Core-mantle Boundary Pressure Research Articles

The direction of core solidification in asteroids: Implications for dynamo generation

Paleomagnetic studies of meteorites over the past two decades have revealed that the cores of multiple meteorite parent bodies, including those of certain chondritic groups, generated dynamo fields as they crystallised. However, uncertainties in the direction and mode of core solidification in asteroid-sized bodies have meant using the timings and durations of these fields to constrain parent body properties, such as size, is challenging. Here, we use updated equations of state and liquidus relationships for Fe-FeS liquids at low pressures to calculate the locations at which solids form in these cores. We perform these calculations for core-mantle boundary (CMB) pressures from 0 - 2 GPa, and Fe-FeS liquid concentrations on the iron-rich side of the eutectic, as well as two values of iron thermal expansivity that cover the measured uncertainties in this parameter, and adiabatic and conductive cooling of these cores. We predict inward core crystallisation from the CMB in asteroids due to their low < 0.5 GPa pressures regardless of the uncertainties in other key core parameters. However, due to low internal pressures in these cores, remelting of any iron snow, as proposed to generate Ganymede’s present-day field, may be unlikely as the cores are approximately isothermal. Therefore a different mode of inward core solidification is possibly required to explain compositionally-driven dynamo action in asteroids. Additionally, we identify possible regimes at higher > 0.6 − 2 GPa pressures in which crystallisation can occur concurrently at the CMB and the centre.

Lorenz number and transport properties of Fe: Implications to the thermal conductivity at Earth's core-mantle boundary

Abstract The electrical resistivity (ρ) and thermal conductivity (κ) of the Earth's core compositions are essential parameters for constraining the core's thermal state, the inner core age, and the evolutionary history of the geodynamo. However, controversies persist between experimental and computational results regarding the electronic transport properties (ρ and κ) of the Earth's core. Iron is the major element in the core, and its transport properties under high pressure and high temperature conditions are crucial for understanding the core's thermal state. We measured the ρ values of solid iron using the four-wire van der Pauw method at 300 K and pressures ranging from 3 to 26 GPa within a multi-anvil press. For comparison, we calculated the ρ and κ values of hexagonal close-packed (hcp) iron at conditions of 300-4100 K and 22-136 GPa using the first-principles molecular dynamics (FPMD) method. Our calculations generally align with prior studies, indicating that the electrical resistivity of solid hcp iron at Earth's core-mantle boundary (CMB) conditions is ~76-83 μΩ∙cm. The resistivity of hcp iron changes small as it melts from solid to liquid at pressures from 98 to 134 GPa. The impact of temperature and pressure on the Lorenz numbers of solid hcp iron are investigated according to our calculation results and previous studies. Under the CMB's pressure conditions, the κ of hcp iron initially decreases with increasing temperature and subsequently increases. The electron-electron scattering plays a dominant role at low temperatures and causes the decrease in κ. At high temperatures, the increase of electronic specific heat significantly increases the Lorentz number and κ. Overall, we estimate the κ of solid hcp iron at CMB's condition to be 114 ± 6 W/m/K, slightly lower than the room temperature value of 129 ± 9 W/m/K at the same pressure. Our model shows that a 0-525 km thickness of a thermally stratified layer may exist beneath the Earth's CMB depending on the core's heat flow and thermal conductivity.

Thermal conductivity of iron and nickel during melting: Implication to the planetary liquid outer core

We report the measurements of the thermal conductivity (κ) of iron (Fe) and nickel (Ni) at high pressures and high temperatures near the core–mantle boundary (CMB) of Mercury and Mars. κ values are estimated from the temperature measurements across the sample surface in a laser-heated diamond anvil cell (LHDAC) and using the COMSOL software. In both metals, values of κ drop suddenly at the melting temperature and can be used as another helpful technique to estimate melting temperatures at high pressures. We have obtained the thermal conductivity values of Fe and Ni at the CMB of Mercury to be about (60–70) ± 28 W m−1 K−1 and (65–70) ± 28 W m−1 K−1, respectively. Similarly, in the case of Mars CMB pressure and temperature conditions, the thermal conductivities of Fe and Ni are found to be 80 ± 32 W m−1 K−1 and (65–70) ± 28 W m−1 K−1, respectively. Using the obtained thermal conductivity values of Fe and Ni, we have estimated the adiabatic heat conduction at the outer core of Mercury and Mars, to be about (59–69) ± 28 mW m−2 and (61–72) ± 29 mW m−2, respectively. Both these values are much larger than the values predicted recently from the estimation of electronic thermal conductivity.

