Fe FeS System Research Articles

Liquid properties in the Fe-FeS system under moderate pressure: Tool box to model small planetary cores

Liquid properties in the Fe-FeS system under moderate pressure: Tool box to model small planetary cores

Effect of Hydrogen on the Melting Temperature of Fe–FeS System at High Pressure

Effect of Hydrogen on the Melting Temperature of Fe–FeS System at High Pressure

Decomposition of Fe3S above 250 GPa

We have determined subsolidus phase relations in the Fe–FeS system up to 271 GPa using laser‐heated diamond‐anvil cell techniques. In situ synchrotron X‐ray diffraction (XRD) measurements performed at high pressure and high temperature demonstrate the coexistence of hexagonal close‐packed (hcp) Fe and tetragonal Fe3S up to 241 GPa and 2510 K. In contrast, the XRD data obtained above 250 GPa show that the hcp phase coexists with the CsCl (B2)‐type phase for three different Fe–S bulk compositions (10, 16, and 20 atm% S). Furthermore, chemical analyses using a scanning transmission electron microscope on a retrieved sample indicate that Fe3S sample decomposes into two phases at 271 GPa and 2530 K, consistent with the XRD data. Theory predicts the presence of extensive solid solution between Fe and FeS at inner core conditions, whereas our results suggest that the Fe–FeS system remains eutectic at least to 271 GPa.

The Fe-rich liquidus in the Fe–FeS system from 1 bar to 10 GPa

The composition and evolution of a metallic planetary core is determined by the behavior with pressure of the eutectic and the liquidus on the Fe-rich side of the Fe–FeS eutectic. New experiments at 6 GPa presented here, along with existing experimental data, inform a thermodynamic model for this liquidus from 1 bar to at least 10 GPa. Fe–FeS has a eutectic that becomes more Fe-rich but remains constant in T up to 6 GPa. The 1 bar, 3 GPa, and 6 GPa liquidi all cross at a pivot point at 1640 ± 5 K and FeS 37 ± 0.5 . This liquid/crystalline metal equilibrium is T– x-fixed and pressure independent through 6 GPa. Models of the 1 bar through 10 GPa experimental liquidi show that with increasing P there is an increase in the T separation between the liquidus and the crest of the metastable two-liquid solvus. The solvus crest decreases in T with increasing P. The model accurately reproduces all the experimental liquidi from 1 bar to 10 GPa, as well as reproducing the 0–6 GPa pivot point. The 14 GPa experimental liquidus ( Chen et al., 2008a,b) deviates sharply from the lower pressure trends indicating that the 0–10 GPa model no longer applies to this 14 GPa data.

Phase relationships of the Fe–FeS system in conditions up to the Earth's outer core

In situ X-ray diffraction experiments in the Fe–FeS system were performed up to 220 GPa and 3300 K using a laser-heated diamond anvil cell. Fe 3S and ε-Fe coexisted stably up to 220 GPa and 3300 K, and thus, Fe 3S is likely to be the stable S-bearing iron alloy under the Earth's core conditions. The solid iron ( ε-Fe) also contained 7.6(0.8) at.% of sulfur at 86 GPa and 2200 K. The amount of sulfur in the solid iron increased with increasing pressure at the eutectic temperatures. If the sulfur content obtained in this study is extrapolated to the conditions at the inner core, all the sulfur in the solid inner core can be stored in ε-Fe. The eutectic composition becomes nonsensitive to pressure and seems to be constant around 20 at.% of sulfur at pressures above 40 GPa. The pressure gradient of the melting curve of the Fe–FeS system is 13.4(0.7) K/GPa. Based on our results of melting relationship, the temperature at the core–mantle boundary should be greater than 2850(100) K, assuming that sulfur is the only light element in the Earth's liquid outer core.

