Experimental studies of oxygen and hydrogen isotope fractionations between precipitated brucite and water at low temperatures

Oxygen and hydrogen isotope fractionation factors between brucite and water were experimentally determined by chemical synthesis techniques at low temperatures of 15° to 120°C. MgCl₂, Mg3N₂, and MgO were used as reactants, respectively, to produce brucite in aqueous solutions. All of the synthesis products were identified by x-ray diffraction (XRD) for crystal structure and by scanning electron microscope (SEM) for morphology. It is observed that oxygen isotope fractionations between brucite and water are temperature dependent regardless of variations in aging time, the chemical composition, and pH value of solutions. Brucites derived from three different starting materials yielded consistent fractionations with water at the same temperatures. These suggest that oxygen isotope equilibrium has been achieved between the synthesized brucite and water, resulting in the fractionation equation of 10³lnα=1.56×10⁶/T²−14.1. When the present results for the brucite–water system are compared with those for systems of gibbsite–water and goethite–water, it suggests the following sequence of ¹⁸O-enrichment in the M−OH bonds of hydroxides: Al³⁺ − OH > Fe³⁺ − OH > Mg²⁺ − OH.Hydrogen isotope fractionations between brucite and water obtained by the different synthesis methods have also achieved equilibrium, resulting in the fractionation equation of 10³lnα=−4.88×10⁶/T²−22.5. Because of the pressure effect on hydrogen isotope fractionations between minerals and water, the present calibrations at atmospheric pressure are systematically lower than fractionations extrapolated from hydrothermal exchange experiments at high temperatures of 510° to 100°C and high pressures of 1060 to 1000 bar. Comparison of the present results with existing calibrations involving other low-temperature minerals suggests the following sequence of D-enrichment in hydroxyl-bearing minerals: Al³⁺ − OH > Mg²⁺ − OH > Fe³⁺ − OH.     

Pressure dependence of Néel transition in (Mg,Fe)O

Pressure dependence of Néel temperature (T _N) in (Mg_0.20Fe_0.80)O, (Mg_0.25Fe_0.75)O, and (Mg_0.30Fe_0.70)O was newly measured up to 1.14 GPa, using superconducting quantum interference device magnetometer and piston–cylinder-type pressure cell under hydrostatic condition. The dT _N/dP values of (Mg_0.20Fe_0.80)O, (Mg_0.25Fe_0.75)O, and (Mg_0.30Fe_0.70)O were determined as 4.0 ± 0.3, 2.7 ± 0.3, and 4.4 ± 0.4 K/GPa, respectively, in linear approximation; however, the T _N deviated from the linearity under nonhydrostatic conditions. The compositional dependence of dT _N/dP in (Mg_1?X Fe _X )O showed a rapid decay with increasing Mg components at X ≥ 0.75 and the trend ended at X = 0.70. The estimated Néel transition pressure at room temperature by extrapolating these linearities are very similar to the rhombohedral distortion determined by previous X-ray diffraction studies for X ≥ 0.75, which suggests that the rhombohedral phase of (Mg_1?X Fe _X )O (X ≥ 0.75) at room temperature is antiferromagnetic under hydrostatic conditions.     

Hydrogen isotope fractionation between kaolinite and water revisited

The equilibrium hydrogen isotope fractionation factor (α) between kaolinite and water in the temperature range 330 to 0°C is 1000 In α_kaol-water = −2.2 × 10⁶T⁻² − 7.7. This monotonic expression is based on a combination of experimental data with >75% of exchange and empirical calibrations. The previously proposed and widely accepted complex fractionation expression is considered to reflect the role of surface and intersite fractionation effects in the low percent of exchange experiments(Liu and Epstein, 1984), and incorrect δD water values for the empirical values (Lambert and Epstein, 1980). There is no measurable fractionation between dickite and kaolinite. The temperature dependence of the kaolinite-water hydrogen isotope fractionation factor can probably be used as a model for other phyllosilicate-water systems below 350°C.     

