Simultaneous determination of melting phase relations of mantle peridotite and mid-ocean ridge basalt at the uppermost lower mantle conditions

Accepted Manuscript Simultaneous determination of melting phase relations of mantle peridotite and mid-ocean ridge basalt at the uppermost lower mantle conditions Hideharu Kuwahara, Ryuichi Nomura, Ryoichi Nakada, Tetsuo Irifune PII: DOI: Reference:

S0031-9201(17)30362-X https://doi.org/10.1016/j.pepi.2018.08.012 PEPI 6188

To appear in:

Physics of the Earth and Planetary Interiors

Received Date: Revised Date: Accepted Date:

31 December 2017 10 June 2018 27 August 2018

Please cite this article as: Kuwahara, H., Nomura, R., Nakada, R., Irifune, T., Simultaneous determination of melting phase relations of mantle peridotite and mid-ocean ridge basalt at the uppermost lower mantle conditions, Physics of the Earth and Planetary Interiors (2018), doi: https://doi.org/10.1016/j.pepi.2018.08.012

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Simultaneous determination of melting phase relations of mantle peridotite and mid-ocean ridge basalt at the uppermost lower mantle conditions Hideharu Kuwaharaa*, Ryuichi Nomuraa*, Ryoichi Nakadab, Tetsuo Irifunea, c a

Geodynamics Research Center, Ehime University, 2-5, Bunkyo-cho, Matsuyama, Ehime 790-8577, Japan

b

Kochi Institute for Core Sample Research, Japan Agency for Marine-Earth Science and Technology, Monobe 200, Nankoku, Kochi 783-8502, Japan

c

Earth-Life Science Institute, Tokyo Institute of Technology, 2-12-1 Ookayama, Meguro, Tokyo 152-8550, Japan

*

Address for editorial correspondence:

Hideharu Kuwahara E-mail: [email protected] Ryuichi Nomura E-mail: [email protected] To be submitted to Physics of the Earth and Planetary Interiors

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Abstract Interpretation of melting phase relationships of mantle peridotite and subducted basaltic crust is important for understanding chemical heterogeneity in the Earth’s interior. Although numerous studies have conducted melting experiments on peridotite and midocean ridge basalt (MORB), and suggested that the solidus temperature of MORB is lower than that of peridotite at whole mantle pressure conditions, both solidus temperatures overlap within their uncertainties. In this study, we conducted simultaneous experiments on KLB-1 peridotite and normal MORB (N-MORB) at pressures from 25 GPa to 27 GPa and temperatures from 2398 K to 2673 K, to compare the solidus temperatures and their melting phase relations. The experimental results show that the solidus temperature of the N-MORB is nearly identical to the KLB-1 peridotite at 25 GPa but lower at 27 GPa. In addition, we found that the crossover of melt fractions between KLB-1 peridotite and N-MORB occurs at 25–27 GPa. These changes are likely to be attributed to the majorite-bridgmanite transition of MORB. This indicates that the dominant melting component may change depending on the location of the uppermost lower mantle. Our calculation result on the density of partial melts along the mantle geotherm suggests that partial melts of KLB-1 peridotite are gravitationally stable around the top of the transition zone, whereas partial melts of NMORB are gravitationally stable even at the top of lower mantle. These results suggest that the distribution of partial melts may be different between KLB-1 peridotite and NMORB in the deep Earth. Our results may be useful for understanding the fate of partial melts of peridotitic mantle and recycled basaltic crust.

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1. Introduction Melting is one of the most important processes that induces chemical heterogeneities in the Earth’s mantle, as represented by the formation of oceanic crust at mid-ocean ridges by partial melting of the shallow mantle. The occurrence of partial melting in the deep mantle has also been implied by seismic observations with seismic low-velocities encountered at the top (Schmandt et al., 2014) and bottom of the lower mantle (e.g., Garnero et al., 1993; Williams and Garnero, 1996). Partial melting of hydrous mantle minerals (Schmandt et al., 2014) at the top of the lower mantle, and of subducted basaltic crust at the core-mantle boundary (e.g., Hirose et al., 1999; Andrault et al., 2014; Pradhan et al., 2015) has been invoked as the origin of the seismic low-velocity regions, though alternative hypotheses have also been proposed for the lower region, such as remnants of a basal magma ocean (e.g., Labrosse et al., 2007; Nomura et al., 2011), the penetration of core materials (Otsuka and Karato, 2012), and the accumulation of iron-rich minerals (Mao et al., 2006). Therefore, melting phase relations of the peridotitic mantle and subducted basaltic crust, especially regarding which components have lower solidus temperature, provide an important clue for understanding chemical heterogeneities in the Earth’s mantle. Numerous melting experiments have been performed on the compositions of a mantle peridotite (e.g., Ito and Takahashi, 1987; Trønnes and Frost, 2002; Fiquet et al., 2010; Nomura et al., 2014; Tateno et al., 2014) and subducted oceanic crust (e.g., Hirose and Fei, 2002; Litasov and Ohtani, 2005; Andrault et al., 2014; Pradhan et al., 2015) at high pressures equivalent to those of the uppermost lower mantle using a multi-anvil apparatus, and of the lowermost lower mantle in a laser-heated diamond anvil cell. However, determining which lithology should melt first at lower mantle conditions remains difficult due to the uncertainties in pressure and temperature evaluations, which make interlaboratory comparisons difficult. In this study, we simultaneously compared melting phase relations of KLB-1 peridotite and N-MORB at the uppermost lower mantle conditions using a multi-anvil apparatus to overcome such problems arisen from the comparisons of the results using different experimental techniques and methods.

2. High-pressure experiments 3

2.1. Starting materials We synthesized two types of starting materials with the chemical compositions of NMORB (Yasuda et al., 1994) and KLB-1 peridotite (Takahashi et al., 1986) from oxide powders (i.e., SiO2, TiO2, Al2O3, FeO, MnO, MgO, CaO, P2O5, Cr2O3, and NiO) and carbonates (Na2CO3 and K2CO3). Oxide powders were dehydrated at 1000 ℃ for 4 h using a muffle furnace before being weighed. The mixtures of dried oxides and carbonates were ground to fine powders under ethanol, and then decarbonated at 1000 ℃ for 10 h using a muffle furnace. The powders were pressed into pellets and reduced in a H2-CO2 gas mixture using a high-temperature furnace. For the KLB-1 peridotite, 2 log-units below the quartz-fayalite-magnetite (QFM) buffer condition was applied at 1000 ℃ for 40 h to reduce the valence state of the iron in the starting material to Fe2+. This is because most of the iron in natural KLB-1 peridotite is present as Fe2+ (e.g., Fiquet et al., 2010). For N-MORB, the Ni-NiO buffer condition was applied at 1400 ℃ for 1 h to buffer the oxidation state near the natural MORB (e.g., Cottrell and Kelley, 2011). The latter sample was quenched into water, and the recovered basaltic glass was again ground to a fine powder. The chemical compositions of starting materials were confirmed by electron probe micro-analysis as summarized in Table 1. It is noted that the synthetic N-MORB glass used in this study is depleted in sodium probably because of the loss of sodium during preparation in a high-temperature furnace. Nevertheless, our results on the melting phase relation of N-MORB can be applied to the interpretation of melting for recycled basaltic crust in the deep Earth because the mineral assemblage at high temperatures near the solidus is consistent with that of previous studies. In the following section 3, we discuss in detail this issue. Moreover, the chemical composition of MORB varies depending on the location (e.g., Klein and Langmuir, 1987). Specifically, the Na2O content of MORB in a region that has thick oceanic crust (i.e. 1.5–2.0 wt. %) is generally lower than the average value of N-MORB (e.g., Klein and Langmuir, 1987). This is because crustal thickness would become increasingly thick with an increase in the degree of melting, whereas the abundance of incompatible elements, such as sodium, in the silicate melt decreased with increasing degree of melting (e.g., Klein and Langmuir, 1987; White and Klein, 2014). Thus, the sodium-depleted basaltic composition used in our study may be relevant for thick

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oceanic crust that has undergone a high degree of melting. Prior to high pressure experiments, synthetic KLB-1 peridotite and N-MORB glass were stored in a vacuum oven at 110 ℃ to avoid water adsorption.