Effect of Ni on Fe–FeS phase relations at high pressure and high temperature

A series of melting experiments in the Fe-rich portion of the Fe–Ni–S system have been conducted at 19–23 GPa and 800–1100 °C. The solubility of S in the Fe–Ni solid alloy and the eutectic melting in the Fe–Ni–S system were determined as a function of Ni content. The maximum S solubility in the Fe–Ni alloy is 2.7 wt.% at 20 GPa and the eutectic temperature. The eutectic melting temperature in the Fe–Ni(5wt.%)–S system is ~1000 °C lower than the melting point of pure Fe at 20 GPa. We also found that Ni can substitute Fe in the Fe3S structure to form (Fe,Ni)3S solid solutions up to at least a Fe/Ni atomic ratio of 0.5. Similar to melting behavior in the Fe–FeS system, the eutectic melting relations in the Fe–Ni–S system could produce inner and outer cores with the right light element balance to account for the density difference between the solid inner core and the liquid outer core.

Percolation threshold of iron-sulfide melts in olivine matrix: An experimental study with a centrifuging piston cylinder

There are several theories about metal-silicate segregation and the formation of the Earth’s core. One of them is the percolation of iron-sulfide melts through a solid silicate matrix, but the efficiency of such a mechanism is still under discussion. The main goal of the present investigation is to determine the percolation threshold for a simplified composition of the Earth core and silicate mantle and to verify whether buoyancy driven segregation of such metal-rich (anion-poor) melts is feasible. Starting materials are mixtures of natural olivine (S. Carlos) and iron-sulfide powder with the eutectic composition of the Fe–FeS system. The experimental apparati used are a standard end-loaded piston cylinder and a newly developed centrifuging piston cylinder. This rotating piston cylinder spins to a maximum speed of 2900 rpm (equivalent to an acceleration of 3000 g) at experimental conditions to 1.5 GPa and 1300 °C (standard 14 mm diameter, salt–pyrex assembly).

Viscosity of the Martian mantle and its initial temperature: Constraints from crust formation history and the evolution of the magnetic field

The rheology of the Martian mantle and the planet's initial temperature is constrained with thermal evolution models that include crust growth and test the conditions for magnetic field generation in the core. As observations we use the present-day average crustal thickness of 50–120 km as estimated from the Mars Global Surveyor gravity and topography data, the evidence for the crust being produced mostly early, with a rate declining from the Noachian to the Hesperian, and the evidence for an early magnetic field that likely existed for less than a billion years. We use the fact that the rate of crust growth is a function of temperature, which must be above the solidus in the sub-lithosphere mantle, and the mantle convection speed because the latter determines the rate at which melt can be replenished. The convection speed is a strong function of viscosity which, in turn, is a strong function of temperature and also of the water content of the mantle. We use a viscosity parameterization with a reference viscosity evaluated at 1600 K the value of which can be characteristic of either a dry or a wet mantle. We further consider the Fe–FeS phase diagram for the core and compare the core liquidus estimated for a sulphur content of 14% as suggested by the SNC meteorite compositions with the core temperatures calculated for our cooling models. Two data sets of the Fe–FeS eutectic temperature have been used that differ by about 200 K [Böhler, R., 1996. Fe–FeS eutectic temperatures at 620 kbar. Phys. Earth Planet. Inter. 96, 181–186; Fei, Y., Bertka, C.M., Finger, L.W., 1997. High-pressure iron–sulphur compound, Fe 3S 2, and melting relations in the Fe–FeS system. Science 275, 1621–1623] at Martian core–mantle boundary pressure and in the eutectic composition by 5 wt%. The differences in eutectic temperature and composition translate into a difference of about 400 K in liquidus temperature for 14 wt% sulphur. We find it premature to rule out specific mantle rheologies on the basis of the presently available crustal thickness and crust growth evidence. Rather a trade-off exists between the initial mantle temperature and the reference viscosity. Both a wet mantle rheology with a reference viscosity less than 10 20 Pas and a dry mantle rheology with a reference viscosity of 10 21 Pas or more can be acceptable if initial mantle temperatures between roughly 1700 and 2000 K are allowed. To explain the magnetic field history, the differences in liquidus temperatures matter. For a liquidus temperature of about 1900 K at the Martian core–mantle boundary as calculated from the Böhler et al. eutectic, a dry mantle rheology can best explain the lack of a present-day dynamo. For a liquidus temperature of about 1500 K at the core–mantle boundary as calculated from the Fei et al. eutectic all models are consistent with the observed lack of dynamo action. The reason lies with the fact that at 14 wt% S the Martian core would be close to the eutectic composition if the Fei et al. data are correct. As inner core growth is unlikely for an almost eutectic core, the early field would have been generated by a thermally driven dynamo. Together with the measured strength of the Martian crustal magnetization this would prove the feasibility of a strong thermally driven dynamo.