Oxygen isotope fractionation between synthetic aragonite and water: Influence of temperature and Mg concentration

Aragonite was precipitated in the laboratory at 0, 5, 10, 25, and 40 °C to determine the temperature dependence of the equilibrium oxygen isotope fractionation between aragonite and water. Forced CO₂ degassing, passive CO₂ degassing, and constant addition methods were employed to precipitate aragonite from supersaturated solutions, but the resulting aragonite-water oxygen isotope fractionation was independent of the precipitation method. In addition, under the experimental conditions of this study, the effect of precipitation rate on the oxygen isotope fractionation between aragonite and water was almost within the analytical error of ±∼0.13‰ and thus insignificant. Because the presence of Mg²⁺ ions is required to nucleate and precipitate aragonite from Na-Ca-Cl-HCO₃ solutions under these experimental conditions, the influence of the total Mg²⁺ concentration (up to ∼0.9 molal) on the aragonite-water oxygen isotope fractionation was examined at 25 °C. No significant Mg²⁺ ion effect, or oxygen isotope salt effect, was detected up to 100 mmolal total Mg²⁺ but a noticeable isotope salt effect was observed at ∼0.9 molal total Mg²⁺.On the basis of results of the laboratory synthesis experiments, a new expression for the aragonite-water fractionation is proposed over the temperature range of 0-40 °C:

1000lnα_{aragonite-water}=17.88±0.13(10³/T)-31.14±0.46     

Effects of cation substitutions in garnet and pyroxene on equilibrium oxygen isotope fractionations

High-grade metamorphic rocks were used to explore oxygen isotope fractionations between pyroxene and garnet, and to investigate the effects on fractionation factors of the cation substitutions Fe³⁺Al_?1 and Ca(Fe,Mg)_?1. Recrystallized, granulite facies (725 °C) wollastonite ores from the northern Adirondack highlands contain essentially only the minerals clinopyroxene (a Di–Hd solid solution)+garnet (a Grs–Adr solid solution)±wollastonite, and exhibit a systematic dependence of measured fractionations on the Fe³⁺ content of calcic garnet: Δ(Cpx–CaGrt)=(0.14±0.12)+(0.78±0.20)X_Adr and Δ(Wo–CaGrt)=(0.15±0.22)+(0.57±0.33)X_Adr. In eclogites formed at T ≤650 °C, measured compositions of Ca-poor garnet and omphacite combined with experimental data indicate that Ca-poor, Fe-rich garnet is enriched in ¹⁸O compared to both diopside and grossular: extrapolating to 1000 K, Δ(Alm–Di)≈c. 0.2 and Δ(Alm–Grs)≈c. 0.5. Orthopyroxene and clinopyroxene from Gore Mountain, New York, show a constant fractionation that is independent of rock type, as expected if they have the same closure temperature. These data imply Δ(Opx-Cpx)≈c. 0.7 at 1000 K. Measured fractionations among Ca-poor garnet, orthopyroxene, clinopyroxene and hornblende in the Gore Mountain rocks further indicate an ¹⁸O enrichment in Ca-poor garnet over Grs (≈c. 0.5 at 1000 K). The new measurements are indistinguishable from expected equilibrium values based on experiments for the minerals enstatite, diopside, grossular, wollastonite and feldspar, but consistently indicate a significant isotope effect for the simple octahedral cation substitutions Fe³⁺Al_?1 (Grs vs. Adr) and Ca(Fe,Mg)_?1 (Ca-poor garnet vs. Grs; Opx vs. Cpx). Neither cation substitution has been directly investigated for its effect on ¹⁸O/¹⁶O fractionation with experiments in silicates. Chemical characterization of minerals is required prior to petrological interpretation of oxygen isotope trends.     