Prediction of the melting curve and phase diagram for CaO using newly developed interatomic potentials

A comprehensive investigation of the melting curve and pressure–temperature (P–T) phase diagram for CaO, a candidate mineral in the Earth's lower mantle, is performed via atomistic simulations applying newly developed interatomic potentials. The efficiency and reliability of the new potentials under high temperature and pressure are verified with lattice dynamics, molecular dynamics, and first–principles approaches, where the fitting values of the potentials are extracted from the lattice parameters, elastic constants, and energy differences of rock–salt (B1) and cesium chloride (B2) phases in CaO. The melting temperatures of CaO with B1 structure at ambient pressure obtained from the single–phase method, void method, and two–phase method are 3080 K, 2775 K, and 2767 K, respectively. The structure and diffusion are evaluated by radial distribution functions, mean square displacements, and self–diffusion coefficients in the solid and liquid states for CaO. Simultaneously, the P–T phase diagram of CaO is calculated up to the core–mantle boundary pressure of 135 GPa, in which the Clausius–Clapeyron relation is introduced to determine the slope of the phase boundary between B1 and B2 phases. Other physical properties, including phase transition pressure, thermal expansion coefficient, isovolumic heat capacity, and entropy, are investigated. The limitations of the new potentials are discussed.

Pressure-induced C23–C37 transition and compression behavior of orthorhombic Fe2S to Earth’s core pressures and high temperatures

Abstract The phase stability of orthorhombic Fe2S was explored to 194 GPa and 2500 K using powder and multigrain synchrotron X-ray diffraction techniques. Between 30 and 120 GPa, a C23-like (Co2P, Pnma, Z = 4) Fe2S structure is observed and determined to exhibit a highly compressible a axis. A softening of the a axis occurs between 120 and 150 GPa and a relative stiffening of the b and c axes accompanies this compressibility change. Above 150 GPa, the a axis stiffens as the b and c axes soften, and a C37-like (Co2Si, Pnma, Z = 4) Fe2S unit cell is measured. On the basis of these changes in unit cell geometry, a pressure-induced C23–C37 Fe2S phase transition is inferred between 120–150 GPa. The C23 and C37 (Pnma, Z = 4) structures are closely related and share the same site symmetries. Forming the C37 structure from the C23 structure requires a shortening of the a axis and lengthening of the b and c axes accompanied by a four- to fivefold coordination change. The softening of the a axis above 120 GPa may therefore indicate the onset of a coordination change, and the final compressibility change above 150 GPa may mark the completion of this phase transition. The presented pressure-temperature (P-T) stabilities of C23 and C37 structures of Fe2S are in agreement with and resolve the differing observations of two previous studies (Tateno et al. 2019; Zurkowski et al. 2022). As C37 Fe2S is observed to core-mantle boundary pressures and high temperatures, the C37 Fe2S density profile through Earth’s outer core was determined by fitting the C23 Fe2S equation of state (&amp;lt;120 GPa) and applying a 1.6% volume reduction based on the C37 Fe2S volume residuals to this fit. Comparing the density of liquid C37 Fe2S with that of liquid hcp-Fe (Dewaele et al. 2006) and the seismologically determined density deficit of Earth’s core (Irving et al. 2018), 13.9 ± 1.5 wt% and 8.6 ± 0.8 wt% sulfur are required to match the densities at the CMB and ICB, respectively, for a purely Fe-S core.

Effects of core formation on the Hf–W isotopic composition of the Earth and dating of the Moon-forming impact