Connectivity of molten Fe alloy in peridotite based on in situ electrical conductivity measurements: implications for core formation in terrestrial planets

The connectivity of molten Fe–S in peridotite has been experimentally investigated by means of in situ electrical conductivity measurements at high temperatures and 1 GPa. Starting materials were powdered mixtures of peridotite KLB-1 with various amounts (0, 3, 6, 13, 19, 24 vol.%) of the 1 GPa eutectic composition in the Fe–FeS binary system. At temperatures above the eutectic point in the Fe–FeS system (∼980 °C) and below the solidus of KLB1 (∼1200 °C), molten Fe–S in a solid silicate matrix interconnects when the volume fraction is over ∼5%. Conductivity–temperature paths indicate that in the presence of partial silicate melting the connectivity of molten Fe–S in a peridotite matrix is inhibited. Based on observations of retrieved samples, the percolation threshold of Fe–S melts in the presence of low to moderate degrees of silicate melt is estimated at 13±2 vol.%. These results indicate that if the volume fraction of Fe-alloy in a planetesimal was initially greater than 5%, and if early heating by decay of radionuclides raised the temperature of the interior above the Fe-alloy melting point, initial metal segregation was controlled by permeable flow of molten iron alloy in a solid silicate matrix. These conditions were likely met by many terrestrial objects in the early solar nebula. Efficient removal of residual Fe-alloy (5 vol.%) from silicate requires high-degree melting of silicate so that metal can segregate as droplets. Giant impacts during the final stage of accretion of large planetary objects could supply the energy required for high-degrees of melting. Alternatively, if initial metal segregation were delayed until a planetary object grew to large size (∼1000 km in diameter), release of gravitational potential energy due to metal segregation could contribute enough heat to form a magma ocean.

Connectivity of molten Fe alloy in peridotite based on in situ electrical conductivity measurements: implications for core formation in terrestrial planets

The connectivity of molten Fe–S in peridotite has been experimentally investigated by means of in situ electrical conductivity measurements at high temperatures and 1 GPa. Starting materials were powdered mixtures of peridotite KLB-1 with various amounts (0, 3, 6, 13, 19, 24 vol.%) of the 1 GPa eutectic composition in the Fe–FeS binary system. At temperatures above the eutectic point in the Fe–FeS system (∼980 °C) and below the solidus of KLB1 (∼1200 °C), molten Fe–S in a solid silicate matrix interconnects when the volume fraction is over ∼5%. Conductivity–temperature paths indicate that in the presence of partial silicate melting the connectivity of molten Fe–S in a peridotite matrix is inhibited. Based on observations of retrieved samples, the percolation threshold of Fe–S melts in the presence of low to moderate degrees of silicate melt is estimated at 13±2 vol.%. These results indicate that if the volume fraction of Fe-alloy in a planetesimal was initially greater than 5%, and if early heating by decay of radionuclides raised the temperature of the interior above the Fe-alloy melting point, initial metal segregation was controlled by permeable flow of molten iron alloy in a solid silicate matrix. These conditions were likely met by many terrestrial objects in the early solar nebula. Efficient removal of residual Fe-alloy (5 vol.%) from silicate requires high-degree melting of silicate so that metal can segregate as droplets. Giant impacts during the final stage of accretion of large planetary objects could supply the energy required for high-degrees of melting. Alternatively, if initial metal segregation were delayed until a planetary object grew to large size (∼1000 km in diameter), release of gravitational potential energy due to metal segregation could contribute enough heat to form a magma ocean.

