Post-spinel transitions in pyrolite and Mg2SiO4 and akimotoite–perovskite transition in MgSiO3: Precise comparison by high-pressure high-temperature experiments with multi-sample cell technique

Earth and Planetary Science Letters 309 (2011) 185–197

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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l

Post-spinel transitions in pyrolite and Mg2SiO4 and akimotoite–perovskite transition in MgSiO3: Precise comparison by high-pressure high-temperature experiments with multi-sample cell technique Takayuki Ishii, Hiroshi Kojitani, Masaki Akaogi ⁎ Department of Chemistry, Gakushuin University, 1-5-1, Mejiro, Toshima-ku, Tokyo 171-8588, Japan

a r t i c l e

i n f o

Article history: Received 5 March 2011 Received in revised form 22 June 2011 Accepted 23 June 2011 Available online 30 July 2011 Editor: L. Stixrude Keywords: pyrolite post-spinel transition ringwoodite akimotoite perovskite 660-km discontinuity

a b s t r a c t We precisely compared phase boundaries of post-spinel transition in pyrolite and Mg2SiO4 and of akimotoite– perovskite transition in MgSiO3 at 21–28 GPa and 1400–1800 °C by detailed phase relation experiments using a multi-anvil apparatus. We used a multi-sample cell technique, in which pyrolite, Mg2SiO4 and MgSiO3 were kept simultaneously at the same pressure–temperature conditions in each run. The experiments were performed in pressure and temperature intervals of 0.3 GPa and 100 °C, respectively. The post-spinel transition boundary in Mg2SiO4 is located at higher pressure by about 0.8 GPa than the akimotoite–perovskite transition boundary in MgSiO3. Both the transition boundaries have the same slope of −0.002 GPa/°C. In pyrolite, the post-spinel transition occurs in a pressure interval within 0.4 GPa at lower pressure by about 0.2–1.0 GPa than that in Mg2SiO4 at 1400–1800 °C. The Clapeyron slope of the post-spinel transition boundary in pyrolite is −0.001 GPa/°C, which is half of −0.002 GPa/°C of Mg2SiO4. When we assume that both the transition zone and the uppermost lower mantle have approximately pyrolitic composition, the above results imply that the Clapeyron slope of the transition boundary in pyrolite is more appropriate than that of Mg2SiO4 to evaluate effects of the post-spinel transition on mantle dynamics and the 660-km discontinuity topography. In pyrolite, the akimotoite–perovskite transition and the post-spinel transition occur at the same pressure at 1400 °C. Above 1700 °C, a part of ringwoodite in pyrolite transforms to garnet + magnesiowüstite at pressure below the postspinel transition, and abundances of garnet and magnesiowüstite increase with increasing temperature. © 2011 Elsevier B.V. All rights reserved.

1. Introduction It is generally accepted that the 660-km depth seismic discontinuity dividing the upper and lower mantle is attributed to the postspinel transition in which ringwoodite (spinel) transforms to the assemblage of Mg-rich perovskite and magnesiowüstite. Because the post-spinel transition has a negative Clapeyron slope, it is widely believed that the transition works as a partial barrier for subduction of slabs which affects mantle dynamics and that the post-spinel transition causes elevation and depression of the 660-km discontinuity due to lateral temperature variations. Therefore, a number of investigations have been made to examine the post-spinel phase relations in pyrolite at the pressure–temperature conditions around the 660-km depth (Hirose, 2002; Irifune, 1994; Litasov et al., 2005a; Nishiyama et al., 2004; Nishiyama and Yagi, 2003; Wood, 2000). The ringwoodite in pyrolite has an approximate composition of (Mg0.9, Fe0.1)2SiO4 at the transition zone conditions. Therefore, as an appropriate model, the post-spinel transition of Mg2SiO4 has been frequently studied by high pressure experiments with quenching

⁎ Corresponding author. Tel.: + 81 3 3986 0221; fax: + 81 3 5992 1029. E-mail address: [email protected] (M. Akaogi). 0012-821X/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2011.06.023

method (Ito and Takahashi, 1989) and in situ X-ray observations (Fei et al., 2004b; Irifune et al., 1998; Katsura et al., 2003) and by thermodynamic calculations (Akaogi et al., 2007, 2008; Akaogi and Ito, 1993; Ito et al., 1990). The Mg2SiO4 post-spinel transition has also been investigated by ab intio calculation (Yu et al., 2007). In the course of studies, discrepancy in pressure between the Mg2SiO4 post-spinel transition observed by the high-pressure experiments and the 660-km depth was extensively discussed in terms of uncertainties of pressure determination with various pressure scales (e.g., Fei et al., 2004a,b; Irifune et al., 1998). Also, for the Clapeyron slope of the Mg2SiO4 postspinel transition, several different values ranging − 0.003 to −0.0004 GPa/°C have been reported, based on quenching and in situ X-ray diffraction methods (Fei et al., 2004b; Irifune et al., 1998; Ito and Takahashi, 1989; Katsura et al., 2003) and by thermodynamic calculations (Akaogi et al., 2007, 2008; Akaogi and Ito, 1993; Ito et al., 1990). Because the value of the Clapeyron slope gives a profound implication for mantle dynamics and the 660-km discontinuity topography, the difference in the Clapeyron slope has been widely discussed (Fei et al., 2004b; Katsura et al., 2003; Litasov et al., 2005a). Although precise determination of the post-spinel transition boundary of Mg2SiO4 has been made by a number of investigators as shown above, it has not yet been fully clarified how much differences may exist in the transition pressure and the Clapeyron slope of the

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post-spinel transition between pyrolite and Mg2SiO4. Minor components such as Fe 2+ in pyrolite may cause the differences. In addition, compared with Mg2SiO4, phase relations of pyrolite around the 660km depth conditions are complex due to formation of akimotoite (ilmenite) at relatively low temperature and increase of abundance of majorite garnet at relatively high temperature (Hirose, 2002). The akimotoite–perovskite transition in MgSiO3 has been widely used as the pressure standard at high temperature at around 23 GPa, which is almost equal to the pressure of the 660-km depth (Hirose et al., 2001b; Ito and Takahashi, 1989; Ono et al., 2001). Except for a few studies (Fei et al., 2004b; Kuroda et al., 2000), pressure difference between the akimotoite–perovskite transition in MgSiO3 and the post-spinel transition in Mg2SiO4 has not yet been precisely examined by direct high-pressure experiments. In this study, we have examined detailed phase relations in pyrolite, Mg2SiO4 and MgSiO3 at 21–28 GPa and 1400–1800 °C, using a Kawai-type 6–8 multianvil apparatus. We have used a multi-sample cell technique (Fei and Bertka, 1999) to compare precisely the phase relations among the three samples. In the method, the three starting materials (pyrolite, Mg2SiO4 and MgSiO3) were loaded in three different small holes in a single metal capsule in a pressure medium, and were kept simultaneously at high-pressure high-temperature conditions. We have conducted the high-pressure high-temperature runs in very small pressure interval of 0.3 GPa and temperature interval of 100 °C to determine small differences in the transition pressures and in the boundary slopes among the postspinel transitions in pyrolite and Mg2SiO4 and the akimotoite–perovskite transition in MgSiO3. Mineral abundances in pyrolite have also been determined on the basis of composition analysis of coexisting phases, and significance of the mineral transitions around the 660-km discontinuity is discussed. 2. Experimental methods 2.1. Starting materials For syntheses of starting materials, reagent-grade chemicals were used throughout. Mg2SiO4 forsterite was synthesized from a mixture of MgO and SiO2 with a 2:1 mol ratio by heating at 1500 °C for 75 h. MgSiO3 enstatite was synthesized as follows. A mixture of MgO and SiO2 with a 1:1 mol ratio was heated at 1670 °C for 1 h and quenched to form MgSiO3 glass. The glass was crystallized into enstatite by heating at 1300 °C for 63 h. Composition of pyrolite by McDonough and Sun (1995) was adopted: SiO2(44.98), TiO2(0.20), Al2O3(4.45), Cr2O3(0.38), FeO(8.05), NiO(0.25), MgO(37.78), CaO(3.55) and Na2O(0.36), where numbers in parentheses are contents in wt.%, excluding very small amounts of MnO, K2O and P2O5. The starting material of pyrolite was prepared in the same composition as above by mixing Mg2SiO4(40.2), MgSiO3(37.5), Fe2SiO4(7.1), CaSiO3(8.0), NaAlSiO4(1.5), TiO2(0.3), Al2O3(4.8), Cr2O3(0.3) and NiO(0.4), where numbers in parentheses are contents in mol%. Fe2SiO4 fayalite was synthesized from a mixture of Fe2O3 and SiO2 with a 1:1 mol ratio by heating at 1180 °C for 24 h in a controlled oxygen fugacity using a mixture of H2, CO2 and Ar with volume ratios of 1:1:2. CaSiO3 pseudo-wollastonite was synthesized from a mixture of CaCO3 and SiO2 with a 1:1 mol ratio by heating at 1450 °C for 131 h. NaAlSiO4 carnegieite was synthesized from a mixture of Na2CO3, Al2O3 and SiO2 with a 1.05:1:2 mol ratio by heating for 34 h at 1300 °C. The addition of 5 mol% Na2CO3 was to avoid Na-loss during the heating process (Kojitani et al., 2011). Identification of these synthetic compounds and confirmation of single-phase materials were made using a powder X-ray diffractometer and a scanning electron microscope with an energy dispersive X-ray spectrometer (SEM-EDS). 2.2. High-pressure experiments High-pressure experiments were made at 21.1–28.0 GPa and 1400– 1800 °C with a Kawai-type 6–8 multianvil apparatus at Gakushuin