2.2. Multi-anvil press High pressure and temperature experiments were conducted using a 3000-ton Kawaitype multi-anvil apparatus at the Geodynamics Research Center, Ehime University, at 25–27 GPa and 2398–2673 K. The assembly was composed of a Co-doped MgO octahedron with an edge length of 11 mm, and tungsten carbide anvils with a truncation edge length of 5 mm. Experiments were conducted at pressures where mineral assemblages and liquidus phases in basaltic compositions were changing to those stable at lower mantle pressures (Hirose et al., 1999; Hirose and Fei, 2002). Pressures at room temperature were calibrated by the previous study up to 23 GPa (Kuwahara et al., 2017). Pressures at high temperatures of the experimental runs were estimated from the mineral assemblage of the recovered samples and the transformational pressures of majorite to bridgmanite constrained by previous studies on pyrolite and MORB compositions (e.g., Hirose et al., 1999; Hirose and Fei, 2002; Nishiyama and Yagi, 2003; Ishii et al., 2018). For example, the absence of majorite and the presence of bridgmanite in the KLB-1 peridotite sample indicates that it was subject to pressures greater than 24 GPa (e.g., Nishiyama and Yagi, 2003; Ishii et al., 2018), whereas the presence of majorite and the absence of Al-bearing bridgmanite in the N-MORB sample under the identical condition indicates that the pressure condition of this sample was less than 26 GPa (e.g., Hirose et al., 1999; Hirose and Fei, 2002). In this case, the pressure condition of the sample was estimated to be approximately 25 GPa. The uncertainties of the estimated pressures are within ±1 GPa based on previous studies. High-temperature conditions were achieved by a LaCrO3 heater (Run Nos. OT2091, OT2102, OT2121, OT2175, OT2251, OT2256, and OT2267) and a Re heater (Run No. OT2072). Temperature was monitored using a W3%Re-W25%Re thermocouple during heating. No correction was applied for the pressure effect on the electromotive force of the thermocouple. The junction of the thermocouple was positioned at the center of the heater. Each sample was located approximately 0.5 mm from the center. Re was used for capsules of both starting materials, which were located across the thermocouple coaxially to determine 5

the melting phase relations simultaneously under identical pressure and temperature conditions. The power–temperature relationship was used to estimate temperature conditions of Run Nos. OT2091, OT2175, and OT2267 because the thermocouple failed. Specifically, the thermocouple became unstable at 2273–2323 K for Run Nos. OT2175 and OT2267. During these runs, we estimated temperatures by the extrapolation of power–temperature relationships. For Run No. OT2091, the thermocouple failed during compression, and the uncertainty in the temperature should be large, probably on the order of ±100–150 K according to the variation in the power–temperature relationships of other runs (see Appendix A). Temperature gradient is a critical issue in investigating the solidus temperature. To estimate the thermal gradient within the heater, two thermocouples located at the center and outer edge of the heater were used during two experiments (Run Nos. OT2256 and OT2267). One experiment (Run No. OT2267) succeeded in obtaining the thermal gradient within the heater up to 2273 K. The thermal gradient within the LaCrO3 heater used in our cell assembly was estimated to be approximately 50 K/mm at a nominal temperature of 1273 K, and 170 K/mm at a nominal temperature of 2273 K (see Appendix B). This value is consistent with previous studies that used similar scale cell assemblies (e.g., Hirose and Fei, 2002; Trønnes and Frost, 2002). Given this thermal gradient, the temperature at the sample position (i.e. 0.5 mm from the center) might be approximately 85 K lower than the nominal temperature recorded by the thermocouple. However, it should be noted that the thermal gradient along the axial distance is nonlinearly distributed (e.g., Zhang and Herzberg, 1994). More specifically, the temperature gradient at regions near the center is much smaller than that of the colder parts. In addition, a Re capsule may reduce the temperature gradient within the sample because of its high thermal conductivity. Thus, the temperature gap between the sample position and thermocouple may be less than 85 K. The samples were pressurized at room temperature and heated at a rate of 100 K/min to target temperatures. Samples were then quenched by cutting off the power supply at the target temperatures after heating for 5– 20 minutes. The experimental conditions and results are summarized in Table 2.

3. Analytical methods The recovered samples were mounted in epoxy resin, and then polished using SiC and 6

diamond powders. After polishing, the surfaces of the recovered samples were coated with carbon for chemical analysis by electron probe micro-analyzers (EPMA). Quenched textures of the recovered samples were observed by scanning electron microscope (SEM; JEOL JSM 6510LV) and field-emission scanning electron microscope (FE-SEM; JEOL JSM-7000F) at the Geodynamics Research Center, Ehime University. The chemical compositions of the recovered samples were analyzed using EPMA with wavelength-dispersive spectrometers (JEOL JXA 8800 at Ehime University and JEOL JXA 8200 at the Kochi Institute of Core Sample Research, Japan Agency for Marine-Earth Science and Technology). EPMA analyses were conducted using an accelerating voltage of 15 kV and a beam current of 60 nA for most minerals and quenched silicate melts. For bridgmanite, a beam current of 5 nA was applied to avoid structural decomposition. The natural silicate glass KL2-G was used as a standard (Jochum et al., 2006). A defocused beam, 10-20 μm in diameter, was employed for quenched silicate melts to average the heterogeneities of the chemical compositions artificially produced during quenching, though a focused beam was used for minerals. It was noted that the chemical analyses were evenly completed in the sample, not for a specific region. Counting times (peak/background) for major (e.g., Si, Mg) and minor (e.g., Mn, Cr) elements are 20/10 sec and 60/30 sec, respectively.

4. Results 4.1. Melting phase relations of KLB-1 peridotite and N-MORB at 25–27 GPa The results of the chemical analyses and back scattered electron images of the recovered samples are shown in Tables 3–4 and Figs. 1–8, respectively. Regarding the KLB-1 peridotite composition, a mineral assemblage of bridgmanite, ferropericlase, and calcium perovskite was observed under sub-solidus conditions at 25–27 GPa and 2398– 2523 K. At 25 GPa and temperatures from 2473 K to 2498 K, a large fraction of silicate melt surrounded by ferropericlase was observed at the hottest part of the sample, where ferropericlase is interpreted as the liquidus phase. The coexistence of bridgmanite and ferropericlase was observed in the colder region, suggesting that calcium perovskite is the first melting phase at 25 GPa. At 25 GPa and 2573 K, most parts of the sample were molten and a trace amount of ferropericlase was found in the coldest part as a liquidus phase. At 27 GPa, the solidus temperature of the KLB-1 peridotite increases above 2523 7

K. At 27 GPa and 2673 K, the sample was nearly molten. The crystallization sequence of ferropericlase, bridgmanite, and calcium perovskite observed in this study is consistent with that reported in Hirose and Fei (2002) at 25–27 GPa. There is a variation in the iron content of the sample among the experimental runs for the KLB-1 peridotite. This might be caused by the reaction of metallic iron with the surrounding Re capsules via the disproportionation reaction of ferric iron (e.g., Frost et al., 2004; Trønnes and Frost, 2002). Indeed, metallic iron is indicated by the bright spot in some recovered samples (see Fig. 6d), and a trace amount of metallic iron (i.e. < 0.16 wt. %) was detected in the inner wall of a Re capsule. This indicates that samples were probably more oxidized than the starting materials and buffered by the Re-ReO2 buffer (e.g., Frost and Langenhorst, 2002). Regarding N-MORB composition, the sub-solidus mineral assemblage was majorite, stishovite, calcium perovskite, and Ti-rich Mg-perovskite at 25 GPa and 2398 K. The result is consistent with a previous study (e.g., Litasov and Ohtani, 2005), although a new aluminous silicate (NAL) phase (e.g., Akagoi et al., 1999; Miyajima et al., 2001) was not observed in this study, presumably beacause of the sodium-depleted starting compositions. At 25 GPa and temperatures from 2473 K to 2498 K, a small amount of Ti- and Fe-concentrated quenched melt was observed. Calcium perovskite existed along with the solid–liquid boundary. Stishovite appeared in the colder region of the sample, followed by majorite. At 2473 K, a Ca-rich aluminous silicate (CAS) phase (Irifune et al., 1994) was also observed. At 25 GPa and 2573 K, a large fraction of the silicate melt and a layer of majorite at the solid–liquid boundary were observed. Majorite, calcium perovskite, and stishovite appeared along the thermal gradients. At 27 GPa and 2448 K, we observed a segregation of calcium perovskite on the hotter side of the capsule (see Fig. 5), whereas the appearance of coexisting CAS phase and stishovite, and of bridgmanite in the colder region was noted. At 27 GPa and temperatures from 2473 K to 2523 K, a large fraction of melt was observed with the crystallization of calcium perovskite, CAS phase, and stishovite. Such a rapid increase in the degree of melting was not observed at 25 GPa. Because the majorite-bridgmanite transition occurs at approximately 26–27 GPa for MORB compositions (e.g., Hirose et al., 1999; Hirose and Fei, 2002), the change in the degree of melting between 25 GPa and 27 GPa may be related to the majorite-bridgmanite transition of MORB. At 27 GPa and 2673 K, most 8

