A speculation on the structure of the D″ layer: The growth of anti-crust at the core–mantle boundary through the subduction history of the Earth

Gondwana Research 15 (2009) 342–353

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A speculation on the structure of the D″ layer: The growth of anti-crust at the core–mantle boundary through the subduction history of the Earth Tetsuya Komabayashi ⁎, Shigenori Maruyama, Shuji Rino Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Meguro, Tokyo 152-8551, Japan

a r t i c l e

i n f o

Article history: Received 17 July 2008 Received in revised form 17 November 2008 Accepted 20 November 2008 Available online 3 December 2008 Keywords: D″ layer Core–mantle boundary Seismic discontinuity Anti-crust Continental subduction

a b s t r a c t The growth curve of the continental crust shows that large amounts of continental crust formed in the early part of the Earth history are missing. In order to test a hypothesis that the former crust was subducted to the deep mantle, we performed phase assemblage analysis in the systems of mid-oceanic ridge basalt (MORB), anorthosite, and tonalite–trondhjemite–granite (TTG) down to the core–mantle boundary (CMB) conditions. Results show that all these materials can be subducted to the CMB leading to the development of a compositional layering in the D″ layer. We speculate that there could be five layers of FeO-enriched melt from partial melting of MORB, MORB crust, anorthosite, TTG, and slab or mantle peridotite in ascending order. Although the polymorphic transformation of perovskite to post-perovskite in (Mg,Fe)SiO3 may explain the seismic discontinuity at the top of the D″ layer (D″ discontinuity), the effects of solid solution on the sharpness of the transformation suggest that the compositional layering is more plausible for the origin of the D″ discontinuity. The D″ layer can be an “anti-crust” made up mostly of TTG + anorthosite derived from the former continental crust. Tectonic style of the anti-crust at the CMB is similar to that at the surface. At both places, chemically distinct layers are density stratified and are also characterized by the processes of accretion, magmatism, and metasomatism. © 2008 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction The top and bottom of Earth's mantle are thermal boundary layers and the tectonic processes operating in these regions are the key to discuss the dynamics, chemistry, and evolution of the mantle. At the top of the mantle, buoyant continental crust is floating, which was formed as a consequence of mantle melting, via slab or lower crust melting. When crossing the solidus, rocks are subjected to partial melting which is one of the most significant chemical processes in the Earth because it leads to the chemical differentiation (Bowen, 1928). The top of the mantle must have extensively and diversely evolved with respect to its chemical characteristics. On the other hand, evolution of the bottom of the mantle has been debated for a long time, and yet very little is known about the processes (e.g., Knittle and Jeanloz, 1991; Ishii and Tromp, 1999; Kellogg et al., 1999; Murakami et al., 2004; Hirose, 2006; Maruyama et al., 2007; Garnero and McNamara, 2008). The so-called D″ layer has been recognized as a distinct region with an average thickness of 250– 265 km (Wysession et al., 1998). This layer is a thermal and chemical boundary layer similar to the lithosphere on the surface since it is in contact with the metallic liquid outer core. Subducted former mid-

⁎ Corresponding author. Department of Earth and Planetary Sciences, Tokyo Institute of Technology, 2-12-1 Ookayama, Meguro, Tokyo 152-8551, Japan. Tel.: +81 3 5734 2618; fax: +81 3 5734 3538. E-mail address: [email protected] (T. Komabayashi).

oceanic ridge basalt (MORB) crust has been considered to form the D″ layer because (1) through geologic time, large amounts of former MORB crust must have been subducted into the mantle (e.g., Reymer and Schubert, 1984; Komiya, 2004), and (2) MORB crust is denser than the surrounding mantle all the way to the core–mantle boundary (CMB), suggesting that MORB crust goes down to the CMB (Hirose et al., 2005). The bottom of the mantle is also undergoing chemical differentiation since it has been seismologically confirmed to be melt-present region. A thin (20–40 km) layer with a great reduction in seismic velocity (N10%) was discovered at the base of the mantle (Mori and Helmberger, 1995; Garnero and Helmberger, 1996). Origin of this ultralow-velocity zone (ULVZ) was discussed in relation to partial melting of silicates. Zerr et al. (1998) discussed the role of a partial melting of the mantle peridotite on the formation of the ULVZ. However, if the subducted former MORB crust exists in the D″ layer, partial melting should preferentially occur in it since its melting temperature is lower than that of the mantle peridotite by ca. 250 K (Hirose et al., 1999). Thus, the D″ layer was thought to be a mixture of peridotite, subducted MORB crust, its partial melt and restite. In this study, we propose there would be other candidates for the chemical variations in the D″ layer. Rino et al. (2004) constructed a growth curve of the continental crust through geologic time. They demonstrated that large amounts of primordial continental crust are missing on the present-day Earth's surface, suggesting that these were subducted into the mantle in spite of a large density barrier. If such