University. Tungsten carbide anvils with a truncated edge length of 2.5 mm were used in combination with a semi-sintered Cr2O3–doped MgO octahedron of 7 mm on edge. In each run, pressure was raised to a targeted pressure at almost constant rate during about 2–4 h, and temperature was raised to 1400–1800 °C at a rate of about 100 °C/min. The sample assembly was kept for 1–6 h at the pressure–temperature conditions, then quenched, and recovered to ambient conditions. Pressure calibration at room temperature was made by using pressurefixed points of ZnS (15.5 GPa), GaAs (18.3 GPa) and GaP (23 GPa) (Dunn and Bundy, 1978; Ito, 2007). The pressure calibration at 1600 °C was made by using wadsleyite–ringwoodite transition in Mg2SiO4 (21.3 GPa) (Suzuki et al., 2000), akimotoite–perovskite transition in MgSiO3 (22.3 GPa) (Fei et al., 2004b; Hirose et al., 2001b), transition of corundum+periclase to MgAl2O4 calcium ferrite (24.9 GPa) (Irifune et al., 2002), and transition of Mg3Al2Si3O12 pyrope to perovskite+ corundum (26.5 GPa) (Fei et al., 2004b; Hirose et al., 2001a). The pressure calibration was also made at 1400 °C and 1800 °C in a similar manner to that at 1600 °C: at 1400 °C by akimotoite–perovskite transition in MgSiO3 (22.9 GPa) (Fei et al., 2004b; Hirose et al., 2001b) and transition of corundum+periclase to MgAl2O4 calcium ferrite (25.2 GPa) (Irifune et al., 2002), and at 1800 °C corundum+periclase=MgAl2O4 calcium ferrite (24.5 GPa) (Irifune et al., 2002). In the pressure calibration at high temperatures, the starting material was put into a hole of a Re capsule in the multi-sample chamber method described below. Reproducibility of pressure in the present quench experiments was estimated to be less than about ±0.2 GPa. Before packing samples in the Re capsule, the magnesia octahedron was dried at 100 °C over-night in a furnace. The three dried starting materials (pyrolite, Mg2SiO4 forsterite, and MgSiO3 enstatite) were packed in three holes of 0.2 mm diameter in the single Re capsule which was 1.0 mm in diameter and 0.7 mm in thickness. Two Re disks of 1.0 mm in diameter and 0.1 mm in thickness were put on both sides of the Re capsule. In some runs, corundum + periclase and/or pyrope were packed in one or two holes of the Re capsule for pressure calibration. The Re capsule with the lids was put in the central part of a tubular rhenium heater in the pressure medium octahedron. A boron nitride capsule and LaCrO3 end-plugs were inserted between the Re capsule + lids and the heater for electrical insulation. A LaCrO3 sleeve was placed outside of the rhenium furnace for thermal insulation. The three starting materials were kept simultaneously at desired pressure–temperature conditions for 1–6 h, then quenched, and recovered to ambient conditions. The run temperature was measured at the central part of outer surface of the rhenium heater by a Pt/Pt-13%Rh thermocouple of 0.1 mm in diameter. Pressure effect on thermoelectromotive force of the thermocouple was ignored. The recovered Re capsule holding the samples was mounted on a slide glass plate with epoxy resin, and polished to expose the samples in the Re capsule for phase identification and compositional analysis. 2.3. Analyses of run products The microfocus X-ray diffractometer and powder X-ray diffractometer (Rigaku RINT 2500 V) were used to identify the recovered run products as well as the starting materials. CrKα radiation with a rotating anode was used at 45 kV and 250 mA. Phase identification of the recovered run products was made near the central part of each hole of the Re capsule using the microfocus X-ray diffractometer with a collimated X-ray beam of 50 μm diameter. The scanning electron microscope (JEOL JMS-6360) with the energy dispersive X-ray spectrometer (Oxford INCA x-sight) was used for chemical composition analysis of phases of the run products, as well as for phase identification. The scanning electron microscope was operated with acceleration voltage of 15 kV and probe current of 0.43 nA. Each phase was analyzed by a fine-focused electron beam of about 1 μm in diameter. Composition analysis was made generally at 4–20 analysis points for each phase, and the average data were shown together with two standard deviations of the mean. Standard materials were natural

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samples of forsterite for Mg and Si, jadeite for Na, wollastonite for Ca, and fayalite for Fe, in addition to synthetic compounds of corundum for Al, rutile for Ti, eskolaite (Cr2O3) for Cr, and nickel oxide for Ni. 3. Results and discussion Fig. 1 shows the pressure calibration curve used in this study. As described above, to calibrate pressure at high temperatures, we used the boundaries of Mg2SiO4 wadsleyite and ringwoodite transition (Suzuki et al., 2000), MgSiO3 akimotoite–perovskite transition (Fei et al., 2004b; Hirose et al., 2001b), dissociation of pyrope to MgSiO3-rich perovskite + corundum (Fei et al., 2004b; Hirose et al., 2001a), and transition of periclase+ corundum to MgAl2O4 calcium ferrite (Irifune et al., 2002). We adopt the pressure calibration curve in Fig. 1 based on the above three boundaries in order to determine mutual, relative differences of the transition pressures among pyrolite, Mg2SiO4 and MgSiO3. The pressure calibration curves determined at 1400 and 1800 °C using the above three boundaries were consistent with that at 1600 °C within the experimental errors. Table 1 summarizes experimental results of the runs for pyrolite, Mg2SiO4 and MgSiO3, together with those for the pressure calibration. Fig. 2 shows the high-pressure cell assembly used for the multi-sample cell experiments. All of pyrolite, Mg2SiO4 and MgSiO3 in the run products were well crystallized. The phases identified by microfocus X-ray diffraction and SEM-EDS analysis are summarized in Table 1, and are shown in Figs. 3–5. The analyzed compositions of coexisting minerals in some run products are listed in Table 2. Fig. 3 illustrates results of Mg2SiO4. With increasing pressure, Mg2SiO4 ringwoodite dissociates into perovskite+ periclase assemblage at 1400–1700 °C, while above 1800 °C wadsleyite directly transforms to the perovskite + periclase without intervening ringwoodite. The triple point among ringwoodite, wadsleyite, and perovskite + periclase is at about 23.3 GPa and 1800 °C with the estimated uncertainties of pressure and temperature of ±0.3 GPa and ±40 °C, respectively. The post-spinel transition boundary in Mg2SiO4 in Fig. 3 is represented as P(GPa) = 27.0–0.002 T(°C), based on the pressure calibration described above. The error of the slope was estimated as ±0.0007 GPa/°C. Although the post-spinel transition boundary of Mg2SiO4 was not determined by in situ X-ray diffraction method in this study, the slope of −0.002 GPa/°C is within the range of slopes of the previous studies from −0.003 GPa/°C (Irifune et al., 1998; Ito et al., 1990) to −0.0004 GPa/°C (Katsura et al., 2003), and is comparable with −0.0026 GPa/°C by the thermodynamic calculation on the basis of enthalpy and entropy measurements (Akaogi et al., 2007). Fig. 4 shows phase relations in MgSiO3. MgSiO3 akimotoite transforms to perovskite at 1400–1700 °C, while majorite directly transforms

30

Pressure/GPa

25

GaP

20 15

GaAs Mg3 Al2 Si 3 O12 Gt Mpv+Co MgAl 2 O4 Co+Pc CF MgSiO3 Ak Mpv Mg 2 SiO 4 Wd Rw

ZnS

10 5 0 0

100

200

300

Oil pressure/kgf

400

500

cm-2

Fig. 1. Pressure calibration curves at room temperature and 1600 °C. A solid line and a dashed line represent pressure calibration curves at 1600 °C and room temperature, respectively. Diamonds show the transitions at room temperature, and other symbols represent the runs at 1600 °C. Gt, garnet; Mpv, MgSiO3-rich perovskite; Co, corundum; CF, calcium ferrite; Pc, periclase; Ak, akimotoite; Wd, wadsleyite; Rw, ringwoodite.

