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dominated by upper mantle and transition zone minerals, which are mostly associated with 37 subducted mafic lithologies rather than peridotite [4][5][6][7]16 . Many superdeep diamonds are made of 38 isotopically light carbon 6,7 and, where measured, their inclusions contain isotopically heavy 39 oxygen 17 , unambiguously indicating an origin from recycled surface material 6,7,17 . The elevated 40 trace element abundances of many silicate inclusions suggest crystallization from a low-degree 41 melt, thought to be generated from melting of subducted oceanic crust 7,18 . Here we examine the fate 42 of subducting carbonated MORB (mid-ocean ridge basalt) as it reaches the transition zone, and the 43 potential for melt-mantle reactions to reproduce superdeep diamonds and their distinctive inclusion 44 assemblages. 45 46Previous experimental studies have investigated the melting behaviour of carbonated basalt at 47 elevated pressures, but only one extends beyond 10 GPa 19. These studies show a remarkable 48 diversity in melting behaviour making extrapolation to higher pressures difficult. In addition, the 49 bulk compositions employed in previous studies often contain considerably more CO 2 than mean 50 oceanic crust, and fall outside the compositional field of natural MORB rocks (see Methods, EDF1 51 and EDT1). To better understand the melting behaviour of deeply subducted oceanic crust we have 52 determined the melting phase relations of a synthetic MORB composition containing 2.5 wt.% CO 2 53 between 3 and 21 GPa (Methods). Our starting composition replicates the major element 54 composition of basaltic rocks from IODP hole 1256D 20 and falls within the range of natural crust 55 compositions 21 (EDF1). 56 57We observe subsolidus phase assemblages containing garnet, clinopyroxene, an SiO 2 polymorph, 58and Ti-rich oxide at all pressures. The carbon component was either CO 2 , dolomite, magnesite or 59 magnesite plus Na-carbonate depending on pressure, and the positions of solid carbonate phase 60 boundaries are consistent with previous studies 22,23 . Near-solidus partial melts are CO 2 bearing 61 silicate melts below 7 GPa, and silica-poor calcic carbonatites above 7 GPa. The alkali component 62 of carbonatite melts increases with pressure (EDF4), and all melts have high TiO 2 /SiO 2 (see 63Methods and extended data items for detailed results). 64 65The melting temperature of carbonated oceanic crust is tightly bracketed from ~ 3 to 21 GPa (figure 66 1). Melting temperatures increase steadily with increasing pressure until about 13 GPa, when the 67 solidus dramatically drops over a narrow pressure interval by ~ 200 °C. This drop in solidus 68 temperature is caused by a change in clinopyroxene composition towards a more Na-rich 69 3 composition above 13 GPa due to dissolution of Na-poor pyroxene components into coexisting 70 garnet. Eventually, clinopyroxene becomes so sodium-rich that a coexisting Na-carbonate mineral 71 Thus, the expulsion of carbonatite melts due to melting of oceanic crust along th...

The Earth took 30-40 million years to accrete from smaller 'planetesimals'. Many of these planetesimals had metallic iron cores and during growth of the Earth this metal re-equilibrated with the Earth's silicate mantle, extracting siderophile ('iron-loving') elements into the Earth's iron-rich core. The current composition of the mantle indicates that much of the re-equilibration took place in a deep (> 400 km) molten silicate layer, or 'magma ocean', and that conditions became more oxidizing with time as the Earth grew. The high-pressure nature of the core-forming process led to the Earth's core being richer in low-atomic-number elements, notably silicon and possibly oxygen, than the cores of the smaller planetesimal building blocks.

[1] We have determined the postspinel transformation boundary in Mg 2 SiO 4 by combining quench technique with in situ pressure measurements, using multiple internal pressure standards including Au, MgO, and Pt. The experimentally determined boundary is in general agreement with previous in situ measurements in which the Au scale of Anderson et al. [1989] was used to calculate pressure: Using this pressure scale, it occurs at significantly lower pressures compared to that corresponding to the 660-km seismic discontinuity. In this study, we also report new experimental data on the transformation boundary determined using MgO as an internal standard. The results show that the transition boundary is located at pressures close to the 660-km discontinuity using the MgO pressure scale of Speziale et al. [2001] and can be represented by a linear equation, P(GPa) = 25.12 À 0.0013T(°C). The Clapeyron slope for the postspinel transition boundary is precisely determined and is significantly less negative than previous estimates. Our results, based on the MgO pressure scale, support the conventional hypothesis that the postspinel transformation is responsible for the observed 660-km seismic discontinuity.

Phase relations of the olivine‐wadsleyite transition in the system (Mg,Fe)2SiO4 have been determined at 1600 and 1900 K using the quench method in a Kawai‐type high‐pressure apparatus. Pressure was determined at a precision better than 0.2 GPa using in situ X‐ray diffraction with MgO as a pressure standard. The transition pressures of the end‐member Mg2SiO4 are estimated to be 14.2 and 15.4 GPa at 1600 and 1900 K, respectively. Partition coefficients for Fe and Mg between olivine and wadsleyite are 0.51 at 1600 K and 0.61 at 1900 K. By comparing the depth of the discontinuity with the transition pressure, the temperature at 410 km depth is estimated to be 1760 ± 45 K for a pyrolitic upper mantle. The mantle potential temperature is estimated to be in the range 1550–1650 K. The temperature at the bottom of the upper mantle is estimated to be 1880 ± 50 K. The thickness of the olivine‐wadsleyite transition in a pyrolitic mantle is determined to be between 7 and 13 km for a pyrolitic mantle, depending on the efficiency of vertical heat transfer. Regions of rapid vertical flow (e.g., convection limbs), in which thermal diffusion is negligible, should have a larger transition interval than stagnant regions, where thermal diffusion is effective. This is in apparent contradiction to short‐period seismic wave observations that indicate a maximum thickness of <5 km. An upper mantle in the region of the 410 km discontinuity with about 40% olivine and an Mg# of at least 89 can possibly explain both the transition thickness and velocity perturbation at the 410 km discontinuity.