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T. Komabayashi et al. / Gondwana Research 15 (2009) 342–353

subduction occurred, this might have given rise to a wide variation of the chemistry and the density structure in the D″ layer. Prediction of the detailed structure of the D″ layer is a key to evaluate the chemistry and dynamics of the bottom of the mantle. In this paper, we perform phase assemblage analysis on several rock systems down to the CMB conditions. From high-pressure mineral physics data, we show that it is possible for primordial continental crust, comprising mostly of anorthosite and tonalite– trondhjemite–granite (TTG), to be major components of the D″ layer. Furthermore, we propose the stratigraphy of the D″ layer based on the density stratification containing subducted continental crust. Subducted continental crust will have a high seismic velocity than the pyrolitic mantle, and therefore, the boundary between them will produce a seismic discontinuity. This discontinuity may correspond to the D″ discontinuity, instead of a (Mg,Fe)SiO3 post-perovskite (PPv) phase transition. Considering the global material circulation and the binary phase transition intervals in the PPv transition, we propose that a bulk chemistry change is more plausible for the origin of the D″ discontinuity than the phase transformation model. Finally, we discuss the tectonic style of the D″ layer as an anti-crust at the bottom of the mantle. 2. The D″ layer and D″ discontinuity At the bottom of the mantle, the D″ layer has long been noticed as seismologically unique layer with an average thickness of 250 km although the thickness varies depending on the place. Historically, D″ layer was first named as the region of the bottom 200 km of the lower mantle where seismic velocity gradient was reduced from the layer right above (D′) (Bullen, 1949). Subsequently, in 1980's, seismologists found an S-wave discontinuity at the top of the D″ layer (Lay and Helmberger, 1983), followed by the discovery of a P-wave velocity jump (Wright et al., 1985). Then, this discontinuity was termed as the D″ discontinuity. The D″ layer was earlier considered to be a thermal anomaly, since it is in contact with the hotter outer core. This idea was consistent with that the seismic velocity of the materials being reduced with increasing temperature. However, it was rather difficult to explain the abrupt increase in the seismic velocity at a certain depth (D″ discontinuity) by the thermal anomaly alone (e.g., Wysession et al., 1998). Later, the D″ layer was reconsidered as a pile of chemically distinct materials. In this scenario, the D″ discontinuity can be explained by a bulk rock chemistry change. This idea was supported by the fact that through geologic time, large amounts of former oceanic plates were subducted into the deep mantle (e.g., Silver et al., 1988; Komiya, 2004). Materials derived through the interaction between the solid mantle and liquid outer core were other alternate candidates to account for the D″ layer (e.g., Knittle and Jeanloz, 1991). Presence of a ULVZ at the bottom of the D″ layer made it more convincing that the D″ layer was a chemically distinct region (e.g., Garnero et al., 1998). On the other hand, Sidorin et al. (1999), from seismic data, proposed a first-order solid–solid phase transition with a Clapeyron slope of 6 MPa/K for the origin of the D″ discontinuity. However, from mineral physics research, until 2004, both high-pressure experiments and theoretical works repeatedly claimed that the major constituent minerals of the lower mantle, (Mg,Fe)SiO3 perovskite and (Mg,Fe)O ferropericlase did not undergo any first-order phase transition all the way to the bottom of the mantle (Knittle and Jeanloz, 1987; Wentzcovitch et al., 1993; Serghiou et al., 1998). In 2004, however, a PPv phase transition in MgSiO3 perovskite shed light on the solid– solid transition scenario (Murakami et al., 2004). This transition reasonably explains the nature of the D″ discontinuity, including the depth and the magnitude of the seismic velocity jump. The depth of the D″ discontinuity estimated by both P- and S-waves varies from 100 to 450 km above the CMB or even absent in some areas (Wysession