Inter-mineral Fe isotope variations in mantle-derived rocks and implications for the Fe geochemical cycle

Iron isotope and major- and minor-element compositions of coexisting olivine, clinopyroxene, and orthopyroxene from eight spinel peridotite mantle xenoliths; olivine, magnetite, amphibole, and biotite from four andesitic volcanic rocks; and garnet and clinopyroxene from seven garnet peridotite and eclogites have been measured to evaluate if inter-mineral Fe isotope fractionation occurs in high-temperature igneous and metamorphic minerals and if isotopic fractionation is related to equilibrium Fe isotope partitioning or a result of open-system behavior. There is no measurable fractionation between silicate minerals and magnetite in andesitic volcanic rocks, nor between olivine and orthopyroxene in spinel peridotite mantle xenoliths. There are some inter-mineral differences (up to 0.2‰ in ⁵⁶Fe/⁵⁴Fe) in the Fe isotope composition of coexisting olivine and clinopyroxene in spinel peridotites. The Fe isotope fractionation observed between clinopyroxene and olivine appears to be a result of open-system behavior based on a positive correlation between the Δ⁵⁶Fe_{clinopyroxene-olivine} fractionation and the δ⁵⁶Fe value of clinopyroxene and olivine. There is also a significant difference in the isotopic compositions of garnet and clinopyroxene in garnet peridotites and eclogites, where the average Δ⁵⁶Fe_{clinopyroxene-garnet} fractionation is +0.32 ± 0.07‰ for six of the seven samples. The one sample that has a lower Δ⁵⁶Fe_{clinopyroxene-garnet} fractionation of 0.08‰ has a low Ca content in garnet, which may reflect some crystal chemical control on Fe isotope fractionation. The Fe isotope variability in mantle-derived minerals is interpreted to reflect subduction of isotopically variable oceanic crust, followed by transport through metasomatic fluids. Isotopic variability in the mantle might also occur during crystal fractionation of basaltic magmas within the mantle if garnet is a liquidus phase. The isotopic variations in the mantle are apparently homogenized during melting processes, producing homogenous Fe isotope compositions during crust formation.     

X-ray diffraction study of Fe2+-Mg order-disorder in orthopyroxene. Some kinetic results

Kinetic rates of Fe²⁺-Mg disordering in three orthopyroxenes (mean value of X_Fe = Fe²⁺/(Fe²⁺+Mg) = 0.175,0.482,0.770 respectively) have been determined employing heating experiments and single crystal X-ray structural refinements. Disordering rate constants \((\vec K)\) (550800° C) for two pyroxenes are given by: ln \((\vec K)\) = 27.107(±5.177)?32062(±783)T^?1(X_Fe = 0.175) ln \((\vec K)\) = 16.142(±0.057)?18227(±423)T^?1(X_Fe = 0.770) The distribution coefficients K_D (representing a steady state of disordering Fe_M2 + Mg_M1 ? Fe_M1 + Mg_M2) are given by: ln K_D = 5.016(±0.223)-7033(±1473) T^?1(X_Fe = 0.175) ln K_D = 1.988(±0.122)-3809(±913)T^?1(X_Fe = 0.770) These distribution coefficients provide the constraint of the disordering reaction on the value of the equilibrium constant for Fe²⁺-Mg order-disorder. Until the low temperature dependence of K_D is well constrained, the calculation of cooling rates of pyroxenes and host rocks cannot be done reliably.     

Mg-Fe云母化学成分的解释和分类(Ⅰ)——基本置换和黑云母平面

 中国东部花岗岩类141个Mg-Fe云母的化学成分将近90%的变化属于八面体层内的类质同象置换,置换矢量Mg ₁Fe⁺²和Fe_-3⁺²(R_Ⅵ⁺³)_-2□_Ⅵ组成了天然黑云母平面,大约80%的变化应当解释为基本置换8Mg ₁Fe⁺²+Fe_-3⁺²(R_Ⅵ⁺³)₂□_Ⅵ.这些是Mg-Fe云母在广泛的自然条件下表现出来的最主要的晶体化学关系。文中还提出了置换矢量的长度、分量和以及电价和三个参数,用以识别矿物化学成分变化的类质同象置换特征。     