Earth's core formation set the initial compositions of the core and mantle. Various aspects of core formation, such as the degree of metal–silicate equilibration, oxygen fugacity, and depth of equilibration, have significant consequences for the resulting compositions, yet are poorly constrained. The Hf–W isotopic system can provide unique constraints on these aspects relative to other geochemical or geophysical methods. Here we model the Hf–W isotopic evolution of the Earth, improving over previous studies by combining a large number of N-body simulations of planetary accretion with a core formation model that includes self-consistent evolution of oxygen fugacity and a partition coefficient of tungsten that evolves with changing pressure, temperature, composition, and oxygen fugacity. The effective average fraction of equilibrating metal is constrained to be k>0.2 for a range of equilibrating silicate masses (for canonical accretion scenarios), and is likely <0.55 if the Moon formed later than 65 Ma. These values of k typically correspond to an effective equilibration depth of ∼0.5–0.7× the evolving core–mantle boundary pressure as the planet grows. The average mass of equilibrating silicate was likely at least 3× the impactor's silicate mass. Equilibration temperature, initial fO2, initial differentiation time, semimajor axis, and planetary mass (above ∼0.9 M⊕) have no systematic effect on the 182W anomaly, or on fHf/W (except for fO2), when applying the constraint that the model must reproduce Earth's mantle W abundance. There are strong tradeoffs between the effects of k, equilibrating silicate mass, depth of equilibration, and timing of core formation, so the terrestrial Hf–W isotopic system should be interpreted with caution when used as a chronometer of Earth's core formation. Because of these strong tradeoffs, the Earth's tungsten anomaly can be reproduced for Moon-forming impact timescales spanning at least 10–175 Ma. Early Moon formation ages require a higher degree of metal–silicate equilibration to produce Earth's 182W anomaly.

First principles and Debye model study of the thermodynamic, electronic and optical properties of MgO under high-temperature and pressure

First principles and quasi-harmonic Debye model have been used to study the thermodynamic properties, enthalpies, electronic and optical properties of MgO up to the core–mantle boundary (CMB) condition (137 GPa and 3700 K). Thermodynamic properties calculation includes thermal expansion coefficient and capacity, which have been studied up to the CMB pressure (137 GPa) and temperature (3700 K) by the Debye model with generalized gradient approximation (GGA) and local-density approximation (LDA). First principles with hybrid functional method (PBE0) has been used to calculate the electronic and optical properties under pressure up to 137 GPa and 0 K. Our results show the Debye model with LDA and first principles with PBE0 can provide accurate thermodynamic properties, enthalpies, electronic and optical properties. Calculated enthalpies show that MgO keep NaCl (B1) structure up to 137 GPa. And MgO is a direct bandgap insulator with a 7.23 eV calculated bandgap. The bandgap increased with increasing pressure, which will induce a blue shift of optical properties. We also calculated the density of states (DOS) and discussed the relation between DOS and band, optical properties. Equations were used to fit the relations between pressure and bandgaps, absorption coefficient ([Formula: see text]([Formula: see text])) of MgO. The equations can be used to evaluate pressure after careful calibration. Our calculations can not only be used to identify some geological processes, but also offer a reference to the applications of MgO in the future.

Fate of MgSiO3 melts at core–mantle boundary conditions

One key for understanding the stratification in the deep mantle lies in the determination of the density and structure of matter at high pressures, as well as the density contrast between solid and liquid silicate phases. Indeed, the density contrast is the main control on the entrainment or settlement of matter and is of fundamental importance for understanding the past and present dynamic behavior of the deepest part of the Earth's mantle. Here, we adapted the X-ray absorption method to the small dimensions of the diamond anvil cell, enabling density measurements of amorphous materials to unprecedented conditions of pressure. Our density data for MgSiO3 glass up to 127 GPa are considerably higher than those previously derived from Brillouin spectroscopy but validate recent ab initio molecular dynamics simulations. A fourth-order Birch-Murnaghan equation of state reproduces our experimental data over the entire pressure regime of the mantle. At the core-mantle boundary (CMB) pressure, the density of MgSiO3 glass is 5.48 ± 0.18 g/cm(3), which is only 1.6% lower than that of MgSiO3 bridgmanite at 5.57 g/cm(3), i.e., they are the same within the uncertainty. Taking into account the partitioning of iron into the melt, we conclude that melts are denser than the surrounding solid phases in the lowermost mantle and that melts will be trapped above the CMB.

Melting of MORB at core–mantle boundary

We investigated the melting properties of natural mid-ocean ridge basalt (MORB) up to core–mantle boundary (CMB) pressures using laser-heated diamond anvil cell. Textural and chemical characterizations of quenched samples were performed by analytical transmission electron microscopy. We used in situ X-ray diffraction primarily for phase identification whereas our melting criterion based on laser power versus temperature plateau combined with textural analysis of recovered solidus and subsolidus samples is accurate and unambiguous. At CMB pressure (135 GPa), the MORB solidus temperature is 3970(±150) K. Quenched melt textures observed in recovered samples indicate that CaSiO3 perovskite (CaPv) is the liquidus phase in the entire pressure range up to CMB. The partial melt composition derived from the central melt pool is enriched in FeO, which suggests that such melt pockets may be gravitationally stable at the core mantle boundary.