187

to perovskite at 1800 °C. At 1600–1700 °C, majorite is stable at lower pressure than akimotoite. The triple point among majorite, akimotoite and perovskite is located at about 22.5 GPa and 1770 °C with the estimated uncertainties of pressure and temperature of ±0.3 GPa and ±30 °C, respectively. The akimotoite–perovskite transition boundary is expressed by P(GPa)= 26.1–0.002 T(°C). The error of the slope was estimated to be ±0.001 GPa/°C. Pressure difference between the postspinel transition in Mg2SiO4 and the akimotoite–perovskite transition in MgSiO3 is about 0.8 GPa, and both the boundaries have the same slope within the experimental errors. The slope of the akimotoite–perovskite transition in this study is close to −0.0024 GPa/°C by the thermodynamic calculation based on entropy measurements of akimotoite and perovskite of MgSiO3 (Akaogi et al., 2008). The above results on Mg2SiO4 and MgSiO3 in this study are generally consistent with those by Kuroda et al. (2000). They carried out highpressure quench experiments, in which Mg2SiO4 and MgSiO3 were put separately in the same sample chamber in the high-pressure cell assembly, and determined both the transition boundaries where pressure was estimated on Mg2SiO4 post-spinel transition boundary on the basis of Anderson et al.'s (1989) gold pressure scale. Kuroda et al. (2000) observed that the akimotoite–perovskite transition boundary was placed at lower pressure by 0.5–1.0 GPa than the post-spinel boundary and that the two boundary slopes were around −0.003 GPa/°C. Fig. 5 shows phase relations in pyrolite at 21.1–28.0 GPa and 1400–1800 °C. For comparison with pyrolite, the transition boundaries of Mg2SiO4 in Fig. 3 are reproduced in Fig. 5. Fig. 6 shows backscattered electron images of typical run products of pyrolite which show six different kinds of the phase assemblages. The minerals in the run products were crystallized generally in size of about 5–10 μm, and the analyzed compositions indicate that the run products were close to chemical equilibrium in the experimental conditions, as shown below. Because CaSiO3-rich perovskite (Cpv) is unquenchable and transforms to aggregate of amorphous phases, it was identified only by SEM-EDS analysis. Compared with the other phases, the compositions of CPv were relatively less accurate because of the aggregate texture and the small-sized amorphized grains. Fig. 7(a) shows changes in microfocus X-ray diffraction patterns with pressure at 1600 °C in the runs from 22.9 GPa to 27.6 GPa. Based on the microfocus X-ray diffraction and SEM-EDS analysis, mineral assemblages at 1600 °C were ringwoodite (Rw) + majorite garnet (Gt) + Cpv at 22.9 GPa, Rw + MgSiO3-rich perovskite (MPv) + magnesiowüstite (Mw) + Gt + Cpv at 23.3 GPa, Mpv + Mw + Gt + Cpv at 23.8 GPa, and Mpv + Mw + Cpv at 27.6 GPa. The results indicate that the post-spinel transition of Rw to Mpv+ Mw occurs in a very narrow pressure interval less than 1 GPa at 1600 °C. At 1400 °C, a part of Gt transforms to akimotoite (Ak) at 22.9 GPa from the assemblage of Rw + Gt + Cpv, as shown in Fig. 7(b) (run no. 32). The assemblage of Rw + Ak + Gt + Cpv changes to Mpv + Mw + Gt+ Cpv (no. 26 in Fig. 7(b)) via the five-phase assemblage Rw + Mpv + Mw + Gt + Cpv (no. 27) at 23.3–23.8 GPa at 1400 °C. No Ak was observed in the runs at 1500–1800 °C. At temperatures of 1700 and 1800 °C at 21.1–22.9 GPa, abundance of Gt increases, that of Rw decreases, and Mw appears in the run products, as shown in Fig. 7(c). This indicates that a reaction of Rw → Gt+ Mw occurs with increase of temperature. The assemblage transforms to Mpv+ Mw + Gt + Cpv via Rw+ Mpv + Mw + Gt + Cpv with increasing pressure. In the whole temperature range of experiments in this study, Gt completely transforms to Mpv and Cpv at about 24.5–26 GPa. The boundary has a positive slope, 0.004 GPa/°C, as shown in Fig. 5. To examine reproducibility of the present results on the phase relations in pyrolite, Mg2SiO4 and MgSiO3, we conducted double runs at the same P–T conditions in the pressure range of 22.9–23.8 GPa at 1500– 1800 °C across the post-spinel transition. As shown in Table 1, for example, two experiments (run nos. 18 and 20) at 22.9 GPa and 1800 °C gave the identical phase assemblages in each of the three samples. Two runs (nos. 19 and 47) at 23.3 GPa and 1800 °C also showed the other

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Table 1 Experimental conditions and run products. Run no.

55 32 27 26 45 10 21 75 24 77 66 71 33 28 31 52 48 44 35 37 50 15 16 39 17 38 13 9 12 67 4 64 63 62 61 59 57 54 58 36 29 43 53 46 51 78 23 18 20 47 19 41 72 73 40 76 69 68 65 60

Oil pressure (kgf cm− 2)

Pressure (GPa)

Temperature (°C)

Time (min)

Phases Pyrolite

Fo

En

270 285 295 310 320 335 350 360 370 380 400 430 285 285 295 310 310 320 335 245 260 270 285 285 295 310 310 320 335 340 350 360 380 400 430 460 500 550 245 270 285 285 295 310 310 245 270 285 285 295 295 310 325 340 350 350 370 400 450 550

22.3 22.9 23.3 23.8 24.1 24.6 25.0 25.3 25.5 25.8 26.2 26.7 22.9 22.9 23.3 23.8 23.8 24.1 24.6 21.1 21.8 22.3 22.9 22.9 23.3 23.8 23.8 24.1 24.6 24.7 25.0 25.3 25.8 26.2 26.7 27.1 27.6 28.0 21.1 22.3 22.9 22.9 23.3 23.8 23.8 21.1 22.3 22.9 22.9 23.3 23.3 23.8 24.3 24.7 25.0 25.0 25.5 26.2 27.0 28.0

1400 1400 1400 1400 1400 1400 1400 1400 1400 1400 1400 1400 1500 1500 1500 1500 1500 1500 1500 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1600 1700 1700 1700 1700 1700 1700 1700 1800 1800 1800 1800 1800 1800 1800 1800 1800 1800 1800 1800 1800 1800 1800

360 360 360 180 360 180 180 180 360 360 360 360 180 360 180 180 360 180 180 180 180 180 180 180 180 180 180 180 180 180 60 180 180 180 180 180 180 180 180 120 360 120 180 120 360 120 120 60 60 120 60 120 120 120 120 120 120 120 120 120

Rw + Gt + Cpv Rw + Gt + Ak + Cpv Rw + Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Rw + Gt + Cpv Rw + Gt + Cpv Rw + Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Rw + Gt Rw + Gt + Cpv Rw + Gt + Cpv Rw + Gt + Cpv Rw + Gt + Cpv Rw + Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Rw + Mw + Gt + Cpv Rw + Mw + Gt + Cpv Rw + Mw + Gt + Cpv Rw + Mw + Gt + Cpv Rw + Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Rw + Mw + Gt Rw + Mw + Gt + Cpv Rw + Mw + Gt + Cpv Rw + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Gt + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv Mpv + Mw + Cpv

Rw Rw Rw Rw Rw + Mpv + Pc Mpv + Pc Mpv + Pc (Co + Pc) Mpv + Pc (CF + Co + Pc) Mpv + Pc Mpv + Pc Rw Rw Rw Rw + Mpv + Pc Rw + Mpv + Pc Mpv + Pc Mpv + Pc Wd† Rw† Rw Rw Rw Rw Rw + Mpv + Pc Rw + Mpv + Pc Mpv + Pc Mpv + Pc (Co + Pc)† Mpv + Pc (CF + Co + Pc)† (CF + Co + Pc) (CF + Co + Pc) (CF + Co + Pc) (CF + Co + Pc) (CF + Co + Pc) (CF + Co + Pc) Wd Wd Rw Rw Rw + Mpv + Pc Mpv + Pc Mpv + Pc Wd Wd Wd + Mpv + Pc Wd + Mpv + Pc Rw + Mpv + Pc Rw + Mpv + Pc Mpv + Pc (Co + Pc) (Co + Pc) Mpv + Pc (Cf + Co + Pc) Mpv + Pc Mpv + Pc Mpv + Pc Mpv + Pc