Effect of grossular on garnet-biotite,Fe–Mg exchange reactions: evidence from garnet with mixed growth and diffusion zoning

Garnets that exhibit mixed growth and diffusion zoning are used to evaluate the effect of grossular content on garnet Fe–Mg exchange reactions. These garnets from the uppermost amphibolite-facies to granulite-facies gneiss of the Wissahickon Group, southeastern Pennsylvania, show variation in grossular content (0.035<X _Ca<0.14) but nearly constant Mg? (X _Mg/(X _Mg+X _Fe) and X _Mn through the interior indicating re-equilibration of garnet and matrix minerals with respect to iron, magnesium, and manganese. Mg? is not correlated with calcium content, evidence that the effect of calcium on garnet Fe–Mg exchange reactions is small or is offset by other interactions in almandine-rich garnets. In either case, the data presented here indicate that correction for calcium content of garnets in the application of garnet-biotite geothermometry to high-grade metapelites is unnecessary and may lead to an overestimate of peak temperature.     

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.     

Quantifying Li isotope fractionation during smectite formation and implications for the Li cycle

Tri-octahedral Li-Mg smectites (hectorites) were synthesized at temperatures ranging from 25 to 250 °C, in the presence of solutions highly enriched in lithium. After removing all the exchangeable lithium from the synthesized clays, Li isotope fractionation (Δ⁷Li_{clay-solution}) was determined. This fractionation was linked to Li incorporation into the structural octahedral site, substituting for Mg²⁺. As predicted, experimental Δ⁷Li_{clay-solution} inversely correlates with temperature, and ranges from −1.6‰ ± 1.3‰ at 250 °C to −10.0‰ ± 1.3‰ at 90 °C, and then stays relatively constant down to 25 °C. The relatively constant isotope fractionation factor below 90 °C may be due to high concentrations of edge octahedra in low crystallinity smectites. The isotopic fractionation factor (α), for a given temperature, does not depend on the solution matrix, nor on the amount of structural Li incorporated into the clay. Empirical linear laws for α as a function of 1/T (K) were inferred. Smectite Li contents and smectite-solution distribution coefficients (D_Li/Mg) increase with temperature, as expected for a substitution process. The fractions of dissolved Li incorporated into the smectite octahedral sites are small and do not depend on the duration of the experiment. In a seawater-like matrix solution, less Li is incorporated into the smectites, probably as a result of competition with dissolved Mg²⁺ ions for incorporation into the octahedral sites. The high Li contents observed in marine smectites are therefore best explained either by a significant contribution from basalts, by adsorption processes, or by the influence of seawater chemical composition on distribution coefficients. We also calculate, using present-day estimates of hydrothermal water and river fluxes, that a steady-state ocean would require a relatively large global clay-water Li isotope fractionation (−12‰ to −21‰). This study demonstrates the ability of laboratory experiments to quantify the impact of secondary phases on the Li geochemical cycle and associated isotope fractionations.     

Li-bearing, disordered Mg-rich tourmaline from a pegmatite-marble contact in the Austroalpine basement units (Styria, Austria)