Accretion and differentiation of the terrestrial planets with implications for the compositions of early-formed Solar System bodies and accretion of water

In order to test accretion simulations as well as planetary differentiation scenarios, we have integrated a multistage core–mantle differentiation model with N-body accretion simulations. Impacts between embryos and planetesimals are considered to result in magma ocean formation and episodes of core formation. The core formation model combines rigorous chemical mass balance with metal–silicate element partitioning data and requires that the bulk compositions of all starting embryos and planetesimals are defined as a function of their heliocentric distances of origin. To do this, we assume that non-volatile elements are present in Solar System (CI) relative abundances in all bodies and that oxygen and H2O contents are the main compositional variables. The primary constraint on the combined model is the composition of the Earth’s primitive mantle. In addition, we aim to reproduce the composition of the martian mantle and the mass fractions of the metallic cores of Earth and Mars. The model is refined by least squares minimization with up to five fitting parameters that consist of the metal–silicate equilibration pressure and 1–4 parameters that define the starting compositions of primitive bodies. This integrated model has been applied to six Grand Tack N-body accretion simulations. Investigations of a broad parameter space indicate that: (1) accretion of Earth was heterogeneous, (2) metal–silicate equilibration pressures increase as accretion progresses and are, on average, 60–70% of core–mantle boundary pressures at the time of each impact, and (3) a large fraction (70–100%) of the metal of impactor cores equilibrates with a small fraction of the silicate mantles of proto-planets during each core formation event. Results are highly sensitive to the compositional model for the primitive starting bodies and several accretion/core-formation models can thus be excluded. Acceptable fits to the Earth’s mantle composition are obtained only when bodies that originated close to the Sun, at <0.9–1.2AU, are highly reduced and those from beyond this distance are increasingly oxidized. Reasonable concentrations of H2O in Earth’s mantle are obtained when bodies originating from beyond 6–7AU contain 20wt% water ice (icy bodies that originated between the snow line and this distance did not contribute to Earth’s accretion because they were swept up by Jupiter and Saturn). In the six models examined, water is added to the Earth mainly after 60–80% of its final mass has accreted. The compositional evolution of the mantles of Venus and Mars are also constrained by the model. The FeO content of the martian mantle depends critically on the heliocentric distance at which the Mars-forming embryo originated. Finally, the Earth’s core is predicted to contain 8–9wt% silicon, 2–4wt% oxygen and 10–60ppm hydrogen, whereas the martian core is predicted to contain low concentrations (<1wt%) of Si and O.

Measurements of lattice thermal conductivity of MgO to core‐mantle boundary pressures

AbstractThe pressure response of lattice thermal conduction in MgO periclase has been a matter of interest for many decades to estimate the thermal conductivity profile of the lower mantle. Using the pulsed light heating thermoreflectance technique, we measured the lattice thermal diffusivity of MgO at pressures up to 137 GPa at 300 K to determine its lattice thermal conductivity under deep lower mantle conditions. Considering the temperature effect estimated by previous high‐temperature measurements, we calculated the lattice part of the thermal conductivity of MgO to be 17.9 ± 1.1 W/m/K at 135 GPa and 3600 K. Additionally, we observed that the lattice conductivity of MgO has little dependence on its grain size under core‐mantle boundary conditions.

Degree one loading by pressure variations at the CMB

Hemispherical asymmetry in core dynamics induces degree-1 pressure variations at the core mantle boundary (CMB), which in turn deforms the overlaying elastic mantle, at the same time keeps center of mass of the whole Earth stationary in space. We develop a systematic procedure to deal with the degree-1 CMB pressure loading. We find by direct calculation a surprisingly negative load Love number h1=−1.425 for vertical displacement. Further analysis indicates that the negative h1 corresponds to thickening above the positive load that defies intuition that pressure inflation pushes overlaying material up and thins the enveloping shell. We also redefine the pressure load Love numbers in general to enable comparison between the surface mass load and the CMB pressure load for the whole spectrum of harmonic degrees. We find that the gravitational perturbations from the two kinds of loads at degrees n>1 are very similar in amplitude but opposite in sign. In particular, if the CMB pressure variation at degree 2 is at the level of ∼1 hpa/yr (1 cm water height per year), it would perturb the variation of Earth’s oblateness, known as the J2, at the observed level.

