystal dissolution and growth in a magma chamber is often accompanied by the descent or ascent of the crystal in the chamber due to gravity. The motion induces convection that enhances mass transport. Such convective dissolution and growth rates have not been quantified before. MgO diffusivity in the melt (D_MgO), MgO concentration of the interface melt (C₀) and the effective thickness of the compositional boundary layer (δ) are necessary parameters to model the convective dissolution. Experiments of non-convective olivine dissolution in a basaltic melt were conducted at 1271-1480 °C and 0.47-1.42 GPa in a piston-cylinder apparatus. At specific temperature and pressure conditions, multiple experiments of different durations show that the interface melt reaches near-saturation within 2 min. Therefore, diffusion, not interface reaction, is the rate-controlling step for non-convective olivine dissolution in basaltic melt. The compositional profile length and olivine dissolution distance are proportional to the square root of experimental duration, consistent with diffusive dissolution. D_MgO and C₀ are obtained from the experimental results. D_MgO displays Arrhenian dependence on temperature, but the pressure dependence is small and not resolved. C₀ increases with increasing temperature and decreases with increasing pressure. Comparison with literature data shows that D_MgO depends strongly on the initial melt composition, while C₀ does not. δ is estimated from fluid dynamics. D_MgO/δ, which characterizes the kinetic and dynamic aspects of convective crystal dissolution, is parameterized as a function of temperature, pressure, and olivine composition. Convective olivine dissolution rate in basaltic melt can be conveniently calculated from the model results. Application to convective crystal growth and xenolith digestion is discussed.     

Iron isotope fractionation during hydrothermal ore deposition and alteration

Iron isotopes fractionate during hydrothermal processes. Therefore, the Fe isotope composition of ore-forming minerals characterizes either iron sources or fluid histories. The former potentially serves to distinguish between sedimentary, magmatic or metamorphic iron sources, and the latter allows the reconstruction of precipitation and redox processes. These processes take place during ore formation or alteration. The aim of this contribution is to investigate the suitability of this new isotope method as a probe of ore-related processes. For this purpose 51 samples of iron ores and iron mineral separates from the Schwarzwald region, southwest Germany, were analyzed for their iron isotope composition using multicollector ICP-MS. Further, the ore-forming and ore-altering processes were quantitatively modeled using reaction path calculations. The Schwarzwald mining district hosts mineralizations that formed discontinuously over almost 300 Ma of hydrothermal activity. Primary hematite, siderite and sulfides formed from mixing of meteoric fluids with deeper crustal brines. Later, these minerals were partly dissolved and oxidized, and secondary hematite, goethite and iron arsenates were precipitated. Two types of alteration products formed: (1) primary and high-temperature secondary Fe minerals formed between 120 and 300 °C, and (2) low-temperature secondary Fe minerals formed under supergene conditions (<100 °C). Measured iron isotope compositions are variable and cover a range in δ⁵⁶Fe between −2.3‰ and +1.3‰. Primary hematite (δ⁵⁶Fe: −0.5‰ to +0.5‰) precipitated by mixing oxidizing surface waters with a hydrothermal fluid that contained moderately light Fe (δ⁵⁶Fe: −0.5‰) leached from the crystalline basement. Occasional input of CO₂-rich waters resulted in precipitation of isotopically light siderite (δ⁵⁶Fe: −1.4 to −0.7‰). The difference between hematite and siderite is compatible with published Fe isotope fractionation factors. The observed range in isotopic compositions can be accounted for by variable fractions of Fe precipitating from the fluid. Therefore, both fluid processes and mass balance can be inferred from Fe isotopes. Supergene weathering of siderite by oxidizing surface waters led to replacement of isotopically light primary siderite by similarly light secondary hematite and goethite, respectively. Because this replacement entails quantitative transfer of iron from precursor mineral to product, no significant isotope fractionation is produced. Hence, Fe isotopes potentially serve to identify precursors in ore alteration products. Goethites from oolitic sedimentary iron ores were also analyzed. Their compositional range appears to indicate oxidative precipitation from relatively uniform Fe dissolved in coastal water. This comprehensive iron isotope study illustrates the potential of the new technique in deciphering ore formation and alteration processes. Isotope ratios are strongly dependent on and highly characteristic of fluid and precipitation histories. Therefore, they are less suitable to provide information on Fe sources. However, it will be possible to unravel the physico-chemical processes leading to the formation, dissolution and redeposition of ores in great detail.     