Pale-blue to pale-green tourmalines from the contact zone of Permian pegmatites to mica schists and marbles from different localities of the Austroalpine basement units (Rappold Complex) in Styria, Austria, are characterized. All these Mg-rich tourmalines have small but significant Li contents, up to 0.29 wt% Li₂O, and can be characterized as dravite, with FeO contents of ?~?0.9–2.7 wt%. Their chemical composition varies from ^X (Na_0.67Ca_0.19?K_0.02?_0.12) ^Y (Mg_1.26Al_0.97Fe²⁺ _0.36Li_0.19Ti⁴⁺ _0.06Zn_0.01?_0.15) ^Z (Al_5.31?Mg_0.69) (BO₃)₃ Si₆O₁₈ ^V (OH)₃? ^W [F_0.66(OH)_0.34], with a?=?15.9220(3), c?=?7.1732(2) Å to ^X (Na_0.67Ca_0.24?K_0.02?_0.07) ^Y (Mg_1.83Al_0.88Fe²⁺ _0.20Li_0.08Zn_0.01Ti⁴⁺ _0.01?_0.09) ^Z (Al_5.25?Mg_0.75) (BO₃)₃ Si₆O₁₈ ^V (OH)₃? ^W [F_0.87(OH)_0.13], with a?=?15.9354(4), c?=?7.1934(4) Å, and they show a significant Al-Mg disorder between the Y and the Z sites (R1?=?0.013–0.015). There is a positive correlation between the Ca content and?<?Y-O?>?distance for all investigated tourmalines (r?≈?1.00), which may reflect short-range order configurations including Ca and Fe²⁺, Mg, and Li. The tourmalines have X_Mg (X_Mg?=?Mg/Mg?+?Fe_total) values in the range 0.84–0.95. The REE patterns show more or less pronounced negative Eu and positive Yb anomalies. In comparison to tourmalines from highly-evolved pegmatites, the tourmaline samples from the border zone of the pegmatites of the Rappold Complex contain relatively low amounts of total REE (~8–36 ppm) and Th (0.1–1.8 ppm) and have low La_N/Yb_N ratios. There is a positive correlation (r?≈?0.91) between MgO of the tourmalines and the MgO contents of the surrounding mica schists. We conclude that the pegmatites formed by anatectic melting of mica schists and paragneisses in Permian time. The tourmalines crystallized from the pegmatitic melt, influenced by the metacarbonate and metapelitic host rocks.     

Carbonate-hosted talc deposits in the contact aureole of an igneous intrusion (Hwanggangri mineralized zone,South Korea): geochemistry,phase relationships,and stable isotope studies

Many talc deposits occur in the Hwanggangri Mineralized Zone (HMZ) in dolomitic marbles of the Cambro-Ordovician Samtaesan Formation within 1 km of the contact with the Cretaceous Muamsa Granite. Talc commonly forms fine-grained, fibrous aggregates, or pseudomorphs after tremolite; abundant tremolite is included as impurities in the talc ore. Talc generally was derived from tremolite in calc-silicate rock within the dolomitic marble. Calc-silicate rock, consisting mainly of tremolite and diopside, was generated from silicic metasomatism during the prograde stage, which promoted decarbonation reactions until dolomite was exhausted locally. Hydrothermal alteration of calc-silicate rock to talc is marked by the addition of Mg and Si, and the leaching of Ca; Cr, Co, and Ni were relatively immobile during the retrograde stage. Contact metamorphism related to the granite intrusion generated the successive appearance of tremolite, diopside, and forsterite, or wollastonite-bearing assemblages in the marble, depending on the bulk rock composition. The X_CO2 content of the metamorphic fluids rose initially above X_CO2=0.6, and decreased steadily toward a water-rich composition with increasing temperature above 600 °C in the calcitic marble, while buffered reaction of the dolomitic marble occurred at higher X_CO2 conditions above 600 °C. Talc mineralization developed under metastable conditions with infiltration of large amounts of igneous fluids along a fault-shattered zone during the retrograde stage and is characterized by the loss of Ca²⁺ with the addition of Mg²⁺. Oxygen and carbon isotopic variations of carbonate and calc-silicate minerals are in agreement with theoretical relationships determined for decarbonation products of contact metamorphism. Talc formation temperatures obtained from oxygen isotope fractionation, T–X_CO2 relationships, and activity diagrams range from 380 to 400 °C.     