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et al., 1998). This variation is interpreted as the temperature difference along the PPv transition with a Clapeyron slope of 5 to 11.5 MPa/K (Hirose, 2006), and its absence may indicate that the temperature is high enough not to cross the transition. This transition can also explain paired D″ discontinuities as a geotherm crosses the PPv transition twice (Hernlund et al., 2005). Thus, the nature of the D″ discontinuity was explained by the post-perovskite phase change, leading to an idea that the D″ layer is dominated by (Mg,Fe)SiO3 postperovskite. However, the discovery of post-perovskite cannot eliminate the possibility of the occurrence of subducted MORB crust in the D″ layer when one considers global material circulation. Subducted MORB crust may be built up to a 300 km thick zone immediately above the CMB if all the MORB material produced during the last 4.0 Ga are deposited on the bottom of the mantle. Ohta et al. (2008) showed that the PPv transition in a MORB system occurs at 4 GPa shallower than in a pyrolite system and considered the difference in chemistry between MORB and pyrolite to explain multiple seismic discontinuities at the bottom of the Pacific superplume. Thus, the PPv transition is now expected to be a good indicator of the chemistry change at the CMB. However, this polymorphic phase transformation belongs to the continuous binary reaction. Binary diagrams in MgSiO3–FeSiO3 (Tateno et al., 2007) and MgSiO3–Mg3Al2Si3O12 (Akber-Knutson et al., 2005; Tateno et al., 2005), demonstrated that the phase loops of the PPv transitions are so broad up to 120–180 km in mantle compositions. Ohta et al. (2008) determined the width of PPv transitions in a pyrolitic mantle and MORB systems. In both systems, the transitions occur over 110–145 km intervals. Even though the seismologically observed sharpness of 50–75 km (Wysession et al., 1998) should be narrower than the experimentally determined phase loop (Stixrude, 1997), these estimations seem too broad. In addition, a seismic anomalous pile with a sharp vertical edge was reported beneath the Indian Ocean (Wen, 2001), which cannot be explained by the pressure-dependent PPv transition. We start our discussion below by presenting the phase relations of a subducted MORB to the CMB conditions, followed by those of continental crust (anorthosite and TTG). In the later sections, we will emphasize the importance of the continental crust for the D″ layer components as a consequence of the global material circulation through geologic time. Finally, we will discuss a possible explanation for the D″ discontinuity by the bulk rock chemistry change between subducted materials and the mantle, which can produce a sharp discontinuity. 3. Subducted MORB crust in the mantle The minimum estimate for the amount of subducted former MORB crust over the last 4.0 Ga is 11 vol.% of the entre mantle (Reymer and Schubert, 1984; Komiya, 2004; Ohta et al., 2008). This value corresponds to about 300 km thick layer at the base of the mantle, comparable to the value of the observed thickness of the D″ layer. The distribution of the former MORB crust in the mantle has been debated. Recent interpretation of the origin of hotspot magma such as in Hawaii supports the idea of the common occurrence of recycled fragments of MORB in the mantle (Takahashi and Nakajima, 2002). Kaneshima and Helffrich (1999, 2003) found platy fragments with a dimension of 500 × 300 km and thickness of 8 km at depths of 1100– 1850 km off Izu–Ogasawara islands under the Pacific Ocean. A key observation in addition to its platy shape is that the S-wave velocity was largely reduced up to 4% while the P-wave velocity showed no significant change. Komabayashi et al. (2007) interpreted this Svelocity drop as the ferroelastic phase transitions in both SiO2 phase and calcium-perovskite (CaPv) in the subducted MORB crusts. This suggested that the subducted MORB crust is likely present in a wide depth range in the lower mantle, without accumulating at the 660-km boundary (e.g., Karato, 1997). From in-situ density measurements,

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Hirose et al. (2005) showed that subducted MORB crust is always denser than the surrounding mantle in the lower mantle conditions. Therefore, the MORB crust at the top of subducting lithosphere is likely going down to the CMB. 4. High-pressure phase relations of the MORB system Here we describe the phase relations of a subducted MORB crust to the CMB conditions. Since the MORB system is likely to have the lowest melting temperature among possible candidates for the D″ layer (Hirose et al., 1999), the nature of the partial melt derived from the MORB crust is also discussed. 4.1. Subsolidus and melting relations of the MORB system MORB crust transforms to garnetite (majorite garnet + stishovite) at 15 GPa (Irifune et al., 1986). Melting phase relations of a MORB were investigated up to 27 GPa (Hirose and Fei, 2002), while the subsolidus relations were estimated up to the CMB pressure (Hirose et al., 2005; Ohta et al., 2008). Beyond the eclogite–garnetite transformation, CaPv and Ca–Al-silicate phase (CAS) are encountered at 22 GPa and 25 GPa, respectively, at subsolidus conditions. Furthermore, at 27 GPa, Mg–Fe– Al-phase with a calcium ferrite structure (CF) was found in the MORB

system. Magnesium-perovskite (MgPv) replaces majorite garnet (Mj) at 26–27 GPa (Hirose and Fei, 2002). In order to predict the phase relations at further higher pressures, we performed phase analysis in the MORB at 27 GPa around 2200 °C, focusing on the melting relations using a metamorphic petrology approach. We analyzed the phase relations in the framework of the ACF diagram where A, C, and F mean Al2O3, CaO, and FeO + MgO, respectively (Fig. 1). Note that A is only Al2O3, since the assemblage does not contain plagioclase at such a high pressure. Since MORB is a silica-saturated system, SiO2-phase, i.e., stishovite (St) is always present. Therefore, the ACF diagram is projected from stishovite. For simplicity, the Fe–Mg binary effect is not considered. Fig. 1 is a Schreinemakers analysis on experimental results (Hirose and Fei, 2002; Ishibashi et al., 2008). An important phase relation above 30 GPa is that two melt-forming reactions occur in a narrow temperature range between 2300 and 2400 °C, with increasing temperature MgPv + CaPv + CF+ St = melt, followed by CaPv + CF + St = CAS + melt. These two reactions behave as solidus reactions. Another notable point is that above these solidi, no Mg–Fe solid phase is stable, indicating that most of the Fe in the system goes into the melt. The composition of the melt is, therefore, FeO-rich such that the melt is dense, most likely to be denser than the solid residue. Moreover, beyond CAS phase stability, i.e., above 40 GPa, Schreinemakers bundle predicts

Fig. 1. Schreinemakers bundle for the MORB system, focusing on the solidus relationships. The web was constructed from experimental data in a MORB system (Hirose and Fei, 2002) and the pure CAS phase (CaAl4Si2O11) stability (Ishibashi et al., 2008). Chemographic analysis was made in the framework of the ACF diagram. Inset at upper left shows the composition of each phase together with the bulk composition (MORB) in Hirose and Fei (2002) which is projected onto the ACF triangle from stishovite (SiO2). Note that the system is silicasaturated, so that stishovite is always present. Solid lines are the reactions constrained by the experiments while broken lines are predicted from the Schreinemakers rule. The Fe–Mg binary effect is not considered for simplicity. The P–T locations of decomposition of CAS were constrained by Ishibashi et al. (2008), although the composition in the web is not pure CaAl4Si2O11. Above 30 GPa, two melting reactions occurs in a narrow temperature range. Beyond CAS stability above 40 GPa, the restite phase assemblage would be CaPv + Cor + St, which is identical to the anorthosite assemblage. See text for details. Abbreviations of the phases are listed in Table 1.