Dissolution rates of phyllosilicates as a function of bacterial metabolic diversity

Weathering experiments using biotite and phlogopite in the presence of bacteria were conducted to better understand biotic dissolution kinetics and processes (proton- and ligand-promoted dissolution) under aerobic conditions. Miniature batch reactors (300 μl in microplate wells) were used at 24 °C for 3 days with and without bacterial strains. Abiotic experiments were performed with organic and nitric acids in order to calibrate the biotite-phlogopite chemical dissolution. An empirical model was used to fit the pH dependence for iron release rate (r_Fe) considering the influence of both protons and ligands from acidic to neutral conditions (pH ranging from 3 to 7): r_Fe=k_H(a_H+)^m+k_L(a_L)¹ where k is the apparent rate constant, a_H+ and a_L are the activities of protons and ligands, and m and l are the reaction orders. For both minerals in most cases at a given pH, the iron release rates in the presence of bacteria were in good agreement with rates determined by the chemical model and could be explained by a combination of proton- and ligand-promoted processes. Bacteria affect mineral dissolution and iron release rates through the quantities and nature of the organic acids they produce. Three domains were differentiated and proposed as biochemical models of mica dissolution: (1) below pH 3, only proton-promoted dissolution occurred, (2) in weakly acidic solutions both ligand- and proton-promoted mechanisms were involved, and (3) iron immobilization occured, at pH values greater than 4 for biotite and greater than 5 for phlogopite. This model allows us to distinguish the “weathering pattern phenotypes” of strains. Bacteria that are isolated from horizons poor in carbon appear more efficient at weathering micas than bacterial strains isolated from environments rich in carbon. Moreover, our results suggest that the mineral could exert a control on the release of organic acids and the “weathering pattern phenotypes” of bacteria.     

Study on the kinetics of iron oxide leaching by oxalic acid

The presence of iron oxides in clay or silica raw materials is detrimental to the manufacturing of high quality ceramics. Although iron has been traditionally removed by physical mineral processing, acid washing has been tested as it is more effective, especially for extremely low iron (of less than 0.1% w/w). However, inorganic acids such as sulphuric or hydrochloric acids easily contaminate the clay products with SO₄²⁻ and Cl⁻, and therefore should be avoided as much as possible. On the other hand, if oxalic acid is used, any acid left behind will be destroyed during the firing of the ceramic products. The characteristics of dissolution of iron oxides were therefore investigated in this study.The dissolution of iron oxides in oxalic acid was found to be very slow at temperatures within the range 25–60 °C, but its rate increases rapidly above 90 °C. The dissolution rate also increases with increasing oxalate concentration at the constant pH values set within the optimum range of pH2.5–3.0. At this optimum pH, the dissolution of fine pure hematite (Fe₂O₃) (105–140 μm) follows a diffusion-controlled shrinking core model. The rate expression expressed as 1 − (2 / 3)x − (1 − x)^2 / 3 where x is a fraction of iron dissolution was found to be proportional to [oxalate]^1.5.The addition of magnetite to the leach liquor at 10% w/w hematite was found to enhance the dissolution rate dramatically. Such addition of magnetite allows coarser hematite in the range 0.5–1.4 mm to be leached at a reasonable rate.     

Dissolution of illite in saline-acidic solutions at 25 °C

The dissolution rate of illite, a common clay mineral in Australian soils, was studied in saline-acidic solutions under far from equilibrium conditions. The clay fraction of Na-saturated Silver Hill illite (K_1.38Na_0.05)(Al_2.87Mg_0.46Fe³⁺_0.39Fe²⁺_0.28Ti_0.07)[Si_7.02Al_0.98]O₂₀(OH)₄ was used for this study. The dissolution rates were measured using flow-through reactors at 25 ± 1 °C, solution pH range of 1.0-4.25 (H₂SO₄) and at two ionic strengths (0.01 and 0.25 M) maintained using NaCl solution. Illite dissolution rates were calculated from the steady state release rates of Al and Si. The dissolution stoichiometry was determined from Al/Si, K/Si, Mg/Si and Fe/Si ratios. The release rates of cations were highly incongruent during the initial stage of experiments, with a preferential release of Al and K over Si in majority of the experiments. An Al/Si ratio >1 was observed at pH 2 and 3 while a ratio close to the stoichiometric composition was observed at pH 1 and 4 at the higher ionic strength. A relatively higher K⁺ release rate was observed at I = 0.25 in 2-4 pH range than at I = 0.01, possibly due to ion exchange reaction between Na⁺ from the solution and K⁺ from interlayer sites of illite. The steady state release rates of K, Fe and Mg were higher than Si over the entire pH range investigated in the study. From the point of view of the dominant structural cations (Si and Al), stoichiometric dissolution of illite occurred at pH 1-4 in the higher ionic strength experiments and at pH ?3 for the lower ionic strength experiments. The experiment at pH 4.25 and at the lower ionic strength exhibited lower R_Al (dissolution rate calculated from steady state Al release) than R_Si (dissolution rate calculated from steady state Si release), possibly due to the adsorption of dissolved Al as the output solutions were undersaturated with respect to gibbsite. The dissolution of illite appears to proceed with the removal of interlayer K followed by the dissolution of octahedral cations (Fe, Mg and Al), the dissolution of Si is the limiting step in the illite dissolution process. A dissolution rate law showing the dependence of illite dissolution rate on proton concentration in the acid-sulfate solutions was derived from the steady state dissolution rates and can be used in predicting the impact of illite dissolution in saline acid-sulfate environments. The fractional reaction orders of 0.32 (I = 0.25) and 0.36 (I = 0.01) obtained in the study for illite dissolution are similar to the values reported for smectite. The dissolution rate of illite is mainly controlled by solution pH and no effect of ionic strength was observed on the dissolution rates.     