Manifestation of the light-induced drift effect in chemically peculiar stellar atmospheres

We have calculated the factor (ν _g ? ν _e )/ν _g in the temperature range T = 300–20 000 K for the ions Be⁺, Mg⁺, Ca⁺, C⁺ in atomic hydrogen and for the ions Mg⁺ in atomic argon using the known interaction potentials. Here ν _e and ν _g are the transport collision frequencies for excited- and ground-state particles respectively. Calculations have shown that at T = 10 000–20 000 K, typical temperatures of the atmospheres of chemically peculiar (CP) stars, the values |ν _g ? ν _e |/ν _g ≈ 0.1–0.2 can be reached for ions. This causes the light-induced drift (LID) velocity of ions up to ～0.1 cm/s in the atmospheres of CP stars with temperatures T < 10 000 K. Therefore the separation of chemical elements due to the LID of ions under the conditions of the atmospheres of such CP stars can be an order of magnitude more efficient in comparison with the separation caused by the radiation pressure. In the atmosphere of more hot stars (20 000 K > T > 10 000 K) it is possible to expect approximately identical magnitude of the LID effect and that of radiation pressure. In the very hot stars (T >20 000 K) the LID effect is manifested very weakly.     

Theoretical calculation of equilibrium Mg isotope fractionation between silicate melt and its vapor

Isotope fractionation during the evaporation of silicate melt and condensation of vapor has been widely used to explain various isotope signals observed in lunar soils, cosmic spherules, calcium–aluminum-rich inclusions, and bulk compositions of planetary materials. During evaporation and condensation, the equilibrium isotope fractionation factor (α) between high-temperature silicate melt and vapor is a fundamental parameter that can constrain the melt’s isotopic compositions. However, equilibrium α is difficult to calibrate experimentally. Here we used Mg as an example and calculated equilibrium Mg isotope fractionation in MgSiO₃ and Mg₂SiO₄ melt–vapor systems based on first-principles molecular dynamics and the high-temperature approximation of the Bigeleisen–Mayer equation. We found that, at 2500 K, δ²⁵Mg values in the MgSiO₃ and Mg₂SiO₄ melts were 0.141?±?0.004 and 0.143?±?0.003‰ more positive than in their respective vapors. The corresponding δ²⁶Mg values were 0.270?±?0.008 and 0.274?±?0.006‰ more positive than in vapors, respectively. The general \(\alpha - T\) equations describing the equilibrium Mg α in MgSiO₃ and Mg₂SiO₄ melt–vapor systems were: \(\alpha_{{{\text{Mg}}\left( {\text{l}} \right) - {\text{Mg}}\left( {\text{g}} \right)}} = 1 + \frac{{5.264 \times 10^{5} }}{{T^{2} }}\left( {\frac{1}{m} - \frac{1}{{m^{\prime}}}} \right)\) and \(\alpha_{{{\text{Mg}}\left( {\text{l}} \right) - {\text{Mg}}\left( {\text{g}} \right)}} = 1 + \frac{{5.340 \times 10^{5} }}{{T^{2} }}\left( {\frac{1}{m} - \frac{1}{{m^{\prime}}}} \right)\), respectively, where m is the mass of light isotope ²⁴Mg and m′ is the mass of the heavier isotope, ²⁵Mg or ²⁶Mg. These results offer a necessary parameter for mechanistic understanding of Mg isotope fractionation during evaporation and condensation that commonly occurs during the early stages of planetary formation and evolution.     

The experimental calibration of the iron isotope fractionation factor between pyrrhotite and peralkaline rhyolitic melt