parts of the sample were molten and a trace amount of calcium perovskite was found in the coldest part as a liquidus phase. The present study shows that the liquidus phase changes from majorite to calcium perovskite near the experimental pressures, consistent with Hirose and Fei (2002), who reported that the liquidus phase changed from garnet majorite at 22 GPa to calcium perovskite at 26 GPa. Although the NAL phase was not observed during our experiments, the results can be applied to the interpretation of melting of recycled basaltic crust under the uppermost lower mantle conditions because the NAL phase is not stable at high temperatures around the solidus of MORB under these conditions (e.g., Hirose and Fei, 2002).

4.2. Solidus temperature and degree of melting Fig. 9 shows a comparison of the solidus temperatures between our experiments and previous studies for KLB-1 peridotite and N-MORB compositions. The dataset was compiled to determine and compare the solidus temperatures of KLB-1 peridotite and N-MORB compositions. It is noted that experimental data for KLB-1 peridotite and NMORB compositions without any additions of trace elements, or metallic iron, are shown in Fig. 9. As expected, interlaboratory compilation yields a large range of solidus temperatures. Specifically, there are inconsistencies between sub-solidus and partial melting temperatures in previous studies at pressures between 22–25 GPa for KLB-1 peridotite composition. For higher pressure conditions corresponding to lower mantle conditions, it is difficult to define an appropriate region of possible solidus temperatures due to the sparse data both on KLB-1 peridotite and N-MORB compositions. Thus, there are virtually no constraints practically regarding which lithology should melt first at these pressures (Fig. 9c). In contrast, it was suggested that the intralaboratory comparisons of experimental data based on identical cell assembly were much more consistent for understanding temperature-melting chemistry relations (Zhang and Herzberg, 1994). However, limited datasets make intralaboratory comparisons of the solidus temperatures between KLB-1 peridotite and N-MORB compositions difficult. Specifically, Hirose and Fei (2002) showed that N-MORB melted at a minimum temperature of 2423 K, yet KLB-1 peridotite did not melt until at least at 2473 K, at 22 GPa. However, comparisons are practically impossible with a temperature resolution of <300 K above 22 GPa due to the 9

sparseness of the dataset. In contrast, our simultaneous melting experiments clearly demonstrated that the solidus temperature of N-MORB is slightly lower than that of KLB-1 peridotite at around 27 GPa. Fig. 10 shows the degree of melting as a function of temperature. The phase proportions were estimated from mass balance calculations as described in the following equation: (1) where,

and

are the mass concentrations of element i in starting materials and

phase j in the sample, respectively.

is the mass fraction (in wt. %) of phase j. The

mass fraction of each phase was estimated via multiple regression using the results of EPMA analyses. The detailed result of the mass balance calculations is summarized in Table 5. In our experiments at 25 GPa, the degree of melting for KLB-1 peridotite rapidly increased up to 80 wt. % within 100 K above the solidus, whereas the degree of melting for N-MORB gradually increased up to approximately 30 wt. % within 100 K above the solidus. At 2573 K, KLB-1 peridotite was nearly fully molten, whereas the degree of melting of N-MORB was approximately 50 wt. %. In contrast to the results at 25 GPa, the degree of melting for N-MORB rapidly increases with a temperature above the solidus at 27 GPa, whereas KLB-1 peridotite was not molten at identical temperature conditions. This indicates that the crossover of the degree of melting between KLB-1 peridotite and N-MORB occurs at approximately 25–27 GPa. As previously mentioned in section 4.1, this change in the degree of melting may be related to the majorite-bridgmanite transition for N-MORB. At 2673 K, a nearly fully molten state was achieved for both KLB-1 peridotite and N-MORB. The narrow solidus– liquidus temperature interval for peridotitic compositions is consistent with previous observations (e.g., Zhang and Herzberg, 1994; Trønnes and Frost, 2002).

5. Discussion 5.1. Solidus temperatures of KLB-1 peridotite and subducted MORB in the deep Earth’s mantle Our results of simultaneous melting experiments clarified that N-MORB has a lower solidus temperature than KLB-1 peridotite at 27 GPa, under conditions where the 10

peridotite has a lower mantle mineral assemblage, and the basalt transforms to a lower mantle mineral assemblage. N-MORB has a lower solidus temperature even at 22 GPa, where the peridotite has the mineralogy of the transition zone based on intralaboratory comparison (Hirose and Fei, 2002). However, the solidus temperatures of KLB-1 peridotite and N-MORB are nearly identical at 25 GPa. Therefore, it is concluded that N-MORB has solidus temperatures lower than those of KLB-1 peridotite at least at pressures corresponding to the lower part of the transition zone and the top of the lower mantle, i.e. greater than 27 GPa, although the situation is complicated at lower pressures beacause of the difficulty and apparent contradictions in the interlaboratory comparisons (Takahashi, 1986; Yasuda et al., 1994; Zhang and Herzberg, 1994; Herzberg and Zhang, 1996). The solidus temperature of the KLB-1 peridotite inferred from our experiments at 25–27 GPa is consistent with the solidus curve at 22–24.5 GPa (Trønnes and Frost, 2002) within ±50 K. In contrast, there is a discrepancy in the solidus temperature of NMORB at 25–27 GPa between our study and that of Hirose and Fei (2002). More specifically, the solidus temperature of N-MORB at 26 GPa (i.e. ~2623 K) inferred from the experiments of Hirose and Fei (2002) is approximately 150 K higher than ours (i.e. ~2473 K). The thermal gradient within the heater is nearly identical between our study (i.e. 170 K/mm at 2273 K) and that of Hirose and Fei (2002) (i.e. 150 K/mm at 2273 K). In addition, the mineral assemblage near the solidus is also the same. Thus, the discrepancy in the solidus temperature of N-MORB might be because of the difference in the location of the thermocouple junction from the sample. It is also noted that our experimental results have relatively large uncertainties in terms of an accurate solidus temperature because of the lack of temperature measurements using a thermocouple during some experimental runs, although the results for the KLB-1 peridotite are relatively consistent with the previous study (Trønnes and Frost, 2002). In multi-anvil experiments, achievable pressure conditions are far below those from the deeper mantle up to the core-mantle boundary, but the results may also be useful for constraining the solidus temperatures for KLB-1 peridotite and N-MORB at the core-mantle boundary. This is because the temperatures determined by laser-heated diamond-anvil cell experiments (Fiquet et al., 2010; Andrault et al., 2014; Nomura et al., 2014; Pradhan et al., 2015) are highly contradictive due to different melting criteria, the 11

inaccuracy of temperature measurements by unknown wavelength-dependent emissivity, and integrated thermal radiation measured along temperature gradients. It should be noted that the solidus temperature of N-MORB estimated from our multi-anvil experiments is consistent with the melting curve of N-MORB determined by Hirose et al. (1999) and Pradhan et al. (2015) within ±100–200 K, but not with that of Andrault et al. (2014). The narrow solidus–liquidus temperature interval obtained in our study is not consistent with the wide solidus–liquids temperature interval reported by Andrault et al. (2014). Such a difference may have been caused by the difference in initial volatile contents in the starting materials because Andrault et al. (2014) used natural MORB glass which contains a moderate amount of H2O (i.e. 2730 ppm) and CO2 (i.e. 165 ppm). The pressure dependence of the solidus temperatures clarified by combining the results of this study and future experiments at higher pressures where both peridotitic and basaltic compositions have a lower mantle mineral assemblage will provide further constraints on the melting relations at deep Earth conditions up to the core-mantle boundary.