Ak Ak Ak + Mpv Mpv Mpv Mpv Mpv (Gt) Mpv (Gt + Mpv + Co) Mpv Mpv Ak Ak Mpv Mpv Mpv Mpv Mpv Ak + Mj Ak Ak† Ak + Mpv† Ak + Mpv Mpv† Mpv Mpv Mpv Mpv (Gt) Mpv (Gt) (Gt)† (Gt + Mpv + Co)† (Gt + Mpv + Co)† (Gt + Mpv + Co) (Mpv + Co) (Mpv + Co) Mj Ak Mpv Mpv Mpv Mpv Mpv Mj Mj Mpv Mpv Mpv Mpv Mpv (Gt) (Gt) Mpv (Gt) Mpv Mpv Mpv Mpv

Abbreviations: Rw, ringwoodite; Mpv, MgSiO3-rich perovskite; Mw, magnesiowustite; Gt, majorite garnet; Cpv, CaSiO3-rich perovskite; Ak, akimotoite; Wd, wadslyite; Pc, periclase; Mj, MgSiO3-rich garnet; Co, Al2O3 corundum; CF, MgAl2O4 calcium ferrite. Phases in parentheses are results of MgAl2O4 (left) and pyrope (right) packed in the Re capsule. Gt in parentheses: pyrope garnet. † The runs used for pressure calibration in Fig. 1.

identical phase assemblages. These results strongly suggest that the reproducibility of pressure was less than about ±0.2 GPa. Other doubleruns (nos. 28 and 33; 48 and 52; 16 and 39; 13 and 38; 46 and 51) in Table 1 also expressed the identical phase assemblages at 22.9–23.8 GPa and 1500–1700 °C.

The compositions of phases in the run products in Table 2 are used to examine whether the phases in the run products were in chemical equilibrium or not. The compositions of the phases in the run at 22.9 GPa and 1700 °C for 2 h (run no. 43) and 6 h (no. 29) were generally in good agreement. The compositions in the two runs (nos.

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189

Fig. 2. (a) A schematic illustration of a cross section of the cell assembly for the multi-sample cell high-pressure experiments. (b) Re capsule in the cell assembly.

46 and 51) at 23.8 GPa and 1700 °C for 2 and 6 h, respectively, were generally consistent with each other. These results suggest that the phase compositions were close to equilibrium. The apparent partition coefficient of Mg and Fe between MPv and Mw is defined by: K

MPv−Mw

MPv

¼ ðXFe =XMg Þ

Mw

=ðXFe =XMg Þ

ð1Þ

where XFe = (Fe2+ + Fe3+) / (Mg2+ + Fe2+ + Fe3+). In our experiments, the KMPv − Mw values were 0.66–0.73 at 24.1–25.3 GPa and 1600 °C (nos. 9 and 64) and 0.70–0.71 at 23.8–24.7 GPa at 1800 °C (nos. 41 and 73). Frost and Langenhorst (2002) reported the K MPv − Mw of 0.69 ± 0.16 at 24–25 GPa and 1650–1900 °C in the Mg-rich compositions with no significant temperature dependence. Therefore, our Mg–Fe partition coefficients are consistent with those by Frost and Langenhorst (2002). The results again suggest that the run products were close to equilibrium in our experimental conditions. Fig. 8 illustrates compositional changes of the phases as a function of pressure at 1800 °C (Table 2). Cation numbers in Gt and Rw are almost constant below 22.3 GPa. However, at pressure above 23.7 GPa, Mg and Si numbers decrease and Al increases in Gt with increasing pressure, and the similar changes occur in MPv. These compositional changes in Gt and MPv correspond to the garnet– perovskite transition. As shown in Fig. 8 and discussed below, Rw and Mw coexist at 21–23 GPa range at 1800 °C. The Mg–Fe partition coefficient between Rw and Mw is defined by: K

Rw–Mw

Rw

¼ ðXFe =XMg Þ

Mw

=ðXFe =XMg Þ

ð2Þ

Fig. 3. Phase relations in Mg2SiO4. Open triangle, ringwoodite(Rw); half-closed triangle, Rw + MgSiO3–perovskite(Mpv) + periclase(Pc); closed triangle, Mpv + Pc; open inverse triangle, wadsleyite(Wd); half-closed inverse triangle, Wd + Mpv + Pc. Solid lines represent phase boundaries in Mg2SiO4.

The K Rw–Mw values are 0.57–0.66 at 21.1–22.3 GPa and 1800 °C (nos. 78 and 23). The data are generally consistent with the K Rw–Mw of 0.61–0.65 at 21 GPa and 1600 °C by the Mg–Fe partition experiments in the Mg-rich compositions (Frost et al., 2001). The volume fractions of minerals in pyrolite were obtained by mass balance calculation using the mineral compositions analyzed. The following zero-pressure volumes in Fabrichnaya et al. (2004)

Fig. 4. Phase relations in MgSiO3. Open square, akimotoite(Ak); half-closed square, Ak+ Mpv; closed square, Mpv; open diamond, MgSiO3–garnet(Mj); half-closed diamond, Ak+ Mj. Solid lines represent phase boundaries in MgSiO3, and dashed lines the phase boundaries in Mg2SiO4 shown in Fig. 3.

Fig. 5. Phase relations in pyrolite. Open circle, Rw + Gt + Cpv; half-closed circle, Rw + Mpv + Mw + Gt + Cpv; closed circle, Mpv + Mw + Gt + Cpv; double triangle, Mpv + Mw + Cpv; double square, Rw + Ak + Gt + Cpv; double circle, Rw + Mw + Gt + Cpv. Rw, ringwoodite; Mpv, MgSiO3-rich perovskite; Mw, magnesiowüstite; Gt, majorite garnet; Cpv, CaSiO3-rich perovskite; Ak, akimotoite. Solid lines represent the phase boundaries in pyrolite, and dashed lines the boundaries in Mg2SiO4 shown in Fig. 3.

190

Table 2 Compositions of coexisting minerals in run products of pyrolite. Run no. 32 22.9 GPa, 1400 °C, 6 h

SiO2 TiO2

Cr2O3 FeO MgO NiO CaO Na2O Total O Si Ti Al Cr Fe Mg Ni Ca Na C. T.

Run no. 44 24.1 GPa, 1500 °C, 3 h

Run no. 9 24.1 GPa, 1600 °C, 3 h

Run no. 64 25.3 GPa, 1600 °C, 3 h

Ak

Rw

Gt

Cpv

Rw

Gt

Mpv

Mw

Cpv

Mpv

Mw

Gt

Cpv

Mpv

Mw

Gt

Cpv

Mpv

57.51 (2.31) 0.24 (0.20) 0.93 (0.59) 0.11 (0.10) 3.11 (0.97) 38.53 (1.02) 0.03 (0.07) 0.11 (0.14) 0.04 (0.06) 100.61 3 0.97 0.00 0.02 0.00 0.04 0.97 0.00 0.00 0.00 2.00

41.10 (0.74) 0.09 (0.09) 0.91 (0.68) 0.17 (0.18) 10.19 (0.36) 46.29 (1.22) 0.14 (0.12) 0.23 (0.12) 0.11 (0.10) 99.23 4 1.02 0.00 0.03 0.00 0.21 1.70 0.00 0.01 0.01 2.98

47.42 (1.83) 0.04 (0.07) 13.53 (2.73) 0.18 (0.13) 6.73 (1.85) 26.06 (0.78) 0.06 (0.11) 4.53 (1.91) 0.77 (0.10) 99.32 12 3.37 0.00 1.13 0.01 0.40 2.76 0.00 0.35 0.11 8.13

64.39 (5.02) 0.18 (0.19) 0.52 (0.28) 0.07 (0.57) 0.57 (0.20) 2.98 (1.09) 0.07 (0.09) 32.10 (6.15) 0.23 (0.09) 101.11 3 1.14 0.00 0.01 0.00 0.01 0.08 0.00 0.61 0.01 1.86

41.37 (0.47) 0.01 (0.01) 0.04 (0.05) 0.10 (0.04) 9.03 (0.53) 48.52 (0.50) 0.43 (0.16) 0.09 (0.09) 0.00 (0.00) 99.59 4 1.01 0.00 0.00 0.00 0.18 1.77 0.01 0.00 0.00 2.97