A first experimental study was conducted to determine the equilibrium iron isotope fractionation between pyrrhotite and silicate melt at magmatic conditions. Experiments were performed in an internally heated gas pressure vessel at 500 MPa and temperatures between 840 and 1000 °C for 120-168 h. Three different types of experiments were conducted and after phase separation the iron isotope composition of the run products was measured by MC-ICP-MS. (i) Kinetic experiments using ⁵⁷Fe-enriched glass and natural pyrrhotite revealed that a close approach to equilibrium is attained already after 48 h. (ii) Isotope exchange experiments—using mixtures of hydrous peralkaline rhyolitic glass powder (∼4 wt% H₂O) and natural pyrrhotites (Fe_1 − xS) as starting materials— and (iii) crystallisation experiments, in which pyrrhotite was formed by reaction between elemental sulphur and rhyolitic melt, consistently showed that pyrrhotite preferentially incorporates light iron. No temperature dependence of the fractionation factor was found between 840 and 1000 °C, within experimental and analytical precision. An average fractionation factor of Δ ⁵⁶Fe/⁵⁴Fe_{pyrrhotite-melt} = −0. 35 ± 0.04‰ (2SE, n = 13) was determined for this temperature range. Predictions of Fe isotope fractionation between FeS and ferric iron-dominated silicate minerals are consistent with our experimental results, indicating that the marked contrast in both ligand and redox state of iron control the isotope fractionation between pyrrhotite and silicate melt. Consequently, the fractionation factor determined in this study is representative for the specific Fe²⁺/ΣFe ratio of our peralkaline rhyolitic melt of 0.38 ± 0.02. At higher Fe²⁺/ΣFe ratios a smaller fractionation factor is expected. Further investigation on Fe isotope fractionation between other mineral phases and silicate melts is needed, but the presented experimental results already suggest that even at high temperatures resolvable variations in the Fe isotope composition can be generated by equilibrium isotope fractionation in natural magmatic systems.     

Hydrogen incorporation in orthopyroxene: interaction of different trivalent cations

The incorporation of hydrogen into ferrosilite, Fe-bearing enstatite and orthopyroxene containing different trivalent cations (Cr³⁺ and Al³⁺, Cr³⁺ and Fe³⁺) was investigated experimentally at 25 kbar. Hydrogen concentration was determined by FTIR-spectroscopy on oriented crystal sections and by secondary ion mass spectroscopy, whereas Mößbauer spectroscopy and optical spectroscopy were used to characterise the valence state of Fe in orthopyroxene. Results suggest that hydrogen incorporation in ferrosilite is achieved by a similar mechanism as in pure enstatite. In Cr-bearing samples, however, hydrogen incorporation is reduced by the presence of other trivalent cations by an increased tendency to form Tschermaks substitutions, e.g. Si _T ⁴⁺ + Mg _M1 ²⁺ ? Al _T ³⁺ + Cr _M1 ³⁺ . Thus, hydrogen solubility in natural orthopyroxenes from the Earth’s mantle, containing significant amounts of Cr³⁺, Al³⁺, and Fe³⁺, may be much more limited than expected from their trivalent cation content, as a large fraction of the trivalent cations does not participate in H-incorporating reactions as 2 Mg _M1 ²⁺ ? M _M1 ³⁺ + V_M1 + H _i ⁺ .     

Influence of cation distribution on the optical absorption spectra of Fe3+-bearing spinel s.s.–hercynite crystals: evidence for electron transitions in VIFe2+–VIFe3+ clusters

Flux-grown Fe³⁺-bearing spinel s.s.–hercynite solid-solution crystals, (Mg _1-y Fe²⁺ _y )Al₂O₄ (0 < y≤ 1), have been investigated by means of electron microprobe technique and Mössbauer and electronic spectroscopy. Obtained results show that different electronic processes cause intense optical absorption bands in the near-infrared spectral region. In addition to an electronic d–d transition in single-ion ^IVFe²⁺, observed at 5200 cm^?1, intense and broad bands at 9500 and 14 500 cm^?1 are assigned to exchange-coupled pair (ECP) and intervalence charge-transfer (IVCT) transitions in ^VI Fe ^{2+ VI} Fe³⁺clusters, respectively. The net linear extinction coefficients of these bands (α) were calibrated against Fe²⁺ and Fe³⁺ concentrations and site distributions previously defined by combined microchemical, Mössbauer, and XRD structural refinement data. The following expressions were obtained: where α is measured in cm^?1 and concentrations are expressed in mol?l^?1. The present results show that optical absorption spectroscopy may be used as a probe to obtain high spatial resolution (?～ 10 μm) information on Fe²⁺ ordering as well as on Fe³⁺ concentrations in minerals belonging to the spinel group.     