5.2. Implications for partial melts at the top of lower mantle Abrupt reductions in seismic shear-velocity of around 2.6 %, which are interpreted as being caused by partial melting of the mantle, have been observed at depths of ~730 km beneath North America (Schmandt et al., 2014). In order to test the hypothesis using seismological and/or geochemical observations, it is important to know whether basaltic components scattered in the transition zone (Lee and Chen, 2007), or surrounding mantle peridotite induce the melting. Although basalt melts at a lower temperature than peridotite (Fig. 10c), the temperature it requires is far higher than the typical mantle geotherm. Therefore, a large amount of water may be necessary to induce partial melting at the top of lower mantle as proposed by previous studies (Schmandt et al., 2014; Liu et al., 2018). Although melting phase relations of hydrous silicate systems have been experimentally investigated (Litasov and Ohtani, 2002; Litasov and Ohtani, 2005), the effect of water on the solidus temperature should be reinvestigated because previous studies used simplified (CaO-MgO-Al2O3-SiO2) pyrolitic compositions (Litasov and Ohtani, 2002), and did not measure the water content in the recovered samples (Litasov and Ohtani, 2002; Litasov and Ohtani, 2005). If dehydration is the 12

major cause of partial melting at the top of lower mantle, it is difficult to constrain the main component of partial melts at the top of lower mantle from the comparison of solidus temperatures between KLB-1 peridotite and N-MORB because the presence/absence of partial melts is mainly controlled by the amount of water. The gravitational stability of partial melts is also an important issue for understanding the distribution of partial melts in the deep Earth’s mantle. Previous studies have investigated the gravitational stability of silicate melts of pyrolite and basalt up to transition zone pressures (e.g., Matsukage et al., 2005; Sakamaki et al., 2006; Lee et al., 2010). However, there are a few studies on the stability of silicate melts at the uppermost lower mantle conditions (e.g., Sakamaki, 2017). Here we estimated the density of partial melts obtained in this study using a thermodynamic model applicable up to 40 GPa (Ghiorso, 2004a-d) (see Appendix C). Fig. 911a and 11b show the densities of partial melts generated from KLB-1 peridotite and N-MORB along the current mantle geotherm (Katsura et al., 2010). It is noted that temperature and iron content are the dominant factors that determine density. Because partial melts of NMORB become less iron and titanium with increasing temperature, the density of a partial melt of basalt decreases with the degree of melting. Therefore, on rising during cooling, the melt becomes enriched with iron and titanium and is eventually gravitationally neutral at the bottom of the transition zone. In contrast to N-MORB, partial melts of KLB-1 peridotite in our study may be gravitationally neutral around the upper part of the transition zone but not stable the bottom of the transition zone. This result is consistent with the previous study (Lee et al., 2010). We note that the degree of melting for silicate melts of the KLB-1 peridotite in our study is much higher than realistic conditions. Given that silicate melts with a lower degree of melting should become richer in iron, the estimated density of the partial melts of the KLB-1 peridotite may provide the lower bound. The difference in the density of partial melts between KLB-1 peridotite and N-MORB may be caused by the large difference in the degree of melting at a given temperature. This is because the concentrations of iron and titanium in partial melts are controlled by the degree of melting. It is also noted that N-MORB contains higher amount of heavy oxides, such as FeO and TiO2, than KLB-1 peridotite. If this is the case, the location of the remnants of partial melting will differ for peridotite and basalt. Given that the effect of water on the density of silicate melts may become 13

negligible at the uppermost lower mantle conditions (e.g., Sakamaki, 2017), our results suggest that the dominant component of partial melts may be different between the top of the transition zone and the top of uppermost lower mantle. Given that the Archean mantle was hotter than that of today, the melting of dry peridotitic mantle may have occurred at a depth corresponding to the uppermost lower mantle conditions in the early Earth (e.g., Herzberg, 1995; Lee et al., 2010 and references therein). More specifically, adiabat intersects the solidus of KLB-1 peridotite at the uppermost lower mantle conditions when the mantle potential temperature is ~2273 K (e.g., Lee et al., 2010). If this is the case, the fate of the peridotitic melts under deep mantle conditions is important in understanding the chemical differentiation of the early Earth’s mantle. Figure 11c shows the density of the peridotitic melts along the adiabat with a slope of 10 K/GPa when the mantle potential temperature is 2273 K. The result suggests that the partial melts of KLB-1 peridotite are gravitationally stable at the top of the transition zone. A previous study has indicated that melts formed at a depth of approximately 14–18 GPa (Lee et al., 2010). Our results suggest that melts formed under the uppermost lower mantle conditions could also have been gravitationally stable. As already proposed by Lee et al. (2010), such dense melts may be important in understanding the distribution of geochemically important incompatible elements, such as U and Th, in the Earth.

6. Summary In this study, we experimentally investigated melting phase relations of KLB-1 peridotite and N-MORB, simultaneously. Our results show that both KLB-1 peridotite and N-MORB have a similar solidus-liquidus region at uppermost lower mantle conditions at 25 GPa, but the solidus temperature of N-MORB is lower than that of KLB-1 peridotite at 27 GPa. The results also show that the crossover of the degree of melting occurs between KLB-1 peridotite and N-MORB at 25–27 GPa. These changes may be related to the majorite-bridgmanite transition for N-MORB composition. In addition, calculation results on the density of partial melts obtained in our experiments suggest that partial melts of N-MORB may be gravitationally stable at the top of lower mantle, whereas partial melts of KLB-1 peridotite may be stable at the top of the transition zone. This is because the degree of melting is largely different between 14

KLB-1 peridotite and N-MORB at a given temperature, and therefore, the concentrations of iron and titanium in partial melts which are the dominant controlling factors for determining density are also different. Our results suggest that the distribution of partial melts in the deep Earth’s mantle may differ between KLB-1 peridotite and N-MORB. The results may be useful for understanding the fate of partial melts in the deep Earth.

Acknowledgments We would like to thank Zhou Youmo for technical assistance with high-pressure experiments. This work was supported by JSPS KAKENHI Grant Numbers JP15H05470, JP15H05830, JP15H05829, 18J00966, and 18K13635. The authors would like to thank three anonymous reviewers and Dr. Yingwei Fei for their helpful comments which greatly improved this study.

15

Appendix A. Uncertainties in the estimation of temperature for Run Nos. OT2091, OT2175, and OT2267 For Run No. OT2091, temperature was estimated from the power–temperature relationships obtained via high-pressure experiments using an identical cell assembly at 25 GPa. As shown in Figure A1, the power–temperature relationship for each run varied at high temperatures of approximately 2273 K within ±100–150 K. For Run Nos. OT2175 and OT2267, temperatures below 2273 K were monitored using the thermocouple (see Figure B2). Temperatures above 2273 K were estimated via extrapolation of the power–temperature relationships obtained from each run. The power–temperature relationship was obtained via a polynomial fitting with a non-fixed intercept.

16

Figure A1. Power–temperature relationships at 25 GPa.

Figure A2. Power–temperature relationships for Run Nos. OT2175 and OT2267. Open squares indicate measured temperatures at a given power input. Dashed curves are the extrapolation of the power–temperature relationships obtained from each run.

17

Appendix B. Thermal gradient within the heater of Run No. OT2267 Figure B1 shows the temperature gradient within the LaCrO3 heater obtained during Run No. OT2267 using two thermocouples.

Figure B1. Temperature gradient within the heater obtained during Run No. OT2267. The X-axis shows the nominal temperature recorded by the thermocouple at the center. The Y-axis shows the temperature gradient between the center and the edge of the heater at a given nominal temperature.