47.33 (0.64) 0.16 (0.11) 18.52 (1.84) 0.23 (0.18) 3.36 (1.70) 26.23 (1.13) 0.78 (1.51) 2.62 (0.30) 0.78 (0.34) 100.01 12 3.26 0.01 1.50 0.01 0.19 2.70 0.04 0.20 0.11 8.02

57.67 (0.93) 0.07 (0.09) 1.18 (0.45) 0.05 (0.09) 2.30 (0.51) 37.87 (0.86) 0.07 (0.10) 0.15 (0.10) 0.02 (0.04) 99.38 3 0.98 0.00 0.03 0.00 0.03 0.96 0.00 0.00 0.00 2.00

0.93 (1.02) 0.01 (0.01) 0.72 (0.77) 1.18 (0.59) 24.93 (1.92) 70.99 (3.64) 0.75 (0.18) 0.31 (0.49) 0.66 (0.34) 100.48 1 0.01 0.00 0.01 0.01 0.16 0.80 0.00 0.00 0.01 1.00

58.33 (2.32) 0.29 (0.16) 0.40 (0.19) 0.21 (0.25) 0.48 (0.15) 1.98 (1.10) 0.06 (0.08) 37.74 (1.59) 0.45 (0.09) 99.94 3 1.08 0.00 0.01 0.00 0.01 0.05 0.00 0.75 0.01 1.91

53.49 (1.48) 0.23 (0.10) 4.23 (0.48) 0.37 (0.12) 6.39 (0.36) 35.03 (0.79) 0.17 (0.17) 0.51 (0.43) 0.07 (0.10) 100.49 3 0.92 0.00 0.09 0.00 0.09 0.90 0.00 0.01 0.00 2.01

2.12 (0.57) 0.06 (0.11) 0.71 (0.19) 0.67 (0.09) 16.95 (0.34) 76.84 (0.85) 1.48 (0.28) 0.10 (0.07) 0.86 (0.16) 99.79 1 0.01 0.00 0.01 0.00 0.10 0.83 0.01 0.00 0.01 0.97

48.21 (1.45) 0.06 (0.09) 15.56 (2.37) 0.65 (0.16) 3.48 (0.30) 27.57 (1.17) 0.14 (0.19) 3.59 (0.38) 0.58 (0.20) 99.84 12 3.34 0.00 1.27 0.04 0.20 2.85 0.01 0.27 0.08 8.06

59.10 (2.16) 0.18 (0.10) 0.62 (0.31) 0.23 (0.22) 0.58 (0.46) 2.60 (1.82) 0.06 (0.05) 36.19 (3.37) 0.96 (0.44) 100.52 3 1.08 0.00 0.01 0.00 0.01 0.07 0.00 0.71 0.04 1.92

54.19 (0.71) 0.20 (0.10) 4.36 (0.21) 0.22 (0.11) 5.50 (0.19) 35.60 (0.41) 0.04 (0.04) 0.36 (0.03) 0.05 (0.05) 100.52 3 0.93 0.00 0.09 0.00 0.08 0.91 0.00 0.01 0.00 2.02

2.27 (0.97) 0.09 (0.06) 1.40 (0.19) 0.78 (0.13) 17.30 (0.60) 76.37 (0.96) 1.12 (0.26) 0.21 (0.06) 0.92 (0.08) 100.46 1 0.02 0.00 0.01 0.00 0.11 0.82 0.01 0.00 0.01 0.98

49.85 (0.52) 0.08 (0.05) 15.63 (0.86) 0.31 (0.12) 3.13 (0.22) 26.92 (0.44) 0.06 (0.05) 3.08 (0.17) 1.01 (0.08) 100.07 12 3.42 0.00 1.27 0.02 0.18 2.75 0.00 0.23 0.14 8.00

70.43 (2.06) 0.15 (0.07) 0.99 (0.66) 0.09 (0.15) 0.43 (0.15) 1.82 (0.87) 0.15 (0.12) 25.94 (1.73) 0.33 (0.11) 100.33 3 1.21 0.00 0.02 0.00 0.01 0.05 0.00 0.48 0.01 1.78

51.35(0.48)

Run no. 64 25.3 GPa, Run no. 61 26.7 GPa, 1600 °C, 3 h 1600 °C, 3 h Mw

Cpv

Mpv

Mw

Run no. 29 22.9 GPa, 1700 °C, 6 h Cpv

Rw

Gt

Run no. 43 22.9 GPa, 1700 °C, 2 h Mw

Rw

Gt

Run no. 46 23.8 GPa, 1700 °C, 2 h Mw

Cpv

Mpv

Mw

0.20(0.12) T. Ishii et al. / Earth and Planetary Science Letters 309 (2011) 185–197

Al2O3

Run no. 31 23.3 GPa, 1500 °C, 3 h

6.09(0.18) 0.26(0.12) 6.45(0.29) 35.39(0.41) 0.12(0.11) 0.29(0.08) 0.07(0.05) 100.22 3 0.89 0.00 0.13 0.00 0.09 0.92 0.00 0.01 0.00 2.04

Run no. 51 23.8 GPa, 1700 °C, 6 h Gt

Cpv

Mpv

Mw

Gt

Cpv

Table 2 (continued) Run no. 64 25.3 GPa, Run no. 61 26.7 GPa, 1600 °C, 3 h 1600 °C, 3 h

SiO2 TiO2

Cr2O3 FeO MgO NiO CaO Na2O Total O Si Ti Al Cr Fe Mg Ni Ca Na C. T.

Run no. 46 23.8 GPa, 1700 °C, 2 h

Run no. 51 23.8 GPa, 1700 °C, 6 h

Mw

Cpv

Mpv

Mw

Cpv

Rw

Gt

Mw

Rw

Gt

Mw

Cpv

Mpv

Mw

Gt

Cpv

Mpv

Mw

Gt

Cpv

1.34 (0.51) 0.12 (0.08) 1.80 (0.17) 1.02 (0.13) 17.77 (0.33) 75.35 (1.08) 1.05 (0.22) 0.15 (0.07) 1.47 (0.07) 100.07 1 0.01 0.00 0.02 0.01 0.11 0.82 0.01 0.00 0.02 1.00

53.38 (1.76) 0.10 (0.20) 0.81 (0.33) 0.12 (0.12) 0.63 (0.30) 1.47 (0.58) 0.11 (0.20) 43.15 (1.74) 0.31 (0.23) 100.08 3 1.01 0.00 0.02 0.00 0.01 0.04 0.00 0.88 0.01 1.97

52.72 (1.35) 0.28 (0.17) 5.75 (0.39) 0.29 (0.13) 5.98 (0.59) 33.97 (0.52) 0.07 (0.11) 0.32 (0.14) 0.05 (0.09) 99.43 3 0.92 0.00 0.12 0.00 0.09 0.88 0.00 0.00 0.00 2.01

1.31 (0.62) 0.19 (0.16) 2.54 (0.30) 0.70 (0.17) 17.46 (0.11) 74.61 (0.68) 1.12 (0.30) 0.08 (0.08) 1.46 (0.16) 99.47 1 0.01 0.00 0.02 0.00 0.11 0.82 0.01 0.00 0.02 0.99

57.46 (0.56) 0.01 (0.02) 0.91 (1.02) 0.05 (0.09) 0.64 (0.29) 1.00 (0.45) 0.02 (0.03) 39.52 (1.53) 0.26 (0.17) 99.87 3 1.07 0.00 0.02 0.00 0.01 0.03 0.00 0.79 0.01 1.93

39.70 (0.36) 0.13 (0.12) 0.36 (0.15) 0.16 (0.12) 9.91 (0.36) 49.30 (0.53) 0.57 (0.43) 0.06 (0.08) 0.10 (0.11) 100.29 4 0.98 0.00 0.01 0.00 0.20 1.81 0.01 0.00 0.00 3.01

50.24 (0.40) 0.26 (0.10) 10.33 (0.38) 1.23 (0.22) 4.41 (0.25) 28.34 (0.34) 0.10 (0.11) 4.39 (0.21) 0.38 (0.06) 99.68 12 3.51 0.01 0.85 0.07 0.26 2.95 0.01 0.33 0.05 8.04

0.69 (0.13) 0.17 (0.20) 0.59 (0.22) 0.72 (0.25) 22.25 (0.48) 73.42 (0.54) 1.14 (0.32) 0.13 (0.09) 0.43 (0.15) 99.54 1 0.00 0.00 0.01 0.00 0.14 0.82 0.01 0.00 0.01 0.99

42.50 (0.71) 0.12 (0.04) 0.75 (0.42) 0.05 (0.08) 7.67 (0.94) 47.81 (1.18) 0.04 (0.08) 0.17 (0.14) 0.15 (0.20) 99.26 4 1.04 0.00 0.02 0.00 0.16 1.73 0.00 0.01 0.01 2.97