Geochemical constraints on the seismic properties of the lunar mantle

Abstract Thermodynamic modeling of phase relations and physical properties in the Na2O–TiO2–CaO–FeO–MgO–Al2O3–SiO2–Fe–FeS system and a Monte-Carlo inversion procedure have been used to calculate the density and velocity distribution in the mantle, and core radii of the Moon from the mass and moment of inertia constraints from recent Lunar Prospector gravity measurements. For the computation of the phase diagram for a given chemical composition, we have used a method of minimization of the total Gibbs free energy and self-consistent thermodynamic data for minerals and solid solutions. For the crustal thickness and density of 60 km and 3.0 g cm−3, the largest core radii are estimated to be 530 km for an FeS (troilite)-core (4.1% of total mass), 490 km for a eutectic Fe–FeS-core (3.4%), 445 km for an Fe–10 wt.% S-core (2.9%), and 350 km for an Fe-core (2%). The existence of an Fe–FeS-core of intermediate radius of 250–400 km is compatible with recent lunar core information and lend support to a chemically stratified lunar mantle. The sensitivity of density and seismic velocities to temperature, pressure and composition variations is examined. The modeling results indicate that we can determine the permissible density and velocity ranges in the zoned lunar mantle with confidence. The maximum and minimum density changes at depths of the lunar mantle, consistent with geochemical constraints and mass and moment of inertia requirements, are found to be: 3.22 g cm −3 ≤ρ≤3.44 g cm−3 at depths of 60–500 km and 3.34 g cm −3 ≤ρ≤3.52 g cm−3 at depths between 500 km and the core mantle boundary. Core-free models of the Moon are possible only if the lower mantle densities are between 3.49 and 3.52 g cm−3 ( 20 wt .% FeO wt.%). The calculated field of seismic velocities in the zoned lunar mantle shows that over a broad range of mantle compositions, temperatures, and densities the maximum variations are as follows: 4.23 km s −1 ≤V S ≤4.69 km s−1 and 7.41 km s −1 ≤V P ≤8.24 km s−1 at depths of 60–500 km, and 4.24 km s −1 ≤V S ≤4.64 km s−1 and 7.63 km s −1 ≤V P ≤8.39 km s−1 at depths between 500 km and the core mantle boundary. These petrologically and geophysically admissible ranges of velocities and densities may constitute a significant constraint for the analysis of internal structure of the Moon using the Apollo seismic data.

Core Sizes and Internal Structure of Earth's and Jupiter's Satellites

Models of the internal structure of the Moon and Galilean satellites of Jupiter—Io, Europa, and Ganymede—based on the geophysical (the mass and moment of inertia from recent Lunar Prospector and Galileo gravity measurements, and seismic velocities for the Moon), geochemical (the chemical composition of meteorites), and thermodynamic (modeling of phase relations and physical properties in the Na 2O–TiO 2–CaO–FeO–MgO–Al 2O 3–SiO 2–Fe–FeS system) constraints are constructed. Bulk composition of the Moon (enrichment in SiO 2, FeO, and refractories, and depletion in MgO) resembles neither the terrestrial matter nor the material of chondrites and achondrites. The radii of a lunar core are estimated to be 310–320 km for the Fe core, and 430–440 km for the eutectic Fe–FeS core. The Moon has the lowest Fe tot/Si weight ratio, 0.40–0.47, among the terrestrial planets, chondrites, and satellites of the outer Solar System. The correspondence between the density and moment of inertia values for bulk Io and rock–iron cores of Europa and Ganymede shows that their bulk compositions are, in general, similar and may be described by the composition close to L and LL chondrites. For such composition and density models, radii of the Fe core and eutectic Fe–FeS core are estimated to be R(Fe core)=590–630 km and R(Fe–FeS core)=830–875 km for Io, R(Fe core)=420–510 km and R(Fe–FeS core)=610–710 km for Europa, and R(Fe core)=580–650 km and R(Fe–FeS core)=820–900 km for Ganymede with an outer shell composed of the ice solid phases. The results of modeling support the hypothesis that Io, Europa, and Ganymede have a massive metallic core in which a magnetic field may be produced. Thickness of an ice–liquid outer shell in Europa is estimated to be 120–140 km (7–8% of total mass); thickness of a solid ice outer shell in Ganymede is expected to be 890–920 km (46–48% of total mass). Models showing the outer shell of Ganymede to consist of a mixture of water and ice are considered, and the thickness of an outer water-ice shell (780–850 km) and Fe–FeS core radii are estimated. The possibility of an inner liquid–water layer beneath the icy surface of Ganymede cannot be ruled out if the bulk composition of its rock–iron core were close to the composition of LL (rather than L) chondrites.