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that the solid phase assemblage above the solidi, i.e., restite, is CaPv + corundum (Cor) + St which is identical to anorthosite assemblage as would be discussed later. Therefore, at the CMB the restite would have anorthosite assemblage. As discussed below, anorthosite is denser than the mantle at the CMB, implying that both partial melt and restite would be denser than the mantle. Thus, at the CMB, no chemical plume is expected to rise through the mantle, if no additional light elements are added from the underlying metallic core. Fig. 2 is a phase diagram in the MORB system to the CMB conditions, inferred from the Schreinemakers bundle (Fig. 1) and experimental studies (Hirose et al., 1999, 2005; Ohta et al., 2008). Three major phase transformations occur beyond 30 GPa. With increasing pressure, stishovite = CaCl 2 -type SiO 2 , CaCl 2 -type SiO2 = α-PbO2-type SiO2, and the PPv transition in (Mg, Fe)SiO3. Note that the PPv transition has a two-phase loop of 8 GPa wide and α-PbO2–SiO2 transition occurs in this loop condition (Ohta et al., 2008). The solidus temperature was placed about 4000–4500 °C at the CMB (Hirose et al., 1999). A major assumption is that the melt composition changes little from 27 GPa to 135 GPa, so as not to alter the Schreinemakers web topology. 4.2. FeO-rich partial melt from the MORB crust Partial melting of the subducting MORB crust at the shallower subduction zone produces granitic melt, which is one of processes for the formation of the continental crust (e.g., Drummond and Defant, 1990). Granitic to andesitic melts have densities of 2.2–2.4 g/cm3 at 1300 °C and 1 atm (Hess, 1989), whereas the surrounding mantle has a density of 3.4 g/cm3 at its top (Irifune and Ringwood, 1987). Therefore, the generated melt has strong buoyancy. On the other hand, the melting behavior of the subducted MORB crust at the deep mantle is markedly different from that at the surface. Hirose and Fei (2002) showed that the partial melt is strongly enriched in FeO compared both with the restite and bulk MORB at about 27 GPa. The Schreinemakers analysis above demonstrates that most of FeO in the system goes into the melt above the solidi. Ohtani and Maeda (2001) predicted that a density crossover between a MORB melt (100% melt of a MORB composition) and the surrounding peridotite occurred at the base of the mantle. Therefore, the partial melt of the MORB crust at the base of the mantle is denser than the surrounding mantle, if the composition of the partial melt changes

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little from 27 GPa to 135 GPa not to alter the Schreinemakers web. The gravitationally stable dense melt would accumulate on the bottom of the CMB and grow through time over 4.0 Ga. 5. Anorthosite as a primordial continent After the collision–amalgamation of planetesimals to form the Earth in the primordial solar system at 4.56 Ga, a primordial crust composed of anorthosite must have been formed by the consolidation of a magma ocean at 4.5–4.4 Ga. On the Moon, anorthosite layer was formed as a primordial continent of 50–60 km thick (e.g., Wieczorek and Phillips, 1998). This layer may have been floating within the magma ocean as a result of melt-crystal density crossover. The volume ratio of anorthosite over the mantle in the Moon gives us a rough estimate about the primordial anorthosite continent in Hadean Earth. The anorthosite layer would have reached a thickness of up to 220 km on the surface of the Earth. Such a huge mass of anorthosite can not be identified at any place on the present-day Earth. One possible reason for the missing anorthosite is that it was removed to the deep mantle. Despite the density barrier at the surface, this was possible because the anorthosite crust was thick enough to transform into grossular garnet + kyanite + quartz at about 30–60 km depths, an assemblage that is denser than the mantle. This deeper part of the anorthosite layer would have removed into the mantle by a convective flow. However, the shallower part would have remained on the surface until plate tectonics began to operate. As discussed below, anorthosite soon becomes much denser with subduction. The anothosite layer would become about 300 km thick right above the CMB if all anorthosite had been accumulated there. Note that the estimate of 300 km thickness is based on the assumption that the volume of the rock at the CMB is a half of that at the surface due to compression, as mentioned later. 6. High-pressure transformations of anorthosite Here we describe the high-pressure phase relations of anorthosite to 135 GPa, the CMB pressure (Fig. 3a). Since no direct experimental data at such high pressures are available, the phase relations discussed here are combinations of each end-member minerals from existing experimental and theoretical works. Phases are listed in Table 1.

