Fig. 2: Comparison between temporal evolutions of Earth’s mantle temperature and the mantle solidus at different pressures.
From: Melting at the base of a terrestrial magma ocean controlled by oxygen fugacity
a, 16 GPa (470 km depth). b, 21 GPa (580 km depth). c, 26 GPa (710 km depth). Closed symbols are calculated using equation (1), based on core formation at 4.5 Ga proceeding from IW − 4 (ref. 26) or IW − 1 (ref. 12) to IW − 2 (ref. 52), followed by mantle self-oxidation according to Stagno and Fei30. The fO2 used for the open symbols in a–c is the highest fO2 value of natural samples at the corresponding ages53,54. The other fO2 buffers relative to IW (Fe-FeO) are expressed as FMQ = IW + 4 log units (FMQ: Fe2SiO4-Fe3O4-SiO2)56. The dashed lines in panels a–c show mantle temperature evolution curves derived from surface potential temperatures Tp of 1,750, 1,725, 1,700, 1,650 and 1,610 °C at 4, 3.5, 3, 2.5 and 2 Ga (refs. 46,47). Adiabatic temperature gradients of 0.45 °C km−1 (<410 km depth), 0.4 °C km−1 (410–660 km depth) and 0.35 °C km−1 (>660 km depth), and changes of adiabatic temperature by 60 °C at 410 km, by 43 °C at 520 km and by −34 °C at 660 km from Katsura et al.57 were adopted to calculate mantle temperatures. All data are presented as mean values. The error bars of modelled temperatures are ±110 °C from equation (1) (Source Data Fig. 2).