52.35 (1.92) 0.17 (0.16) 9.53 (3.24) 0.42 (0.12) 3.46 (0.18) 29.12 (1.68) 0.08 (0.17) 3.50 (0.50) 0.50 (0.10) 99.13 12 3.63 0.01 0.78 0.02 0.20 3.01 0.00 0.26 0.07 7.98

1.04 (1.48) 0.28 (0.17) 0.40 (0.22) 0.32 (0.10) 20.66 (2.73) 75.38 (3.34) 0.71 (0.27) 0.14 (0.05) 0.57 (0.16) 99.50 1 0.01 0.00 0.00 0.00 0.13 0.84 0.00 0.00 0.01 0.99

59.94 (3.29) 0.30 (0.24) 0.60 (0.17) 0.11 (0.13) 0.60 (0.21) 2.36 (0.61) 0.20 (0.23) 34.36 (3.30) 1.17 (0.31) 99.64 3 1.10 0.00 0.01 0.00 0.01 0.07 0.00 0.68 0.04 1.91

52.53 (0.53) 0.24 (0.11) 3.45 (0.30) 0.27 (0.12) 5.99 (0.22) 35.64 (0.70) 0.09 (0.11) 0.51 (0.14) 0.10 (0.08) 98.82 3 0.92 0.00 0.07 0.00 0.09 0.93 0.00 0.01 0.00 2.02

0.60 (0.31) 0.47 (0.24) 0.54 (0.13) 0.13 (0.12) 21.57 (0.30) 73.98 (1.50) 1.27 (0.36) 0.13 (0.06) 0.57 (0.16) 99.26 1 0.00 0.00 0.00 0.00 0.14 0.83 0.01 0.00 0.01 0.99

50.71 (0.53) 0.73 (0.19) 10.54 (0.49) 0.12 (0.09) 4.50 (0.25) 28.98 (0.46) 0.25 (0.23) 4.09 (0.32) 0.53 (0.09) 100.45 12 3.52 0.04 0.86 0.01 0.26 2.99 0.01 0.30 0.07 8.06

56.32 (1.02) 0.11 (0.14) 0.34 (0.19) 0.06 (0.09) 0.49 (0.24) 1.18 (0.31) 0.09 (0.13) 40.42 (1.00) 0.21 (0.07) 99.22 3 1.06 0.00 0.01 0.00 0.01 0.03 0.00 0.82 0.01 1.94

51.83 (1.28) 0.44 (0.15) 5.16 (0.46) 0.43 (0.14) 6.20 (0.58) 35.48 (0.76) 0.26 (0.26) 0.24 (0.05) 0.04 (0.04) 100.08 3 0.90 0.01 0.11 0.01 0.09 0.92 0.00 0.00 0.00 2.04

0.91 (0.30) 0.08 (0.17) 1.10 (0.27) 1.07 (0.19) 21.30 (0.94) 73.01 (0.35) 1.31 (0.44) 0.06 (0.06) 0.86 (0.19) 99.70 1 0.01 0.00 0.01 0.01 0.13 0.82 0.01 0.00 0.01 1.00

50.37 (0.82) 0.26 (0.12) 10.73 (0.88) 0.53 (0.16) 4.51 (0.65) 31.55 (0.64) 0.20 (0.22) 1.66 (0.13) 0.42 (0.14) 100.23 12 3.48 0.01 0.87 0.03 0.26 3.25 0.01 0.12 0.06 8.09

55.71 (0.98) 0.18 (0.20) 0.84 (0.18) 0.15 (0.15) 0.53 (0.32) 1.61 (0.19) 0.00 (0.00) 41.85 (0.75) 0.20 (0.20) 101.07 3 1.04 0.00 0.02 0.00 0.01 0.04 0.00 0.84 0.01 1.96

Run no.78 21.1 GPa, 1800 °C, 2 h

SiO2

Run no. 43 22.9 GPa, 1700 °C, 2 h

Run no. 23 22.3 GPa, 1800 °C, 2 h

Run no. 41 23.8 GPa, 1800 °C, 2 h

Run no. 73 24.7 GPa, 1800 °C, 2 h

Run no. 68 26.2 GPa, 1800 °C, 2 h

T. Ishii et al. / Earth and Planetary Science Letters 309 (2011) 185–197

Al2O3

Run no. 29 22.9 GPa, 1700 °C, 6 h

Run no. 60 28.0 GPa, 1800 °C, 2 h

Rw

Mw

Gt

Rw

Mw

Gt

Mpv

Mw

Gt

Mpv

Mw

Gt

Cpv

Mpv

Mw

Cpv

Mpv

Mw

Cpv

40.60

1.30

50.92

39.63

0.44

49.64

55.01

1.55

51.58

53.34

1.58

47.18

56.94

51.58

1.55

60.00

50.78

5.64

57.15

191

192

Table 2 (continued) Run no.78 21.1 GPa, 1800 °C, 2 h

TiO2

Cr2O3 FeO MgO NiO CaO Na2O Total O Si Ti Al Cr Fe Mg Ni Ca Na C. T.

Run no. 41 23.8 GPa, 1800 °C, 2 h

Run no. 73 24.7 GPa, 1800 °C, 2 h

Run no. 68 26.2 GPa, 1800 °C, 2 h

Run no. 60 28.0 GPa, 1800 °C, 2 h

Rw

Mw

Gt

Rw

Mw

Gt

Mpv

Mw

Gt

Mpv

Mw

Gt

Cpv

Mpv

Mw

Cpv

Mpv

Mw

Cpv

(0.44) 0.14 (0.09) 0.20 (0.17) 0.09 (0.08) 9.27 (0.22) 48.78 (0.38) 0.48 (0.17) 0.15 (0.13) 0.04 (0.03) 99.75 4 1.00 0.00 0.01 0.00 0.19 1.79 0.00 0.00 0.00 2.99

(0.89) 0.17 (0.14) 0.57 (0.31) 0.40 (0.10) 23.15 (0.32) 73.13 (0.74) 0.86 (0.28) 0.22 (0.14) 0.51 (0.18) 100.31 1 0.01 0.00 0.00 0.00 0.15 0.81 0.01 0.00 0.01 0.99

(0.44) 0.28 (0.10) 9.56 (0.40) 0.58 (0.08) 5.17 (0.17) 28.16 (0.27) 0.10 (0.09) 5.49 (0.23) 0.41 (0.08) 100.67 12 3.54 0.01 0.78 0.03 0.30 2.92 0.00 0.41 0.05 8.04

(0.30) 0.14 (0.14) 0.32 (0.13) 0.23 (0.24) 9.72 (0.49) 49.23 (0.50) 0.62 (0.21) 0.12 (0.09) 0.13 (0.12) 100.14 4 0.97 0.00 0.01 0.01 0.20 1.81 0.02 0.00 0.01 3.03

(0.21) 0.18 (0.08) 0.43 (0.20) 0.55 (0.31) 22.12 (0.45) 75.54 (1.14) 0.50 (0.29) 0.15 (0.12) 0.34 (0.12) 100.25 1 0.00 0.00 0.00 0.00 0.14 0.84 0.00 0.00 0.00 0.98

(0.58) 0.36 (0.10) 11.03 (0.58) 0.59 (0.15) 4.54 (0.29) 27.77 (0.50) 0.09 (0.08) 5.41 (0.30) 0.50 (0.05) 99.93 12 3.47 0.02 0.91 0.03 0.27 2.90 0.01 0.41 0.07 8.09

(0.85) 0.29 (0.15) 4.08 (0.23) 0.31 (0.23) 5.38 (0.23) 35.94 (0.69) 0.05 (0.09) 0.34 (0.04) 0.07 (0.12) 101.47 3 0.94 0.00 0.08 0.01 0.08 0.91 0.00 0.01 0.00 2.03

(1.29) 0.14 (0.12) 2.97 (0.47) 1.21 (0.47) 16.2 (1.71) 73.96 (2.07) 0.85 (0.41) 0.36 (0.36) 1.88 (0.09) 99.12 1 0.01 0.00 0.03 0.01 0.10 0.81 0.01 0.00 0.03 1.00

(0.80) 0.11 (0.05) 10.75 (1.23) 0.56 (0.13) 1.34 (0.11) 32.93 (0.67) 0.07 (0.07) 2.95 (0.23) 0.16 (0.07) 100.45 12 3.51 0.01 0.86 0.03 0.08 3.34 0.00 0.22 0.02 8.07