18

Appendix C. The calculation of silicate liquid density The density of partial melts for both the KLB-1 peridotite and N-MORB was estimated using the equation of state (EOS) for silicate liquid proposed by Ghiorso (2004a-d). In this method, the volume of multicomponent silicate liquids at the temperature of interest is first calculated under 1 bar. Then, the volume of the silicate liquids at the pressure of interest is calculated. Based on Ghiorso’s EOS for silicate liquid, the pressure dependence on the volume of multicomponent silicate liquids is described by the following equation: (C.1) where

is the volume of the multicomponent silicate liquids at a given pressure ;

the reference pressure (i.e. 1 bar);

,

pressure-derivatives of the volume; and

, and and

is

are the model parameters describing the are also parameters as follows: (C.2)

(C.3) where

, and

are model parameters describing the pressure-derivatives of the

volume. At 1 bar, the equation (C.1) can be reduced to

.

is provided by a

combination of the linear mixing relation of the volumes for a multicomponent system and its temperature dependence is described in the following equation: (C.4) where

is the volume of the multicomponent silicate liquids at the reference

temperature (i.e. 1673 K) and

is the thermal expansion. Here,

and

are provided

via the following equations: (C.5)

(C.6) where

and

are the mole fraction and the partial molar volume of element i at the

19

reference temperature, respectively.

is the temperature-derivative form of the

volume at the reference temperature as follows: (C.7) where

is the temperature derivative of the volume of each component i. If the

silicate liquids contain a large amount of TiO2, the partial molar volume and its temperature derivative form of TiO2 are described using the following equations: (C.8) (C.9) Next, we estimate the parameter

.

is expressed via the following equation: (C.10)

where

is the weight in g,

is the sound speed, and

is the heat capacity of the

multicomponent silicate liquid. These parameters are described via the following equations: (C.11) (C.12)

(C.13)

In equation (C.11),

is the heat capacity of element i. In equation (C.12),

molecular weight of element i. In equation (C.13), the sound speed of element i, respectively.

,

parameters. Other volume parameters ( ,

, and

and

is the

are the mole fraction and

, and

are the interaction

) are provided by a linear function

of volumetric parameters for each component i as follows: (C.14)

20

(C.15) (C.16) Parameters used in this study are listed in Table C.1. In addition, we compared the estimated densities to experimental results obtained from high-pressure experiments using the sink-flotation method of olivine and diamond (Agee and Walker, 1993; Suzuki et al., 1995; Smith and Agee, 1997; Agee, 1998; Ohtani et al., 1998; Ohtani and Maeda, 2001) to assess the uncertainty of this method. The result is shown in Figure C.1. For a peridotitic composition, Ghiorso’s EOS slightly overestimates the melt density. In contrast, Ghiorso’s EOS underestimates the melt density of basaltic compositions. Thus, the density of partial melts for the KLB-1 peridotite obtained from this study may be slightly lower than the actual density, whereas the partial melts of N-MORB may be denser than our calculated results.

21

Figure C.1. Comparison of melt densities between Ghiorso’s EOS for silicate liquid and experimental results. Melt densities obtained from calculation and experiments are shown as an open diamond. The solid and dashed lines indicate the 1:1 correspondence and 5% error bounds, respectively. The experimental data on the peridotitic melt were taken from Agee and Walker (1993), Suzuki et al. (1995), and Ohtani et al. (1998). The experimental data on the MORB melt were taken from Agee (1998) and Ohtani and Maeda (2001). The experimental data on the picritic melt were taken from Smith and Agee (1997) and Ohtani and Maeda (2001). The density of each silicate melt was calculated under identical conditions (i.e. pressure, temperature, and composition).

22

Table C.1. Parameters for density calculation used in this study (cm3/mol)

(cm3/mol)

(cm3/mol)

(cm3/mol)

(cm3/K/mol) (m/sec)

SiO2 TiO2

26.71 23.45

0.21995 0.059857

0.01022 0.1035

-0.00025985 -0.054133

(g/mol)

(m/sec/K) (J/K/mol)

1.00687×10-3

2321.75

0.39934

82.6

60.0848

-3

1693.60

0.81199

109.2

79.8988

-3

2738.35

0.50394

170.3

101.9612

6.80672×10

Al2O3

37.62

0.15738

-0.030193

0.0091947

-6.48602×10

FeO

13.90

0.22771

-0.12968

0.060629

1.53203×10-3

2399.53

-0.10726

78.8

71.8464

-0.037501

2.88655×10

-3

3349.96

0.27564

94.2

40.3114

3.14295×10

-3

3967.42

-0.20526

89.8

56.0794

6.07700×10

-3

3080.69

-2.16757

97.6

61.9790

-3

-1325.21

-

-

-

5800.72

-

-

-

MgO CaO Na2O Na2O-

12.02 16.67 29.12

-0.025979 0.28439 3.4298

0.045354 -0.18258 -0.58834

0.043285 -2.351

-

-

-

-

9.69858×10

-

-

-

-

-

TiO2 Na2OAl2O3

23

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Ghiorso, M., 2004d. An equation of state for silicate melts. Ⅳ. Calibration of a multicomponent mixing model to 40 GPa. Am. J. Sci. 304, 811-838. Herzberg, C., 1995. Generation of plume magmas through time: an experimental perspective. Chem. Geol. 126, 1-16. Herzberg, C., Zhang, J., 1996. Melting experiments on anhydrous peridotite KLB-1: Compositions of magmas in the upper mantle and transition zone. J. Geophys. Res. 101, 8271-8295. Hirose, K., Fei, Y., Ma, Y., Mao, H.-K., 1999. The fate of subducted basaltic crust in the Earth’s lower mantle. Nature 397, 53-56. Hirose, K., Fei, Y., 2002. Solidus and melting phase relations of basaltic composition in the uppermost lower mantle. Geochim. Cosmochim. Acta, 66, 2099-2108. Irifune, T., Ringwood, A. E., Hibberson, W. O., 1994. Subduction of continental crust and terrigenous and pelagic sediments: An experimental study. Earth Planet. Sci. Lett. 126, 351-368. Ito, E., Takahashi, E., 1987. Melting of peridotite at uppermost lower-mantle conditions. Nature 328, 514-517. Ito, E., Kubo, A., Katsura, T., Walter, M. J., 2004. Melting experiments of mantle minerals under lower mantle conditions with implications for magma ocean differentiation. Phys. Earth Planet. Inter. 143-144, 397-406. Katsura, T., Yoneda, A., Yamazaki, D., Yoshino, T., Ito, E., 2010. Adiabatic temperature profile in the mantle. Phys. Earth Planet. Inter. 183, 212-218. Kuwahara, H., Gotou, H., Shinmei, T., Ogawa, N., Yamaguchi, A., Takahata, N., Sano, Y., Yagi, T., Sugita, S., 2017. High pressure experiments on metal-silicate partitioning of chlorine in a magma ocean: Implications for terrestrial missing chlorine.

Geochemistry,

Geophysics,

Geosystems

18,

https://doi.org/10.1002/2017GC007159. Klein, E. M., Langmuir, C. H., 1987. Ocean ridge basalt chemistry, axial depth, crustal thickness and temperature variations in the mantle. J. Geophys. Res. 92, 8089-8115. Labrosse, S., Hernlund, J. W., Coltice, N., 2007. A crystallizing dense magma ocean at the base of the Earth’s mantle. Nature 450, 866-869. Lee, C.-T. A., Chen, Q.-P., 2007. Possible density segregation of subducted oceanic lithosphere along a weak serpentine layer and implications for compositional 25

stratification of the Earth’s mantle. Earth Planet. Sci. Lett. 255, 357-366. Lee, C.-T. A., Luffi, P., Hoink, T., Li, J., Dasgupta, R., Hernlund, J., 2010. Upsidedown differentiation and generation of a ‘primordial’ lower mantle Nature 463, 930933. Litasov, K., Ohtani, E., 2002. Phase relations and melt compositions in CMAS-pyroliteH2O system up to 25 GPa. Phys. Earth Planet. Inter. 134, 105-127. Litasov, K. D., Ohtani, E., 2005. Phase relations in hydrous MORB at 18-28 GPa: Implications for heterogeneity of the lower mantle. Phys. Earth Planet. Inter. 150, 239-263. Liu, Z., Park, J., Karato, S., 2018. Seismic evidence for water transport out of the mantle transition zone beneath the European Alps. Earth Planet. Sci. Lett. 482, 93104. Mao, W. L., Mao, H.-K., Sturhahn, W., Zhao, J., Prakapenka, V. B., Meng, Y., Shu, J., Fei, Y., Hemley, R. J., 2006. Iron-rich post-perovskite and the origin of ultralowvelocity zones. Science 312, 564-565. Matsukage, K. N., Jing, Z., Karato, S.-I., 2005. Density of hydrous silicate melt at the conditions of Earth’s dep upper mantle. Nature 438, 488-491. Miyajima, N., Yagi, T., Hirose, K., Kondo, T., Fujino, K., Miura, H., 2001. Potential host phase of aluminum and potassium in the Earth’s lower mantle. Am. Mineral. 86, 740-746. Nomura, R., Ozawa, H., Tateno, S., Hirose, H., Hernlund, J., Muto, S., Ishii, H., Hiraoka, N., 2011. Spin crossover and iron-rich silicate melt in the Earth’s deep mantle. Nature 473, 199-202. Nomura, R., Hirose, K., Uesugi, K., Ohishi, Y., Tsuchiya, A., Miyake, A., Ueno, Y., 2014. Low core-mantle boundary temperature inferred from the solidus of pyrolite. Science 343, 522-525. Nomura, R., Youmo, Z., Irifune, T., 2017. Melting phase relations in the MgSiO3– CaSiO3 system at 24 GPa. Progress in Earth and Planetary Science 4:34, DOI 10.1186/s40645-017-0149-2. Ohtani, E., Suzuki, A., Kato, T., 1998. Flotation of olivine and diamond in mantle melt at high pressure: Implications for fractionation in deep mantle and ultradeep origin of diamond. Geophys. Monogr. 101, 227-239. 26