In situ measurement of viscosity of liquids in the Fe-FeS system at high pressures and temperatures

The viscosity of liquid FeS and Fe-FeS eutectic was measured at pressures between 0.5 and 5.0 GPa using a synchrotron-based falling sphere technique. We obtain viscosities of 2 x 10(-2) to 4 x 10(-3) Pa-s in FeS at 1450 to 1700 degreesC and 2 x 10(-2) to 8 x 10(-3) Pa-s in Fe-S-eut at 1150 to 1380 degreesC. These results are consistent with recent viscosity measurements in Fe-S-eut at 5 to 7 GPa (Urakawa, in preparation), measured diffusivities (Dobson 2000) and ab initio simulated viscosity (Vocadlo et al. 2000). The results are also similar to the values for pure iron at low pressure (Shimoji and Itami 1986). A systematic increase in viscosity and activation energy is seen with increasing sulfur content. Interpolation between the data presented yields a viscosity of 1.4 x 10(-2) Pa-s for an outer core composition with similar to 10 wt% S at the melting temperature. There is good evidence of homologous behavior for Fe-S liquids which implies that the liquid alloy at the inner core boundary may have a similar viscosity.

Resemblance and difference between the constitution of the Moon and Io

The Moon and Io are comparable in the mean density, the mass and moment of inertia in spite of the principle differences in bulk compositions and masses of central planets. The Moon is the only extraterrestrial body for which we have information about the mantle seismic velocities as a function of depth. Volcanic activity on the Moon was completed in the early stage of its evolution. Io is the most active of all known satellites and it is much richer in volatiles and iron than the Moon. In this work, internally consistent models of the constitution of the Moon and Io based on the geophysical (the seismic velocities, the mass and moment of inertia) and geochemical (chemical and phase composition) constraints are compared. Thermodynamic modeling of the phase relations and physical properties in the CaO–FeO–MgO–Al 2O 3–SiO 2–Fe–FeS system was used to estimate the bulk composition, the density distribution in the mantle and radius of a core from the geophysical data. The solution of the inverse problem is based on the minimization of the deviations between the calculated and observed geophysical parameters. The smallest and largest core radii are estimated to be: 500–655 km for an Fe core (6–10%) and 810–1080 km for an FeS core (12–28% of total mass) for Io; 290–345 km for an Fe core (1.1–1.9%) and 450–530 km for an FeS core (2.4–4% of total mass) for the Moon. The atomic ratios of total iron to silicon for Io are estimated to be (Fe/Si)=0.40–0.66 in comparison with 0.22–0.24 for the Moon, 0.51–0.59 for the LL and L chondrites, and 0.74–0.8 for CV3 and CM2 meteorites. The current results show that the proposed Io’s compositions similar to that of carbonaceous chondrites do not satisfy to bulk Fe/Si ratios inferred from the geophysical constraints. The bulk composition of Io is estimated to be close to that of the L and LL chondrites having 8–13 wt.% of iron and iron sulfide; a change of the core composition from pure iron to pure iron sulfide leads to a change of the core radius from 600 to 625 km (8–9%) to 930–965 km (18–20%); the radius of the eutectic Fe–FeS-core is estimated to be 840–870 km (14–16% of Io’s mass; (Fe/Si) at=0.55–0.57). If Io does have such a composition, then the core is probably Fe or Fe-rich, whereas a large FeS core is excluded by the composition of the L and LL chondrites. A comparison of estimated bulk chemistry of Io with much more reliable composition of the Moon shows that Earth’s and Jupiter’s satellites have absolutely different bulk compositions.