Fig. 2. Phase relations for the MORB composition. The bulk composition is from Hirose and Fei (2002). Phase transitions above 40 GPa (Fig. 1) is St = CaCl2–SiO2 (Hirose et al., 2005), CaCl2–SiO2 = α-PbO2–SiO2, and the post-perovskite change (Ohta et al., 2008). The phase assemblage in each sector is listed in Table 2. A hypothetical mantle geotherm is shown as a grey line. St, stishovite; MgPv, magnesium–perovskite; MgPPv, magnesium-post-perovskite; CMB, core–mantle boundary.

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Fig. 3. Phase relations for the continental crust: (a) anorthosite represented by anorthite and (b) TTG represented by albite + quartz (1:7 in mole). Reference for each reaction is, An = Gr+ Ky+ Qz (Koziol and Newton, 1988), Ab= Jd+ Qz (Holland, 1980), Qz= Coe (Bose and Ganguly, 1995), Coe = St (Zhang et al., 1996), Jd= NAL + St (Akaogi et al., 2002), Ky= Cor + St (Ono et al., 2007), Gr= CaPv+ Cor (tentatively assumed) (Takafuji et al., 2002), St = CaCl2–SiO2 (Ono et al., 2002), CaCl2–SiO2 = α-PbO2–SiO2 (Murakami et al., 2003), Cor = Rh2O3–Al2O3 (Oganov and Ono, 2005), and Rh2O3–Al2O3 = PPv-Al2O3 (Oganov and Ono, 2005). The phase assemblage in each sector is listed in Table 2. A hypothetical mantle geotherm is shown as a grey line.

Anorthosite is represented by anorthite (CaAl2Si2O8) as a first order of approximation. As mentioned above, at 1–2 GPa, the phase assemblage is grossular (Gro) + kyanite (Ky) + quartz (Qz) (Koziol and Newton, 1988). Quartz undergoes polymorphic phase transition twice up to 10 GPa, Qz = coesite (Coe) (Bose and Ganguly, 1995) and Coe = St (Zhang et al., 1996). Kyanite is decomposed at about 12 GPa by a reaction kyanite = corundum + stishovite (Ono et al., 2007). A decomposition of grossular garnet to a mixture of calcium-perovskite + corundum was found at 23–25 GPa by Takafuji et al. (2002). In Fig. 3, we tentatively put a reaction grossular = calcium-perovskite + corundum at 25 GPa, although the P–T location of this reaction was not tightly constrained. At 60–80 GPa, stishovite is transformed to CaCl2type SiO2 phase (Ono et al., 2002). This reaction is a second-order phase transition where no discontinuous volume change is accompanied. At the similar pressures, a transformation of corundum to Rh2O3-type Al2O3 was suggested with a negative dP/dT slope by a first principles theory (Oganov and Ono, 2005). At further higher pressures of 110–120 GPa, both SiO2 and Al2O3 undergo phase transitions again, CaCl2–SiO2 = α-PbO2–SiO2 (Murakami et al., 2003) and Rh2O3–Al2O3 = post-perovskite (PPv)-type Al2O3 (Oganov and Ono, 2005). The PPv-Al2O3 phase transition has a negative dP/dT slope. The highest pressure assemblage is Ca-Pv + PPv-Al2O3 + α-PbO2–SiO2. The phase assemblage in each sector in Fig. 3 is listed in Table 2 together with zero-pressure density. Zero-pressure density profile to about 30 GPa is illustrated in Fig. 4. 7. TTG into the mantle: arc subduction, tectonic erosion and sediment subduction In this section, we discuss the growth history of continental crust, which involves the creation from the mantle and removal back into the mantle. In addition to these two processes, recycling of continental

crust is an important process for the estimation of the growth rate. Recycling involves the return of subducted crust components to the surface as magmas in subduction zones. In this paper we focus on the CMB, and the processes important for our discussion are the creation and subduction of continental crust. We do not discuss here the recycling aspect, which is considered in detail by Rino et al. (2004). Fyfe(1978), Brown (1979), O'Nions et al. (1979), Armstrong (1981), Dewey and Windley (1981), McLennan and Taylor (1982), Reymer and Schubert (1984), and several other workers considered that the most of continental crust, more than 80%, was formed by the end of Archean because of high heat-flow in the Archean and the slow growth rate in

Table 1 Phases in anorthosite and TTG systems. Phase

Abbreviation

Composition

Density (g/cm3)

Reference

Quartz Coesite Stishovite CaCl2-type SiO2

Qz Coe St CaCl2–SiO2

SiO2 SiO2 SiO2 SiO2

2.65 2.91 4.29 4.29

α-PbO2-type SiO2 anorthite albite jadeite Na–Al-phase corundum Rh2O3-type Al2O3 PPv-type Al2O3 kyanite grossular calsium perovskite Ca–Al-silicate

α-PbO2–SiO2 An Ab Jd NAL Cor Rh2O3–Al2O3 PPv-Al2O3 Ky Gro CaPv CAS

SiO2 CaAl2Si2O6 NaAlSi3O8 NaAlSi2O6 NaAlSiO4 Al2O3 Al2O3 Al2O3 Al2SiO5 Ca3Al2Si3O12 CaSiO3 CaAl4Si2O11

4.316a 2.76 2.62 3.34 3.917 3.99 4.09 4.2 3.675 3.595 4.23 3.888

Robie et al. (1978) Robie et al. (1978) Nishihara et al. (2005) Assumed same as stishovite Murakami et al. (2003) Robie et al. (1978) Robie et al. (1978) Robie et al. (1978) Akaogi et al. (2002) Oganov and Ono (2005) Oganov and Ono (2005) Oganov and Ono (2005) Robie et al. (1978) Robie et al. (1978) Zhang et al. (2006) Ono et al. (2005)

a

The density was estimated as 0.6% denser than CaCl2-type SiO2.