A Detailed Model Grid for Solid Planets from 0.1 through 100 Earth Masses

ABSTRACTThis article describes a new grid for the mass–radius relation of three-layer exoplanets within the mass range of 0.1–100 M⊕. The three layers are: Fe (!ϵ-phase of iron), MgSiO3 (including both the perovskite phase, post-perovskite phase, and its dissociation at ultrahigh pressures), and H2O (including Ices Ih, III, V, VI, VII, X, and the superionic phase along the melting curve). We discuss the current state of knowledge about the equations of state (EOS) that influence these calculations and the improvements used in the new grid. For the two-layer model, we demonstrate the utility of contours on the mass–radius diagrams. Given the mass and radius input, these contours can be used to quickly determine the important physical properties of a planet including its ! p0 (central pressure), ! p1/p0 (core–mantle boundary pressure over central pressure), CMF (core mass fraction) or CRF (core radius fraction). For the three-layer model, a curve segment on the ternary diagram represents all possible relative mass proportions of the three layers for a given mass–radius input. These ternary diagrams are tabulated with the intent to make comparison to observations easier. How the presence of Fe in the mantle affects the mass–radius relations is also discussed in a separate section. A dynamic and interactive tool to characterize and illustrate the interior structure of exoplanets built upon models in this article is available online.

Crystal structures of (Mg 1- x ,Fe x )SiO 3 postperovskite at high pressures

X-ray diffraction experiments on postperovskite (ppv) with compositions (Mg(0.9)Fe(0.1))SiO(3) and (Mg(0.6)Fe(0.4))SiO(3) at Earth core-mantle boundary pressures reveal different crystal structures. The former adopts the CaIrO(3)-type structure with space group Cmcm, whereas the latter crystallizes in a structure with the Pmcm (Pmma) space group. The latter has a significantly higher density (ρ = 6.119(1) g/cm(3)) than the former (ρ = 5.694(8) g/cm(3)) due to both the larger amount of iron and the smaller ionic radius of Fe(2+) as a result of an electronic spin transition observed by X-ray emission spectroscopy (XES). The smaller ionic radius for low-spin compared to high-spin Fe(2+) also leads to an ordered cation distribution in the M1 and M2 crystallographic sites of the higher density ppv structure. Rietveld structure refinement indicates that approximately 70% of the total Fe(2+) in that phase occupies the M2 site. XES results indicate a loss of 70% of the unpaired electronic spins consistent with a low spin M2 site and high spin M1 site. First-principles calculations of the magnetic ordering confirm that Pmcm with a two-site model is energetically more favorable at high pressure, and predict that the ordered structure is anisotropic in its electrical and elastic properties. These results suggest that interpretations of seismic structure in the deep mantle need to treat a broader range of mineral structures than previously considered.

Melting of Fe–Ni–Si and Fe–Ni–S alloys at megabar pressures: implications for the core–mantle boundary temperature

High pressure melting behavior of three Fe-alloys containing 5 wt% Ni and (1) 10 wt% Si, (2) 15 wt% Si or (3) 12 wt% S was investigated up to megabar pressures by in situ X-ray diffraction and laser-heated diamond anvil cell techniques. We observe a decrease in melting temperature with increasing Si content over the entire investigated pressure range. This trend is used to discuss the melting curve of pure Fe. Moreover, our measurements of eutectic melting in the Fe–Fe3S system show a change in slope around 50 GPa concomitant with the fcc–hcp phase transition in pure solid iron. Extrapolations of our melting curve up to the core–mantle boundary pressure yield values of 3,600–3,750 K for the freezing temperature of plausible outer core compositions.

The melting curve of MgSiO3 perovskite from molecular dynamics simulation

The high-pressure melting curve of MgSiO3 perovskite is simulated by using the constant temperature and pressure molecular dynamics method combined with effective pair potentials. The simulated structural properties of MgSiO3 perovskite at ambient conditions reproduce the experiments and agree well with other theoretical works. The calculated equation of state is very successful in reproducing accurately the recent experimental data over wide pressure ranges. The predicted high-pressure melting curve is in good agreement with the recent experimental and the latest theoretical ones, and the melting curve up to the core–mantle boundary pressure, being very steep at lower pressures, rapidly flattens on increasing pressure. The present results also suggest the validity of the experimental data of Zerr and Boehler (1993 Science 262 553) and Shen and Lazor (1995 J. Geophys. Res. 100 17699).