Distribution of cations at two tetrahedral sites in Ca2MgSi2O7-Ca2Fe3+AlSiO7 series synthetic melilite and its relation to incommensurate structure

Synthetic melilites on the join Ca₂MgSi₂O₇ (åkermanite: Ak)-Ca₂Fe³⁺AlSiO₇ (ferrialuminium gehlenite: FAGeh) were studied using X-ray powder diffraction and ⁵⁷Fe Mössbauer spectroscopic methods to determine the distribution of Fe³⁺ between two different tetrahedral sites (T1 and T2), and the relationship between ionic substitution and incommensurate (IC) structure. Melilites were synthesized from starting materials with compositions of Ak₁₀₀, Ak₈₀FAGeh₂₀, Ak₇₀FAGeh₃₀ and Ak₅₀FAGeh₅₀ by sintering at 1,170–1,350 °C and 1 atm. The average chemical compositions and end-member components, Ak, FAGeh and Geh (Ca₂Al₂SiO₇), of the synthetic melilites were Ca_2.015Mg_1.023Si_1.981O₇ (Ak₁₀₀), Ca_2.017Mg_0.788Fe _0.187 ³⁺ Al_0.221Si_1.791O₇ (Ak₇₈FAGeh₁₉Geh₃), Ca_1.995Mg_0.695Fe _0.258 ³⁺ Al_0.318Si_1.723O₇ (Ak₆₉FAGeh₂₅Geh₆) and Ca_1.982Mg_0.495Fe _0.449 ³⁺ Al_0.519Si_1.535O₇ (Ak₄₉FAGeh₄₄Geh₇), respectively. Rietveld refinements using X-ray powder diffraction data measured using CuK _α -radiation at room temperature converged successfully with goodness-of-fits of 1.15–1.26. The refined Fe occupancies at the T1 and T2 sites and the Mg and Si contents determined by electron microprobe analysis gave the site populations of [0.788Mg + 0.082Fe³⁺ + 0.130Al]_T1[0.104Fe³⁺ + 0.104Al + 1.792Si]_T2 for Ak₇₈FAGeh₁₉Geh₃, [0.695Mg + 0.127Fe³⁺ + 0.178Al]_T1[0.132Fe³⁺ + 0.144Al + 1.724Si]_T2 for Ak₆₉FAGeh₂₅Geh₆ and [0.495Mg + 0.202Fe³⁺ + 0.303Al]_T1[0.248Fe³⁺ + 0.216Al + 1.536Si]_T2 for Ak₄₉FAGeh₄₄Geh₇ (apfu: atoms per formula unit), respectively. The results indicate that Fe³⁺ is distributed at both the T1 and the T2 sites. The mean T1–O distance decreases with the substitution of Fe³⁺ + Al³⁺ for Mg²⁺ at the T1 site, whereas the mean T2–O distance increases with substitution of Fe³⁺ + Al³⁺ for Si⁴⁺ at the T2 site, causing decrease in the a dimension and increase in the c dimension. However, in spite of the successful Rietveld refinements for the X-ray powder diffraction data measured using CuK _α-radiation at room temperature, each Bragg reflection measured using CuK _α1-radiation at room temperature showed weak shoulders, which were not observed in those measured at 200 °C. The Mössbauer spectra of the melilites measured at room temperature consist of two doublets assigned to Fe³⁺ at the T1 site and two or three doublets to Fe³⁺ at the T2 site, implying the existence of multiple T1 and T2 sites with different site distortions. These facts can be interpreted in terms of the IC structure in all synthetic melilites at room temperature, respectively. The results of Mössbauer analysis indicate that the IC structure in melilite is caused by not only known multiple T1 site, but also multiple T2 site at room temperature.