T. Komabayashi et al. / Gondwana Research 15 (2009) 342–353

7.2. Selective subduction of juvenile arcs in the western Pacific and its implication for the Archean arc subduction

Table 2 Phase assemblage and zero-pressure density. Rock type

Sector in phase

Phase assemblagea

MORB

b

Anorthosite

TTG

1 2 3 4 5 6 7 8 9 1 2 3 4 5 6 7 8 9 10 11 12 1 2 3 4 5 6 7

Zero–P Density (g/cm3)

diagram MgPv + CaPv + CAS + St Mj/MgPv + CaPv + CAS + Cor + St MgPv + CaPv + CF + St CaPv + CAS + St + Melt CaPv + Cor + St + Melt MgPv + CaPv + CF + CaCl2–SiO2 CaPv + Cor + CaCl2–SiO2 + Melt MgPPv + CaPv + CF + α-PbO2–SiO2 CaPv + Cor + α-PbO2–SiO2 + Melt An Gro + Ky + Qz Gro + Ky + Coe Gro + Ky + St Gro + Cor + St CaPv + Cor + St CaPv + Cor + CaCl2–SiO2 CaPv + Rh2O3–Al2O3 + St CaPv + Rh2O3–Al2O3 + CaCl2–SiO2 CaPv + PPv-Al2O3 + CaCl2–SiO2 CaPv + Rh2O3–Al2O3 + α-PbO2–SiO2 CaPv + PPv-Al2O3 + α-PbO2–SiO2 Ab + QZ Jd + Qz Jd + Coe Jd + St NAL + St NAL + CaCl2–SiO2 NAL + α-PbO2–SiO2

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4.29 – – – – – – – – 2.76 3.53 3.56 3.67 3.82 4.15 4.15 4.19 4.19 4.23 4.20 4.24 2.64 2.82 3.03 3.96 4.21 4.21 4.23

a Abbreviations not listed in Table 1 are, MgPv, magnesium-perovskite; Mj, majorite garnet; CF, Mg–Fe–Al-phase with the CF structure; MgPPv, magnesium-post-perovskite. b Zero-pressure density for MORB is not available at higher pressures.

the Phanerozoic. However, the newly estimated crust formation curve by Rino et al. (2004, 2008) does not support these previously established concepts. 7.1. Missing Archean continental crust on the present Earth's surface The rock records on the Earth indicate that the Archean continental crust was predominated by TTG, accreted MORB, and plateau basalts including komatiite and fractionated rocks. The Archean continental crust was formed as an intra-oceanic island arc, presumably after the change of plate boundary from transform fault to subduction zone, due to change of relative plate motion. Intra-oceanic arcs collide each other with time to become much bigger composite arcs or primitive continents, as best documented in Pilbara (Kitajima, 2003), North America (Hoffman, 1989), and many other Archean cratons. For example, 6–7 intra-oceanic arcs were amalgamated together to generate a primitive continent in the Pilbara craton during 3.5–3.1 Ga. Style of orogeny is similar to that of the on-going examples such as those in the western Pacific triangular region (Maruyama et al., 2007). The size of Archean oceanic plate was as small as in the western Pacific of about 700 km across or less, equivalent to the thickness of the upper mantle. The size of orogen is also in the same order of magnitude b600 km long. The number of oceanic microplates (700 km across) was calculated to be about 400. The total length of plate boundary must have been twice as that of the modern one, ca. 40,000 km. This suggests that the production rate of the continental crust was double as well. However, the estimated continental growth curve (Rino et al., 2004, 2008) demonstrated that the net growth of the continental crust by the end of Archean (i.e., amount of the continental crust formed in Archean and still staying on the present surface) accounts for only 20 vol.% or less of the present amount. This discrepancy suggests that most intra-oceanic arcs were subducted into deep mantle, as seen in the modern western Pacific.