Otsuka, K., Karato, S., 2012. Deep penetration of molten iron into the mantle caused by a morphological instability. Nature 492, 243-246. Pradhan, G. K., Fiquet, G., Siebert, J., Auzende, A.-L., Morard, G., Antonangeli, D., Garbarino, G., 2015. Melting of MORB at core-mantle boundary. Earth Planet. Sci. Lett. 431, 247-255. Sakamaki, T., Suzuki, A., Ohtani, E., 2006. Stability of hydrous melt at the base of the Earth’s upper mantle. Nature 439, 192-194. Sakamaki, T., 2017. Density of hydrous magma. Chem. Geol. 475, 135-139. Schmandt, B., Jacobsen, S. D., Becker, T. W., Liu, Z., Dueker, K. G., 2014. Dehydration melting at the top of the lower mantle. Science 344, 1265-1268. Smith, J. R., Agee, C. B., 1997. Compressibility of molten ‘green glass’ and crystalliquid density crossover in low-Ti lunar magma. Geochim. Cosmochim. Acta 61, 2139-2145. Suzuki, A., Ohtani, E., Kato, T., 1995. Flotation of diamond in mantle melt at high pressure. Science 269, 216-218. Takahashi, E., 1986. Melting of a dry peridotite KLB-1 up to 14 GPa: Implications on the origin of peridotitic upper mantle. J. Geophys. Res. 91, 9367-9382. Tateno, S., Hirose, K., Ohishi, Y., 2014. Melting experiments on peridotite to lowermost mantle conditions. J. Geophys. Res. 119, 4684-4694. Trønnes, R. G., Frost, D. J., 2002. Peridotite melting and mineral-melt partitioning of major and minor elements at 22-24.5 GPa. Earth Planet. Sci. Lett. 197, 117-131. White, W. M., Klein, E. M., 2014. Composition of the oceanic crust. In: Holland, H. D., Turekian, K. K. (Eds.), Treatise on Geochemistry 4, 457-496. Williams, Q., Garnero, E. J., 1996. Seismic evidence for partial melt at the base of Earth’s mantle. Science 273, 1528-1530. Yasuda, A., Fujii, T., Kurita, K., 1994. Melting phase relations of an anhydrous midocean ridge basalt from 3 to 20 GPa: Implications for the behavior of subducted oceanic crust in the mantle. J. Geophys. Res. 99, 9401-9414. Zhang, J., Herzberg, C., 1994. Melting experiments on anhydrous peridotite KLB-1 from 5.0 to 22.5 GPa. J. Geophys. Res. 99, 17729-17742.

27

Table 1. Chemical compositions (wt.%) of synthesized anhydrous starting materials and comparisons with previous studies KLB-1 peridotite Takahashi (1986)

N-MORB

Herzberg and

This study

Yasuda et al.

Zhang (1996)

Hirose et al. (1999)

(1994) a

Litasov and Ohtani

This study

(2005)

SiO2

44.48

44.30

45.28(1.09 )

49.71

49.64

51.09

49.56(0.11a)

TiO2

0.16

0.12

0.18(0.01)

1.71

1.64

1.50

1.68 (0.00)

Al2O3

3.59

3.54

3.92(0.16)

15.68

14.88

15.70

15.88(0.00)

FeO

8.1

8.59

7.88(0.57)

9.36

11.43

9.81

10.29(0.00)

MnO

0.12

0.14

0.09(0.00)

0.18

0.18

-

0.15(0.03)

MgO

39.22

39.50

37.84(0.83)

8.43

8.51

7.74

8.89(0.00)

CaO

3.44

3.03

3.14(0.78)

11.73

10.55

11.51

11.34(0.00)

Na2O

0.30

0.30

0.38(0.08)

2.76

2.90

2.48

0.83(0.01)

b

0.23

0.12

0.17

n.d.

K2O

0.02

-

P2O5

-

-

-

0.02

-

-

n.d.

Cr2O3

0.31

0.38

0.33(0.02)

-

-

-

-

NiO

0.25

0.21

n.d.

-

-

-

-

Total

99.72

100.11

99.03

99.79

99.85

100.00

98.61

a

One standard deviation

b

Not detected

n.d.

28

Table 2. Summary of experimental conditions and observed phase assemblage Run No.

P (GPa)

Nominal T (K)a

Duration (min)

OT2072

25

2398

15

OT2256

25

2473

10

OT2102

25

2498

15

OT2091c

25

2573

10

OT2175c

27

2448

20

OT2251

27

2473

15

OT2267 c

27

2523

10

OT2121

27

2673

5

Phase assemblageb

Starting material KLB-1 peridotite N-MORB KLB-1 peridotite N-MORB KLB-1 peridotite N-MORB KLB-1 peridotite N-MORB KLB-1 peridotite N-MORB KLB-1 peridotite N-MORB KLB-1 peridotite N-MORB KLB-1 peridotite N-MORB

Sub-solidus: Bdg + Fp + CaPv Sub-solidus: Maj + St + CaPv, Ti-MgPv Melt, Fp, Bdg Melt, CaPv, CAS + St, Maj Melt, Fp, Bdg Melt, CaPv, St, Maj Melt, Fp Melt, Maj, CaPv. St Sub-solidus: Bdg + Fp + CaPv Melt, CaPv, CAS + St, Bdg Sub-solidus: Bdg + Fp + CaPv Melt, CaPv, CAS + St Sub-solidus: Bdg + Fp + CaPv Melt, CaPv, CAS + St Melt, Fp Melt, CaPv

a

Nominal temperature was determined by thermocouple.

b

Bdg, bridgmanite; Fp, ferropericlase; Maj, majorite; St, stishovite; CaPv, calcium perovskite; CAS, Ca-Al phase; Ti-MgPv, Ti-rich magnesium perovskite. Phases are shown in

order of their appearance from hot to cold part of the sample. Coexisting phases are linked by “+.” c

Temperature was estimated by the power–temperature relationship

29

Table 3. Chemical compositions of recovered samples for KLB-1 peridotite Sample

P

Nominal

No.

(GPa)

T (K)

25

2398

SiO2

TiO2

Al2O3

FeO

MnO

MgO

CaO

Na2O

Cr2O3

Total

Na

51.91(1.35

0.20(0.03)

4.03(0.41)

5.49(0.67)

0.11(0.04)

37.01(0.79)

0.77(032)

0.14(0.02)

0.30(0.04)

99.97(1.02)

9

0.14(0.04)

n.d.c

1.11(0.08)

17.29(1.25)

0.18(0.00)

78.03(1.73)

0.06(0.02)

1.64(0.08)

0.28(0.03)

98.72(1.15)

5

46.60(0.86)

0.18(0.01)

4.36(0.18)

7.52(0.60)

0.10(0.01)

35.66(0.92)

4.14(0.30)

0.50(0.12)

0.37(0.01)

99.41(0.80)

9

54.74(0.91)

0.22(0.00)

3.61(0.16)

3.08(0.20)

0.04(0.02)

39.52(0.72)

0.80(0.04)

n.d.