(0.17) 0.39 (0.18) 4.40 (0.59) 0.39 (0.27) 4.43 (0.48) 36.52 (0.46) 0.04 (0.06) 0.33 (0.13) 0.07 (0.16) 99.91 3 0.92 0.01 0.09 0.00 0.07 0.94 0.00 0.01 0.00 2.04

(1.61) 0.15 (0.14) 3.01 (0.57) 1.04 (0.35) 16.58 (1.87) 73.74 (2.52) 0.94 (0.44) 0.38 (0.44) 1.89 (0.11) 99.31 1 0.01 0.00 0.03 0.01 0.10 0.80 0.01 0.00 0.03 0.99

(0.74) 0.07 (0.07) 17.84 (1.11) 0.59 (0.11) 2.70 (0.46) 28.84 (0.44) 0.10 (0.12) 2.87 (0.13) 0.28 (0.10) 100.47 12 3.23 0.00 1.44 0.03 0.15 2.94 0.01 0.21 0.04 8.05

(1.18) 0.02 (0.05) 0.34 (0.44) 0.09 (0.10) 0.28 (0.05) 2.94 (1.39) 0.24 (0.19) 39.38 (2.22) 0.30 (0.21) 100.53 3 1.06 0.00 0.01 0.00 0.00 0.08 0.00 0.78 0.01 1.94

(0.52) 0.23 (0.14) 6.36 (0.31) 0.27 (0.10) 6.47 (0.54) 35.09 (0.49) 0.19 (0.13) 0.46 (0.09) 0.02 (0.02) 100.67 3 0.89 0.00 0.13 0.00 0.10 0.91 0.00 0.01 0.00 2.04

(0.96) 0.14 (0.09) 2.97 (0.35) 1.21 (0.35) 16.20 (1.27) 73.96 (1.54) 0.85 (0.30) 0.36 (0.27) 1.88 (0.07) 99.12 1 0.01 0.00 0.03 0.01 0.10 0.81 0.01 0.00 0.03 1.00

(1.67) 0.08 (0.05) 0.69 (0.07) 0.14 (0.11) 0.55 (0.07) 2.57 (0.83) 0.09 (0.08) 36.22 (2.40) 0.38 (0.05) 100.72 3 1.09 0.00 0.02 0.00 0.01 0.07 0.00 0.71 0.01 1.91

(0.53) 0.24 (0.10) 6.22 (0.27) 0.29 (0.11) 6.70 (0.22) 35.22 (0.46) 0.14 (0.10) 0.38 (0.11) 0.03 (0.02) 100.00 3 0.89 0.00 0.13 0.00 0.10 0.92 0.00 0.01 0.00 2.05

(2.45) 0.08 (0.06) 2.71 (0.25) 0.89 (0.11) 14.25 (0.65) 72.85 (2.01) 1.56 (0.24) 0.35 (0.31) 1.76 (0.11) 100.09 1 0.04 0.00 0.02 0.01 0.08 0.77 0.01 0.00 0.02 0.95

(2.51) 0.10 (0.12) 0.85 (0.36) 0.05 (0.10) 0.33 (0.31) 2.79 (1.12) 0.00 (0.00) 38.24 (1.85) 0.37 (0.18) 99.88 3 1.06 0.00 0.02 0.00 0.00 0.08 0.00 0.76 0.02 1.94

Abbreviation: Rw, ringwoodite; Ak, akimotoite; Mpv, MgSiO3-rich perovskite; Mw, magnesiowüstite; Gt, majorite garnet; Cpv, CaSiO3-rich perovskite; C. T., cation total. Total iron as FeO. Numbers in parentheses indicate two standard deviations of the mean in the analyses.

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Al2O3

Run no. 23 22.3 GPa, 1800 °C, 2 h

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Fig. 6. Back-scattered electron images of typical recovered run products of pyrolite. (a) Run no.33 (22.9 GPa, 1500 °C). (b) Run no.31 (23.3 GPa, 1500 °C). (c) Run no.9 (24.1 GPa, 1600 °C). (d) Run no.62 (26.2 GPa, 1600 °C). (e) Run no.32 (22.9 GPa, 1400 °C). (f) Run no.18 (22.9 GPa, 1800 °C).

were used for the calculations of volume fractions of minerals, where numbers in parentheses are volumes in cm 3/mol: periclase (11.25), wüstite (12.25), Mg2SiO4 spinel (39.65), Fe2SiO4 spinel (42.02), MgSiO3 garnet (28.50), FeSiO3 garnet (29.43), almandine (115.28), pyrope (113.28), MgSiO3 ilmenite (26.35), FeSiO3 ilmenite (27.60), MgSiO3 perovskite (24.45), and FeSiO3 perovskite (25.59), as well as grossular (125.31) of Saxena et al. (1993), Al2O3 perovskite (25.94) of Hemley and Cohen (1992), and CaSiO3 perovskite (27.32) of Mao et al. (1989). Fig. 9(a–c) shows the volume fractions of minerals in pyrolite at 1400–1800 °C with increase of pressure. These calculations were made in a pressure interval of about 1 GPa at each temperature. Fig. 9 (d) shows the volume fractions of minerals in pyrolite at 22.9 GPa with increasing temperature. This figure indicates that Rw dissociates partially into Gt + Mw at 22.9 GPa above about 1650 °C. Here, we compare the post-spinel phase relations between pyrolite and Mg2SiO4 in details. The following differences are observed in Fig. 5. (1) The post-spinel transition in pyrolite occurs at pressure interval within about 0.4 GPa, compared with the univariant transition boundary of Mg2SiO4. (2) The post-spinel transition in pyrolite occurs by about 0.2–1.0 GPa lower than that in Mg2SiO4 at 1400–1800 °C. (3) The Clapeyron slope of the post-spinel transition in pyrolite is about −0.001 GPa/°C, which is half of the slope of Mg2SiO4. The uncertainty of the slope of pyrolite was estimated as ±0.0005 GPa/°C. These results are compared with those of the previous studies and discussed below.

It should be noted that the above transition pressures and slopes depend on the transition boundaries of Mg2SiO4, MgSiO3, MgAl2O4, and Mg3Al2Si3O12 adopted for the pressure calibration in this study. These boundaries are based on different pressure standards: the Wd–Rw transition (Suzuki et al., 2000) on Brown's (1999) NaCl scale, the MgAl2O4 transition (Irifune et al., 2002) on Anderson et al.'s (1989) Au scale, and the transitions of Ak–MPv and pyrope (Fei et al., 2004b) on Speziale et al.'s (2001) MgO scale. The Brown's NaCl and Speziale et al.'s MgO scales are generally consistent within about 0.5 GPa at about 20 GPa at high temperature (Fei et al., 2004a). However, several studies indicated that Anderson et al.'s Au scale gives lower pressure than that of Speziale et al.'s MgO scale by about 1 GPa (Fei et al., 2004a; Matsui, 2010; Matsui and Nishiyama, 2002). Therefore, it is highly desirable in future studies to establish the transition boundaries at high pressures and high temperatures by a unified pressure calibration standard with the accurate equation of state, in order to determine accurately the pressure and slope of the post-spinel transition boundary. However, even if we tentatively increase the pressure of Pc + Cor to MgAl2O4 calcium ferrite by 1 GPa in the calibration curve in Fig. 1, leaving the other calibration points unchanged, the slopes of the post-spinel transition boundaries in Mg2SiO4 and in pyrolite slightly change to − 0.0024 and − 0.0011 GPa/°C, respectively, and that the pressure difference in the post-spinel transition between pyrolite and Mg2SiO4 is about 0.2–1.0 GPa. These results indicate that the slope

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Fig. 7. X-ray diffraction profiles of pyrolite. (a) X-ray diffraction profiles at 1600 °C with increase of pressure (22.9, 23.3, 23.8, and 27.6 GPa from bottom to top). Phase assemblages of run nos. 16, 17, 38 and 57 are Rw+ Gt+ Cpv, Rw+ Mpv + Mw+ Gt+ Cpv, Mpv + Mw + Gt+ Cpv and Mpv+ Mw+ Cpv, respectively. (b) Run no. 32 (22.9 GPa, 1400 °C) and no. 26 (23.8 GPa, 1400 °C). Phase assemblages are Rw + Gt+ Ak+ Cpv and Mpv + Mw + Gt+ Cpv, respectively. (c) Run no. 43 (22.9 GPa, 1700 °C). Phase assemblage is Rw+ Mw+ Gt+ Cpv. Rw, ringwoodite; Mpv, MgSiO3-rich perovskite; Mw, magnesiowüstite; Gt, majorite garnet; Cpv, CaSiO3-rich perovskite; Ak, akimotoite; Re, Re capsule.

values depend a little on the pressure calibration points but that the relative differences in the slopes and transition pressures are hardly affected.