As noted above, most of the Archean arcs are not preserved on the present surface of the Earth. The reason for the missing arcs can be investigated in the western Pacific since this is the modern example of microplate region. In this region, juvenile island arcs are present that account for over 70% of the arcs in the world (Maruyama et al., 2007). An important point is that among of the arc–arc collisions in this area, amalgamation can only be possible in parallel collision such as Celebes. Otherwise, one arc is subducted beneath the other (e.g., Yamamoto et al., 2009-this issue). The parallel collision and amalgamate of intra-oceanic arcs were well-documented in the Archean orogenic belts over the world (e.g., Yamamoto et al., 2009this issue). Thus, the modern analogues of arc–arc collision indicate that most of the arcs were subducted into the mantle. This may be the reason for the missing Archean arcs on the present day globe. However, once the critical size of continental crust is exceeded, such as 35 km in thickness, the continental crust began to increase its volume effectively at the surface. The timing for this was probably around 2.7 and 2.0 Ga (Rino et al., 2004; Maruyama et al., 2007). Presumably 5.8 times the mass of the present continental crust had been subducted into deep mantle before 2.0 Ga, and thereafter the continental crust has accumulated efficiently on the surface of the Earth, although considerable amounts have also been subducted subsequently, as seen even today in the western Pacific, through processes of tectonic erosion, arc subduction, and sediment-trapped subduction (Maruyama et al., 2007). 7.3. Continental subduction, tectonic erosion, and sediment subduction in the Proterozoic and Phanerozoic During the Proterozoic, the formation of the continental crust was operated most effectively, generating over 70% of the present volume (Rino et al., 2004). From 2.8–2.7 to 2.2–2.3 Ga, the collision of the primordial continents (composite arcs) built up more rectangularshaped continents in addition to arc–arc collisions. The collision-type orogenic belts such as the Limpopo and Capricorn orogens were developed in this period. After the mantle overturn at 2.8–2.7 Ga, the episodic whole mantle convection would have started, to create the large continental plates as seen today. The clear evidence for ca. 3000 km wide continental plate is the presence of 1.9–1.8 Ga Laurentia (Hoffman, 1989). The continental crust, however, continued to be subducted into the mantle similar to that in the Archean.

Fig. 4. Zero-pressure density profile to 27 GPa for TTG (this study), anorthosite (this study), MORB (Irifune and Ringwood, 1987; Hirose et al., 1999), pyrolite (Irifune and Ringwood, 1987), and harzburgite (Irifune and Ringwood, 1987). The profiles for anorthosite and TTG are calculated along the geotherms in Fig. 3. The density data are listed in Table 2. See text for details.

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By the time the Paleoproterozoic supercontinent Columbia was assembled, ca. 50% of the present-day continental crust was completed. Proterozoic effective growth of continental crust occurred in the late Proterozoic time after 1.5 Ga, specifically at 1.0 and 0.8– 0.6 Ga (Rino et al., 2008). At 1.0 Ga the supercontinent Rodinia was completed. The late Proterozoic was marked by the breakup of Rodinia and the formation of Gondwana, the latter considered as a semisupercontinent (Senshu et al., 2009-this issue). Completion of Gondwana was the time of the onset of the Phanerozoic, (the magic number of 542 Ma) marking the closure of the Mozambique Ocean between the East and West Gondwana. Since then Baltica collided during the early Paleozoic and Africa collided to form the Hercynian orogens around North America. The major growth time of continental crust has had a peak period during the Cretaceous, ca. 120–80 Ma (Rino et al., 2008). New continental crust included those of the circum-Pacific and Tethyan orogenic belts and the oceanic regions with a number of arcs in the Altai–Alaska–Middle America. In the Phanerozoic, tectonic erosion at trench is common around the active trenches, and even in the case where there is no tectonic erosion, trench–turbidite can be transported into deep mantle, because subducted oceanic lithosphere bends before trench-axis to yield horst-graben structure reaching to a 500 m high wall. These grabens are filled by trench–turbidite at trench, and move down into deep mantle transporting granitic sediments (Senshu et al., 2009-this issue).

7.5. Subduction of the continental crust: Discovery of UHP rocks The Archean continental crust is TTG whose major constituents are albite + quartz (1:7, in mole). Albite breaks down to jadeite + quartz at 1–2 GPa. Hence volumetrically 75% of TTG crust is composed of SiO2 phase at depths deeper than 40–50 km. Until 1984, subduction of continental crust beyond the Moho depth had not been accepted because of the large density barrier. Field geology such as those in the Himalayas (England and Thompson, 1984) supported the notion that collision and subduction never exceed Moho depth. However, the discovery of coesite, a highpressure polymorph of quartz (Chopin, 1984), followed by the discovery of diamond (Sobolev and Shatsky, 1990) from regional metamorphic belts revolutionized the paradigm about the continental subduction. Researches on the ultrahigh-pressure (UHP) metamorphism in collisional orogenic belts suggest that the continent subduction reaches to the depth 200–300 km depth (Liou et al., 2002). The discovery of the UHP rocks indicates that the subduction of continental lithosphere is not controlled by the density of continental crust alone but by the net density of subducting materials, i.e., continental crust, lower mafic crust and peridotitic layer. The continental crust is being subducted, coupled with the subducting lithosphere. Moreover, the tectonic erosion and sediments-trapped graben structures mentioned above are also mechanically coupled with the lithosphere. 8. High-pressure transformations of TTG