0.20(0.01)

102.21(0.52)

3

0.23(0.04)

0.02(0.01)

1.37(0.11)

11.68(0.37)

0.09(0.02)

84.33(0.74)

0.02(0.01)

0.27(0.02)

0.49(0.02)

98.50(0.97)

3

46.45(0.49)

0.16(0.01)

3.92(0.17)

7.03(0.34)

0.10(0.00)

34.87(0.59)

5.03(0.33)

0.27(0.03)

0.31(0.01)

98.12(0.16)

5

54.78(0.92)

0.23(0.03)

3.55(0.11)

3.44(0.32)

n.d.

37.48(0.67)

1.28(0.16)

n.d.

0.30(0.05)

101.06(1.43)

5

0.12(0.02)

n.d.

1.25(0.04)

11.07(0.16)

0.09(0.01)

87.95(0.69)

0.05(0.01)

0.33(0.05)

n.d.

101.00(1.00)

7

45.44(0.21)

0.17(0.01)

4.18(0.07)

7.87(0.11)

0.10(0.01)

34.76(0.22)

4.25(0.07)

0.56(0.05)

0.31(0.01)

97.64(0.25)

9

0.14(0.02)

n.d.

1.56(0.04)

11.04(0.17)

0.08(0.00)

85.53(0.84)

0.03(0.01)

0.30(0.01)

0.14(0.02)

98.80(0.82)

10

53.61(2.16)

0.06(0.01)

4.33(0.55)

6.24(0.46)

0.11(0.02)

33.63(0.93)

1.40 (0.18)

0.18(0.06)

0.20(0.03)

99.75(1.07)

4

0.31(0.14)

n.d.

2.51(0.06)

19.91(0.76)

0.19(0.02)

75.94(2.42)

0.15(0.11)

2.73(0.09)

0.57(0.04)

102.32(2.26)

4

54.95(0.95)

0.21(0.01)

4.03(0.32)

6.98(0.30)

0.10(0.02)

34.04(1.22)

1.28(0.17)

0.21(0.06)

n.d.

101.83(0.56)

4

0.28(0.02)

n.d.

2.53(0.04)

23.54(0.10

0.17(0.01)

72.60(0.95)

0.13(0.03)

2.21(0.05)

0.81(0.01)

102.26(1.07)

3

50.23(0.82)

0.18(0.03)

5.02(0.60)

7.83(0.71)

0.08(0.02)

37.28(1.27)

1.07(0.26)

0.18(0.05)

0.26(0.05)

102.15(0.71)

6

0.45(0.25)

n.d.

2.07(0.05)

20.82(0.31)

0.17(0.02)

76.73(0.54)

0.08(0.01)

1.90(0.05)

0.69(0.04)

102.93(0.49)

3

45.43(0.67)

0.16(0.01)

4.01(0.13)

9.24(0.49)

0.09(0.01)

37.56(0.81)

3.86(0.25)

0.56(0.16)

0.33(0.01)

101.24(0.79)

7

0.16(0.03)

n.d.

1.50(0.06)

11.92(0.31)

0.04(0.02)

87.22(0.68)

0.05(0.01)

0.38(0.03)

0.63(0.28)

101.90(0.79)

3

OT2072A Bdg

Fp

b

)

OT2256A Melt Bdg

25

2473

Fp OT2102A Melt Bdg

25

2498

Fp OT2091A Melt

25

2573

Fp OT2175A Bdg

27

2448

Fp OT2251A Bdg

27

2473

Fp OT2267A Bdg

27

2523

Fp OT2121A Melt Fp

27

2673

30

a

The number of analyses

b

One standard deviation

c

Not detected

31

Table 4. Chemical compositions of recovered samples for N-MORB SiO2

TiO2

Al2O3

FeO

MnO

MgO

CaO

Na2O

Total

Na

Maj

43.41(1.52 b)

0.35(0.15)

21.16(1.98)

10.38(0.29)

0.21(0.03)

15.31(1.09)

7.92(1.27)

1.17(0.21)

99.92(1.30)

6

St

90.75(2.04)

0.17(0.06)

6.67(1.52)

0.38(0.08)

n.d.c

0.23(0.06)

0.69(0.19)

n.d.

98.90(0.77)

3

47.65(1.79)

5.82(0.41)

4.45(0.78)

3.28(0.10)

0.04(0.00)

0.62(0.08)

36.09(1.58)

0.11(0)

98.06(0.76)

2

27.66(1.03)

8.05(0.41)

18.06(0.27)

31.68(0.91)

0.06(0.01)

13.81(0.38)

1.30(0.73)

0.22(0.01)

100.83(0.67)

6

39.71(1.02)

5.60(0.62)

14.03(0.33)

23.42(0.72)

0.24(0.01)

10.06(0.20)

7.03(0.21)

1.38(0.03)

101.47(0.65)

3

48.69(0.50)

2.38(0.17)

3.18(0.16)

2.72(0.30)

0.08(0.00)

1.62(0.07)

42.62(0.25)

0.13(0.01)

101.43(0.61)

3

39.29(2.11)

0.33(0.04)

42.03(1.39)

2.12(0.08)

n.d.

2.81(0.25)

13.58(0.54)

0.47(0.05)

100.65(0.77)

3

St

99.82(1.15)

0.09(0.02)

0.86(0.15)

0.18(0.03)

n.d.

0.05(0.04)

0.35(0.24)

n.d.

101.34(1.12)

4

Maj

45.38(0.97)

0.28(0.04)

19.29(0.60)

9.70(0.32)

0.20(0.03)

17.54(0.72)

7.37(0.36)

1.14(0.14)

100.90(0.58)

4

41.83(1.30)

3.33(0.62)

14.70(0.61)

17.78(1.19)

0.21(0.01)

11.83(0.46)

9.29(0.34)

1.17(0.05)

100.14(0.88)

11

Sample No.

P (GPa)

Nominal T (K)

OT2072B

CaPv

25

2398

Ti- Al-rich Pv OT2256B Melt CaPv CAS

25

2473

OT2102B Melt Maj

45.24(1.41)

0.40(0.15)

18.66(0.52)

9.71(0.67)

0.20(0.01)

15.45(0.63)

8.65(0.61)

1.25(0.24)

99.55(1.75)

7

St

25

2498

97.95(1.94)

0.10(0.02)

2.35(0.89)

0.46(0.08)

n.d.

0.22(0.11)

0.39(0.46)

n.d.

101.47(2.19)

4

CaPv

49.60(0.61)

2.64(0.34)

4.92(1.79)

3.09(0.23)

0.08(0.01)

1.94(0.53)

39.35(1.02)

0.20(0.06)

101.84(0.89)

3

42.08(2.70)

3.62(0.64)

13.49(0.56)

15.27(1.30)

0.19(0.01)

14.62(0.59)

7.89(0.28)

1.11(0.03)

99.58((1.71)

11

OT2091B Melt Maj

46.18(0.80)

0.33(0.03)

20.71(2.70)

5.79(0.54)

0.15(0.02)

19.36(1.92)

7.26(0.74)

0.95(0.07)

100.74(0.67)

12

St

25

2573

97.23(0.77)

0.05(0.01)

0.89(0.31)

0.11(0.02)

n.d.

0.13(0.11)

0.14(0.06)

n.d.

98.57(1.25)

3

CaPv

47.89(0.14)

1.27(0.08)

3.28(0.01)

1.36(0.04)

0.07(0.01)

1.85(0.01)

42.08(0.36)

0.10(0.00)

97.89(0.30)

3

42.09(0.89)

1.01(0.03)

15.54(1.19)

15.54(0.34)

0.27(0.02)

20.72(0.84)

2.29(0.78)

0.27(0.03)

97.74(0.23)

3

95.51(2.34)

0.10(0.05)

1.34(0.80)

0.65(0.21)

0.01(0.01)

0.29(0.19)

1.79(1.02)

0.04(0.03)

99.74(0.71)

3

CaPv

49.45(2.65)

1.12(0.17)

4.60(1.09)

4.96(0.11)

0.11(0.02)

2.28(0.45)

33.34(3.11)

0.51(0.03)

96.37(0.14)

2

CAS

40.88(1.55)

0.20(0.04)

41.54(1.96)

2.04(0.31)

0.02(0.01)

3.02(0.21)

12.99(0.46)

1.01 (0.15)

101.73(1.55)

7

OT2175B Al-Bdg St

27

2448

32

OT2251B Melt

43.87(0.92)

2.48(0.12)

15.95(0.87)

15.27(0.53)

0.22(0.01)

12.81(0.66)

7.95(0.21)

1.09(0.09)

99.64(0.53)

11

49.11(0.40)

1.35(0.04)

3.38(0.39)

2.34(0.06)

0.09(0.00)

2.10(0.02)

42.05(1.10)

0.23(0.08)

100.65(0.80)

3

CAS

38.98(0.80)

0.24(0.08)

42.71(0.79)

1.98(0.43)

0.02(0.01)

3.51(0.27)

13.79(0.17)

0.53(0.02)

101.75(0.83)

4

St

98.02(0.96)

0.05(0.00)

0.68(0.05)

0.21(0.01)

n.d.