The exploration of Martian interiors using the spheroidal oscillation method

The diagnostic possibilities of the free oscillation method for probing the global pattern of the structure of the Mars interior (the radius and aggregate (solid or liquid) state of the core, the crust thickness, the depth of the beginning of the olivine-spinel phase transition, and the width of this transition) are investigated. A set of two-layer models for different compositions of the FeFeS system are used to investigate the aggregate state of the core and its radius. The effects of the crust thickness and the depth of the beginning of the olivine-spinel phase transition are investigated for four- and five-layer models. The high-(HTM) and low-(LTM) temperature models are considered. For a large number of models spheroidal oscillations are calculated for n= 2–200. Dispersion curves are constructed for Rayleigh waves, providing a valuable tool for diagnosing the core thickness. The oscillation periods with low values n⩽7 depend substantially on the core parameters, while the spheroidal modes with n ≈ 7–10 diagnose reliably the position and structure of the transition zone.

The thermal regime of Venus

Models for the thermal evolution of Venus are constructed numerically. The mantle is proposed to consist of two independently convecting layers (upper and lower mantles); the floating crust keeps the temperature of the upper surface of the convecting mantle near 1200°C and the core may solidify at temperatures below the liquidus. Previously, these models were calculated with the usual parameterization of convection based mostly on investigations of constant-viscosity convection. Now this parameterization is modified to take into account new numerical investigations of convection in media with complicated rheology. The thermal evolution of Venus is assumed to have begun 4.6 billion years (b.y.) ago. It can be divided into three periods: (1) adjustment of the upper mantle to the thermal regime of the lower mantle (∼0.5 b.y.); (2) transition of the entire mantle to the asymptotic regime (∼3–4 b.y.); (3) asymptotic regime. Parameters of the convecting planet in the asymptotic regime do not depend on initial conditions (the planet “forgets” its initial state) and are found analytically. The present-day heat flux is equal to ∼50 erg cm −2s −1. The results are close to those obtained previously. Comparison of the present thermal models of Venus and Earth with the recent studies of melting of the Fe FeS system suggest that the absence of a dipole magnetic field is closely connected to the absence of solidification of the Venusian core at present. The position of the iron triple point (γ-ε-liquid) makes this conclusion very probable. The thermal regime of the Venusian crust is also discussed. Convection in the lower part of the crust is shown to play a role in certain regions with specific crustal composition. The prevailing mechanism of heat transfer to the surface is advection by magmas produced by melting the lower layers of the crust.

The distribution of trace quantities of germanium between metal, silicate and sulflde phases

The distribution of trace quantities of Ge between metal, silicate and sulfide phases has been studied hydrothermally at controlled P O 2 by means of oxygen buffer techniques. Radioactive Ge 68-labeled olivine of varying Mg and Fe composition and Ge 68-labeled iron were synthesized for this study. At 900°C and 500 bars, Ge is strongly siderophile in the iron-olivine system even at the equilibrium P O 2 of iron-wüstite. Similar strongly siderophilic behavior of Ge was also observed in the Fe-FeS system at this temperature and pressure. Thermodynamic data indicate that the free energy of formation of germanium dioxide lies between that of the iron oxide and nickel oxide at all temperatures. In the presence of metallic iron, the reduction of GeO 2 to Ge and the formation of Fe-Ge alloy are spontaneous reactions. If nickel is added to the system, the metal phase (Fe-Ni alloy) can exist at a higher P O 2 than pure iron. At the low Ni content of about 10 per cent, as found in most iron meteorites, the change in P O 2 Of metal-wüstite from iron-wüstite is small and is unlikely to affect the siderophilic behavior of Ge. A discussion is given to indicate that the diffusion of Ge out of the metal phase is possible if the Ni content of the metal is very high. From these results, it seems unlikely that a simple one-stage separation of metal from silicate would cause the extremely low Ge content observed in some iron meteorites e.g. iron meteorites of chemical groups IVA and IVB. Other possible fractionation processes involving Ge are also discussed.