Genesis of itabirite-hosted Au–Pd–Pt-bearing hematite-(quartz) veins, Quadrilátero Ferrífero, Minas Gerais, Brazil: constraints from fluid inclusion infrared microthermometry, bulk crush-leach analysis and U–Pb systematics

Fluid inclusions hosted in quartz and specular hematite from auriferous (jacutinga) and barren veins in the Quadrilátero Ferrífero (QF) have been studied using conventional and near infrared microscopy, respectively. The mineralization consists of veins that cross-cut metamorphosed iron formation (itabirite) of the Paleoproterozoic Itabira Group. The sample suite comprises hematite from veins from the low-strain domain in the W and SW of the study area, as well as hematite samples from the eastern high-strain domain in the central and NE parts of the QF. Halogen ratios of fluid inclusions in quartz and hematite from all studied deposits are consistent with a fluid evolved from dissolving and reprecipitating halite that was subsequently diluted. Fluid inclusions hosted in quartz and hematite are characterized by consistent Na/K ratios and considerable SO₄ contents, and suggest similar formation conditions and, perhaps, fluid origin from a common source. Na/K and Na/Li fluid mineral geothermometers indicate water–rock interaction at approximately 340±40°C. Hematites from the high-strain domain contain fluid inclusion assemblages of high-temperature aqueous-carbonic and multiphase high-salinity, high-temperature aqueous inclusions probably due to fluid immiscibility in the system H₂O–NaCl–CO₂. Fluid inclusions hosted in hematite from barren veins in the low-strain domain, as well as in hematite from jacutinga-type mineralization from the central part of the QF, only host multiphase aqueous fluid inclusions all showing narrow ranges of salinity (7.2–11.7 wt.% NaCl equiv.) and homogenization temperatures (148 to 229°C). Lower homogenization temperatures and the absence of CO₂-rich inclusions in specular hematite from these occurrences are attributed to carbonate precipitation and/or CO₂ escape due to cooling during fluid migration from the high- to the low-strain domain. Pb–Pb and U–Pb systematics of gold, hematite and hematite-hosted fluid inclusions in combination with geochemical evidence indicate distinct sources for Pd, Au, and Pb. The formation of specular hematite veins may be related to retrograde metamorphic fluids being released during the Brazilian orogenic cycle (600–700 Ma). The Pb isotopic characteristics of all samples are readily reconciled in a simple model that involves two different Paleoproterozoic or Archean source lithologies for lead and reflects contrasting depths of fluid percolation during the Brasiliano orogeny.     

First-principles study of the structural and isotopic properties of Al- and OH-bearing hematite

A series of hematite structures containing various amounts of aluminum and hydroxyl groups was modeled using first-principles methods based on the density functional theory. Evolution of the lattice parameters was quantified as a function of Al and H concentrations. The a and c lattice parameters decrease with the aluminum content and increase with the water content. This allows explaining experimental data reported for synthetic hematites, in particular the observed deviation from the Vegard’s rule. Investigation of several hydroxyl configurations associated with cationic vacancies suggests that the speciation of water also significantly affects the hematite structure. The ⁵⁷Fe and ¹⁸O reduced partition functions (β-factor) were determined. Results show a linear dependence of the iron and oxygen β-factors on the aluminum content. An incorporation of 18 mole % Al₂O₃ in hematite would increase the iron β-factor of about 0.6‰ and the oxygen β-factor of about 5.5‰ at 0 °C. This effect is sufficiently large to be measurable and to affect the interpretation of natural isotopic compositions. On the other hand, the effect of water is found to be negligible for the hydroxyl configuration investigated.     

Experimental and natural early diagenetic mobility of Sr and Mg in biogenic carbonates

Ammonium acetate dissolution experiments were performed on shell material of the modern bivalves, Crassostrea virginica (calcite) and Mercenaria mercenaria (aragonite). Their purpose was to determine the order of preferential dissolution of Sr, Mg and Ca; these results subsequently were compared with Sr, Mg and Ca data from other Recent as well as ancient mollusks.Results from these experiments suggest the following relative order of abundance of readily exchangeable Mg and Sr in biogenic carbonates: Mg(arag) > Mg(calc) > Sr(calc) > Sr(arag). It is apparent that incongruent dissolution of minerals with different solubilities cannot entirely explain the observed dissolution patterns for Sr, Mg and Ca in these biogenic carbonates. Secular changes in whole shell Mg and Sr concentrations for Recent and unrecrystallized fossil mollusks suggest an order of “ionic mobility” in diagenesis identical to the order of abundance for readily exchangeable ions found in the NH₄Ac dissolution experiments. It is concluded that this “ionic mobility” is due to a post mortem, early diagenetic (pre-recrystallization) approach to equilibrium with the surrounding chemical environment.