0.03(0.03)

0.14(0.03)

n.d.

99.13(0.91)

4

47.05(2.29)

2.09(0.28)

15.31(0.89)

11.72(1.73)

0.19(0.02)

16.65(1.80)

7.96(0.76)

0.94(0.48)

101.92(0.75)

10

CaPv

27

2473

OT2267B Melt CaPv

50.34(1.49)

1.14(0.07)

3.04(0.28)

1.77(0.06)

0.08(0.01)

2.59(0.15)

43.49(0.34)

0.11(0.01)

102.58(1.56)

5

CAS

27

2523

37.49(0.74)

0.17(0.01)

42.77(0.58)

1.42(0.11)

0.02(0.01)

3.22(0.19)

14.05(0.13)

0.49(0.04)

99.62(0.65)

4

St

99.02(1.49)

0.05(0.02)

1.78(0.79)

0.18(0.09)

n.d.

0.24(0.12)

0.61(0.60)

n.d.

101.88(0.85)

6

46.96(0.53)

1.65(0.23)

15.86(1.36)

9.03(1.12)

0.13(0.01)

15.40(1.17)

11.80(0.39)

0.68(0.17)

101.52(0.35)

11

46.11(0.74)

0.90(0.02)

4.34(0.06)

1.11(0.04)

0.01(0.01)

2.36(0.09)

43.09(0.32)

0.09(0.01)

98.01(0.60)

5

OT2121B Melt

27

CaPv a

The number of analyses

b

One standard deviation

c

Not detected

2673

33

Table 5. Phase proportions by mass balance calculation a Sample No.

P

Nominal

(GPa)

T (K)

25

2473

Bdg b

Fp

0.109(0.035

0.046(0.004)

CaPv

Maj

CAS

St

Ti-MgPv

Melt

R2 c

-

-

-

-

0.831(0.041)

1.000

KLB-1 peridotite OT2256A

d

)

-

OT2102A

25

2498

0.138(0.103)

0.051(0.010)

-

-

-

-

-

0.812(0.120)

0.999

OT2091A

25

2573

-

0.038(0.007)

-

-

-

-

-

0.993(0.010)

1.000

OT2121A

27

2673

-

0.004(0.009)

-

-

-

-

-

0.993(0.013)

0.999

OT2072B

25

2398

-

-

0.190(0.021)

0.523(0.046)

-

0.159(0.017)

0.124(0.030)

-

0.999

OT2256B

25

2473

-

-

0.123(0.005)

0.318(0.015)

0.127(0.007)

0.131(0.004)

-

0.277(0.011)

1.000

OT2102B

25

2498

-

-

0.115(0.041)

0.460(0.147)

-

0.104(0.031)

-

0.309(0.146)

0.997

OT2091B

25

2573

-

-

0.132(0.064)

0.247(0.190)

-

0.105(0.052)

-

0.514(0.217)

0.992

OT2175B

27

2448

0.434(0.051)

-

0.247(0.042)

-

0.180(0.038)

0.126(0.028)

-

trace

0.998

OT2251B

27

2473

-

-

0.108(0.001)

-

0.122(0.002)

0.116(0.001)

-

0.640(0.003)

1.000

OT2267B

27

2523

-

-

0.102(0.043)

-

0.140(0.053)

0.105(0.038)

-

0.611(0.090)

0.996

-

-

0.044(0.077)

-

-

-

-

0.967(0.091)

0.984

N-MORB

OT2121B a

Because the grain size of CaPv was smaller than a focused analytical spot of EPMA, we cannot determine the phase proportion of Run No. OT2172A, OT2175A, OT2251A, and

OT2267A. b

Bdg, bridgmanite; Fp, ferropericlase; Maj, majorite; St, stishovite; CaPv, calcium perovskite; CAS, Ca-Al phase; Ti-MgPv, Ti-rich magnesium perovskite

c

The coefficient of determination

d

One standard error

34

Figure 1. Back scattered electron images of the recovered sample of Run No. OT2072 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d) Magnified view of recovered KLB-1 peridotite sample. (e) Magnified view of NMORB sample.

35

Figure 2. Back scattered electron images of the recovered sample of Run No. OT2256 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d, e), Magnified views of recovered KLB-1 peridotite N-MORB samples.

36

Figure 3. Back scattered electron images of the recovered sample of Run No. OT2102 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d, e), Magnified views of recovered KLB-1 peridotite N-MORB samples.

37

Figure 4. Back scattered electron images of the recovered sample of Run No. OT2091 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d, e), Magnified views of recovered KLB-1 peridotite N-MORB samples.

38

Figure 5. Back scattered electron images of the recovered sample of Run No. OT2175 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d, e), Magnified views of recovered KLB-1 peridotite N-MORB samples.

39

Figure 6. Back scattered electron images of the recovered sample of Run No. OT2251 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d, e), Magnified views of recovered KLB-1 peridotite N-MORB samples.

40

Figure 7. Back scattered electron images of the recovered sample of Run No. OT2267 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d, e), Magnified views of recovered KLB-1 peridotite N-MORB samples.

41

Figure 8. Back scattered electron images of the recovered sample of Run No. OT2121 (a), The overview (b, c), back scattered electron images of recovered KLB-1 peridotite and N-MORB samples. (d, e), Magnified views of recovered KLB-1 peridotite N-MORB samples.

42

Figure 9. Comparison of possible solidus temperature regions. Each region is represented by connecting lines of partial melting and sub-solidus temperatures reported by previous studies. Filled triangle and open inverted triangle indicate sub-solidus and above solidus, respectively. a) dry KLB-1 peridotite. Green shaded region indicates possible solidus temperature regions. b) dry (blue shaded region) and hydrous (grey shaded region) N-MORB. c) The comparison of this study with previous studies. Filled circle and open circle indicate sub-solidus and above solidus, respectively. Half-filled circle indicates sub-solidus KLB-1 peridotite and partial melting of N-MORB. Data for KLB-1 peridotite were taken from Ito and Takahashi (1987), Zhang and Herzberg (1994), Hirose and Fei (2002), Trønnes and Frost (2002), and Ito et al. (2004). Data for N-MORB were taken from Yasuda et al. (1994), Hirose and Fei (2002), and Litasov and Ohtani (2005). It should be noted that the data containing trace elements, such as metallic iron, were excluded.

43

Figure 10. The degree of melting for KLB-1 peridotite and N-MORB at 25–27 GPa as a function of temperature. Green and blue squares indicate KLB-1 peridotite and N-MORB, respectively. The error bar indicates 1 standard error.

44

Figure 11. Calculated densities of the partial melts of a) KLB-1 peridotite and b) N-MORB along the mantle geotherm estimated from Katsura et al. (2010). (c) Calculated density of partial melts for KLB-1 peridotite along the adiabat with the slope of 10 K/GPa at the mantle potential temperature Tp = 2273 K.

45

Highlights Solidus temperatures of KLB-1 peridotite and N-MORB were simultaneously determined under the identical pressure-temperature conditions. Solidus temperature of N-MORB is nearly identical to the KLB-1 peridotite at 25 GPa but lower at 27 GPa. The crossover of melt fractions between KLB-1 peridotite and N-MORB occurs at 25– 27 GPa. Partial melts of N-MORB may be gravitationally stable at the uppermost lower mantle, whereas partial melts of KLB-1 peridotite may be stable at the top of the transition zone. The distribution of partial melts in the deep Earth's mantle may be different between KLB-1 peridotite and N-MORB.

46