The transition pressure interval within about 0.4 GPa and the lower pressure by about 0.2–1.0 GPa in pyrolite than in Mg2SiO4 can be attributed to that pyrolite is a multi-component system, compared

Fig. 8. Changes of mineral compositions (cation numbers) with pressure at 1800 °C. Gt, majorite garnet; Mpv, MgSiO3-rich perovskite; Rw, ringwoodite; Mw, magnesiowüstite; Triangle, Ca; inverse triangle, Mg; diamond, Si; square, Fe; circle, Al. Errors of cation numbers are close to or less than sizes of the symbols.

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Fig. 9. Mineral proportions in pyrolite with increase of pressure at (a)1400 °C, (b)1600 °C, and (c)1800 °C. (d) Mineral proportions in pyrolite at 22.9 GPa with increase of temperature. Closed squares represent the calculated mineral fractions. Rw, ringwoodite; Mpv, MgSiO3-rich perovskite; Mw, magnesiowüstite; Gt, majorite garnet; Cpv, CaSiO3-rich perovskite; Ak, akimotoite.

with pure Mg2SiO4. This implies that minor components such as Fe 2+, Fe 3+ and Al 3+ affect the post-spinel phase relations. Ito and Takahashi's (1989) experimental results indicate that the effect of Fe 2+ on the post-spinel transition pressure of Mg2SiO4 is negligibly small. However, thermodynamic calculation by Fabrichnaya et al. (2004) indicates that Fe 2+ increases slightly the transition pressure. Incorporation of Al 3+ and Fe 3+ in Mg–perovskite (Frost and McCammon, 2008) would affect the post-spinel transition pressure and interval, which has not yet been examined. Therefore, the effects of the minor elements should be examined in detail as a future investigation. Considering the uncertainty of pressure, the transition pressure interval in this study is consistent with those determined with in situ X-ray diffraction studies by Nishiyama et al. (2004) and Litasov et al. (2005a). The lower pressure by about 0.2–1.0 GPa for the post-spinel transition in pyrolite than that in Mg2SiO4 is also almost consistent with the result by Litasov et al. (2005a). Our results clearly indicate that the post-spinel transition in pyrolite has a smaller boundary slope than that of Mg2SiO4. At 23.3 GPa and 1700 °C (run no. 53), Rw + Mpv + Mw (or periclase, Pc) were observed in both Mg2SiO4 and pyrolite. However, at 23.8 GPa and 1500 °C (no. 48), Rw + Mpv + Pc were observed in Mg2SiO4, while all of Rw transformed to Mpv + Mw in pyrolite at 23.8 GPa and 1400– 1500 °C (nos. 26, 48 and 52) (Fig. 7b). As the results, the post-spinel transition boundary in pyrolite has a slope of about −0.001 GPa/°C, which is half of the slope of Mg2SiO4. Litasov et al. (2005a) compared the post-spinel transition slopes in pyrolite and Mg2SiO4 in a number of previous studies, and concluded that the compositional difference between pyrolite and Mg2SiO4 would not change the slope signifi-

cantly. However, the previous studies did not directly compare the transitions precisely between pyrolite and Mg2SiO4 in the simultaneous high-pressure high-temperature experiments. The results on the different Clapeyron slopes of pyrolite and Mg2SiO4 suggest the following issues. Here, we assume that both the transition zone and the uppermost lower mantle have approximately pyrolitic composition. To evaluate effect of the post-spinel transition slope on mantle dynamics and to estimate lateral temperature variations based on elevation and depression of the 660-km discontinuity, the postspinel transition slope of pyrolite rather than Mg2SiO4 should be used, although the slope of Mg2SiO4 has been generally used in the previous studies (e.g., Davies, 1998; Fukao et al., 2009; Helffrich, 2000). Even if the composition of the transition zone is a little different from pyrolite, our results suggest that the minor components may change the slope of the post-spinel transition boundary from that of Mg2SiO4. The above Clapeyron slope of the post-spinel transition in pyrolite, –0.001 GPa/°C, is smaller than the usually assumed slope range, –0.002––0.003 GPa/°C, in geodynamics studies (e.g., Davies, 1998). In addition, Fukao et al. (2009) seismologically estimated the Clapeyron slope at the 660 km discontinuity to be about −0.0025 GPa/°C. The difference between Fukao et al.'s and ours may be derived at least in part from the uncertainty of the seismological estimate, and also from the errors of the pressure calibration, as described above. As another possibility, effect of water which would allow the post-spinel transition slope more negative may be considered (Litasov et al., 2005b). In Hirose's (2002) experiments of natural peridotite KLB-1, a very small amount of akimotoite (Ak) was observed at 22 GPa and 1600 °C, and the Ak completely transformed to Mpv at pressure about 1 GPa

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lower than the post-spinel transition. However, Fig. 5 indicates that in our experiments the Ak–Mpv transition and the post-spinel transition occur in pyrolite at the same pressure within the pressure uncertainty at 1400 °C. As discussed by Nishiyama and Yagi (2003) and Nishiyama et al. (2004), the different pressures of the Ak–Mpv transition and the post-spinel transition in Hirose's (2002) experiments might result from non-equilibrium run products, particularly in Al2O3 contents in Gt and MPv, in his experiments in which the starting material of the pulverized rock was used. In this study, Gt is stable in pyrolite at pressure up to 24.5–26 GPa at 1400–1800 °C (Fig. 5). The pressure is a little smaller than 26.5 GPa at 1600 °C for dissociation of pyrope. Hirose (2002) reported that Gt is stable in the peridotite KLB-1 at pressure only to 24 GPa at 1400–1800 °C. The pressure difference between Hirose's and our studies may be attributed to difference of Al2O3 + Cr2O3 contents between pyrolite (4.83 wt.%) and KLB-1 (3.90 wt.%), because the larger amount of Al2O3 + Cr2O3 stabilizes Gt relative to Mpv at higher pressure. Recent seismological studies indicate that detailed structure of the 660-km discontinuity is complicated and double seismic reflections have been observed at the depth ranges of about 660 km and around 720 km (e.g., Deuss et al., 2006). The depths of the double discontinuities generally correspond to the post-spinel and Gt–Mpv transitions, as shown in Figs. 5 and 9. Fig. 9(a–c) indicates that abundance of Gt coexisting with MPv increases with increasing temperature above about 23 GPa. This is generally consistent with the results by Hirose (2002) and Nishiyama and Yagi (2003). This may be explained, based on the garnet–perovskite transition loop in the simple system Mg4Si4O12–Mg3Al2Si3O12 (Kubo and Akaogi, 2000; Hirose et al., 2001a). In the system, the transitions of Gt → Gt + Mpv→ Mpv have positive dP/dT slopes, which indicate that abundance of Gt relative to MPv decreases with pressure and increases with temperature. Fig. 9(d) demonstrates that, at 22.9 GPa and about 1600–1700 °C, abundance of Rw decreases, that of Gt increases, and Mw becomes stable. Fig. 5 shows that above about 1650 °C Rw + Mw + Gt + Cpv are the stable phases at 21.1–22.9 GPa which is below the post-spinel transition pressure. Therefore, the increase of Gt abundance with increasing temperature occurs in not only the Mpv-stable field but also the Rw-stable one. Although we observed the increase of Gt with increasing temperature, the increase is not so remarkable at temperature below 1700–1800 °C as Gt is more abundant than Rw. Therefore, compared with the post-spinel transition, the Gt–Mpv transition is not the dominant phase transition below 1700–1800 °C at around the 660-km depth. From seismological studies, Houser and Williams (2010) found that the Clapeyron slope associated with the 660-km discontinuity may change the sign from negative to positive with increasing temperature in relatively warm mantle regions. They suggest that the dominant phase transition associated with the 660-km discontinuity shifts from the post-spinel to the Gt–Mpv. To clarify the effect of the Gt–Mpv transition at the 660-km discontinuity, further study is necessary to examine the detailed phase relations at temperature range above about 1700–1800 °C. Acknowledgments We are grateful to E. Ito and M. Matsui for valuable discussion. We express our thanks to anonymous referees for their constructive reviews and to L. Stixrude for editorial handling. This research was supported in part by the Grant-in-Aid (No. 22340163) of the Scientific Research of the Japan Society for the Promotion of Science to M. Akaogi. References Akaogi, M., Ito, E., 1993. Refinement of enthalpy measurement of MgSiO3 perovskite and negative pressure–temperature slopes for perovskite-forming reactions. Geophys. Res. Lett. 20, 1839–1842.

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