Solidus and liquidus profiles of chondritic mantle: Implication for melting of the Earth across its history

We investigated the melting properties of a synthetic chondritic primitive mantle up to core–mantle boundary (CMB) pressures, using laser-heated diamond anvil cell. Melting criteria are essentially based on the use of X-rays provided by synchrotron radiation. We report a solidus melting curve lower than previously determined using optical methods. The liquidus curve is found between 300 and 600 K higher than the solidus over the entire lower mantle. At CMB pressures (135 GPa), the chondritic mantle solidus and liquidus reach 4150 (± 150) K and 4725 (± 150) K, respectively. We discuss that the lower mantle is unlikely to melt in the D″-layer, except if the highest estimate of the temperature profile at the base of the mantle, which is associated with a very hot core, is confirmed. Therefore, recent suggestions of partial melting in the lowermost mantle based on seismic observations of ultra-low velocity zones indicate either (1) a outer core exceeding 4150 K at the CMB or (2) the presence of chemical heterogeneities with high concentration of fusible elements. Our observations of a high liquidus temperature as well as a large gap between solidus and liquidus temperatures have important implications for the properties of the magma ocean during accretion. Not only complete melting of the lower mantle would require excessively high temperatures, but also, below liquidus temperatures partial melting should take place over a much larger depth interval than previously thought. In addition, magma adiabats suggest very high surface temperatures in case of a magma ocean that would extend to more than 40 GPa, as suggested by siderophile metal–silicate partitioning data. Such high surface temperature regime, where thermal blanketing is inefficient, points out to a transient character of the magma ocean, with a very fast cooling rate.

The high-pressure melting curve of CaO

Shell-model molecular dynamics simulation has been performed to investigate the melting of the major Earth-forming mineral CaO at elevated temperatures and high pressures, based on thermal instability analysis. The interatomic potential is taken to be the sum of effective pair-wise additive Coulomb, van der Waals attraction, and repulsive interactions. It is shown that the simulated molar volume of CaO is successful in reproducing recent experimental data and our DFT-GGA calculations up to the core–mantle boundary pressure of 135 GPa. The pressure dependence of the simulated high pressure melting temperature of CaO is in good agreement with the results obtained from the Lindemann melting equation at a pressure of below 7 GPa. The extrapolated melting temperatures are in good agreement with the results obtained from Wang’s empirical model up to 60 GPa. The predicted high pressure melting curve, being very steep at lower pressures, rapidly flattens on increasing pressure. The thermodynamic properties of the rocksalt phase of CaO are summarized in the 0–135 GPa pressure range and for temperatures up to 9300 K.

The melting curve of [formula omitted] perovskite under lower mantle pressures

The high pressure melting curve of CaSiO 3 perovskite is simulated by using the constant temperature and pressure molecular dynamics method combined with effective pair potentials for the first time. The simulated results for the partial radial distribution function all compare well with experiment. The calculated equation of state is very successful in accurately reproducing the recent experimental data over a wide pressure range. The predicted high pressure melting curve is in good agreement with the experimental ones, and the melting curve up to the core–mantle boundary pressure, being very steep at lower pressures, rapidly flattens on increasing pressure. The present results also suggest the validity of the experimental data of Zerr and Boehler.

Melting of iron–silicon alloy up to the core–mantle boundary pressure: implications to the thermal structure of the Earth’s core

The melting temperature of Fe–18 wt% Si alloy was determined up to 119 GPa based on a change of laser heating efficiency and the texture of the recovered samples in the laser-heated diamond anvil cell experiments. We have also investigated the subsolidus phase relations of Fe–18 wt% Si alloy by the in-situ X-ray diffraction method and confirmed that the bcc phase is stable at least up to 57 GPa and high temperature. The melting curve of the alloy was fitted by the Simon’s equation, P(GPa)/a = (T m(K)/T 0) c , with parameters, T 0 = 1,473 K, a = 3.5 ± 1.1 GPa, and c = 4.5 ± 0.4. The melting temperature of bcc Fe–18 wt% Si alloy is comparable with that of pure iron in the pressure range of this work. The melting temperature of Fe–18 wt% Si alloy is estimated to be 3,300–3,500 K at 135 GPa, and 4,000–4,200 K at around 330 GPa, which may provide the lower bound of the temperatures at the core–mantle boundary and the inner core–outer core boundary if the light element in the core is silicon.