7.4. Mantle metasomatism by subducted granite for the origin of tectosphere The foregoing discussion strongly suggests that the missing ancient TTG are stored in the deep mantle. One possible site is the root of the tectosphere. We estimated the amounts of TTG absorbed by the shallow mantle for the origin of the tectosphere. High-geothermal gradients were recorded in the Archean subduction zones. Therefore, the slab-melting to produce adakitic TTG melt might have commonly operated, which is rare in the modern Earth. If arc subduction occurred, hydrated granitic rocks must have selectively melted to yield silicic melt which metasomatically reacted with the mantle to form orthopyroxenite. The reacted zone may correspond to the root of the tectosphere (Jordan, 1979). The tectosphere is only developed under the cratons which formed before 2.0 Ga. Mantle xenoliths from the tectosphere are different in composition from abyssal peridotites. Those xenoliths are enriched in orthopyroxene up to 40 modal %, compared to a lherzolite with only 10% (Jordan, 1979), indicating that the tectosphere is enriched in SiO2, by about 2–10 wt.% compared with pyrolite. Assuming this excess SiO2 is derived from pre-2.0 Ga TTG by subduction zone metasomatism, we made a simple calculation for the amount of Archean granite as a metasomatic agent. An S-wave tomography of the whole mantle (Grand, 2002) provides an estimation of the volume of the tectosphere. Assuming the bottom of the tectosphere is at 300 km depth and the surface extension of pre-2.0 Ga craton of 5.7 × 107 km2, we calculated the accreted granitic mass to be about 5–10% of the present-day granitic crust. Although the above calculation is based on the simple assumptions, this estimate is too small to account for the total mass of the subducted Archean TTG. Another possible storage site in the mantle for the huge amounts of TTG is the base of the mantle (D″ layer) as discussed later. The D″ layer can accommodate 20 times as much granite as the surface continental crust, since the density of SiO2 phase at the CMB (5.60 g/cm3) is much higher than at the surface (2.65 g/cm3), i.e. the mass can be compressed at the CMB. Note that the value of 5.60 g/cm3 is the density at a high pressure, not the zero-pressure density used in the later discussions (Murakami et al., 2003).

Here we investigate the phase assemblage in the system for TTG which does not contain potassium, down to the CMB conditions. Highpressure phase transformation of TTG is summarized in Fig. 3b, and the phases are listed in Table 1. The TTG is approximated by albite + quartz (1:7 molar ratio). At 10 GPa, the phase assemblage is jadeite + stishovite. Jadeite is decomposed into sodium–aluminum phase (NAL) + stishovite at about 22 GPa (Akaogi et al., 2002). No study has yet been reported about further high-pressure phase transition of NAL. Therefore, we assume this phase to be stable up to the CMB pressure. The SiO2 phase undergoes polymorphic transitions at both 60–80 and 110–120 GPa as shown in the case of anorthosite. The highest pressure assemblage is NAL+ α-PbO2–SiO2. The zero-pressure density profile is shown in Fig. 4. 9. Density arrangement of slab-peridotite, TTG, anorthosite, MORB, and FeO-enriched melt derived from MORB In this section, we discuss the densities of several possible rock candidates at the CMB. Absence of in-situ density data for most of minerals of anorthosite and TTG systems constrains us to make comparison on the basis of the zero-pressure density data (Table 2). Anorthosite and TTG have zero-pressure densities of 4.24 and 4.23 g/cm3, respectively at depths of 2700–2900 km. Zero-pressure densities of MORB (Hirose et al., 1999), pyrolite (Irifune and Ringwood, 1987), and harzburgite (Irifune and Ringwood, 1987) at 27 GPa are 4.29, 4.19, and 4.10 g/cm3, respectively. At the bottom of the mantle, these three rocks are subjected to further phase transition in SiO2 or (Mg,Fe)SiO3. Therefore, zero-pressure densities at the CMB should be higher than at 27 GPa. The density of partial melt of the MORB at the CMB conditions is unknown. However, it should be denser than the MORB because the melt is FeO-rich compared to the MORB crust as discussed above. Therefore, we tentatively assume that the partial melt is the densest material among these rock types. In addition, as discussed in the previous section from the Schreinemakers bundle, the restite of MORB partial melting would have the anorthosite assemblage at such high pressures. Thus, on the basis of the density, the order of arrangement at the CMB can be inferred to be the partial melt of the MORB, MORB crust, anorthosite, TTG, pyrolite, and harzburgite.

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10. Discussions 10.1. Mechanism to transport anorthosite to deep mantle In the present day Earth's surface, the huge mass of anorthosite expected to be present cannot be identified at any place. Thus, the anorthositic crust may have been subducted to the deep mantle despite the density barrier. However, as mentioned earlier, the subduction is controlled by the net density. Therefore the subduction of a lighter rock is possible if it is coupled with a denser rock. In addition, in the case of anorthosite, it is less dense than the mantle, only at the surface. It becomes an assemblage at 1–2 GPa of garnet + kyanite + quartz which is denser than the surrounding mantle. Fig. 4 shows zero-pressure density profiles of rocks with depth. It illustrates that anorthosite is always denser than the mantle (pyrolite) down to 660 km depth (23.5 GPa). At 660 km depth, pyrolite is subjected to the post-spinel transition giving rise to a large density jump. For anorthosite, due to the decomposition of grossular garnet at about 25 GPa, the density of the rock jumps from 3.82 to 4.15 g/cm3. Below this depth, anorthosite is slightly less dense than the surrounding mantle in terms of zero-pressure density (Fig. 4). However, subducting materials are colder than the mantle. This thermal effect is significant (Irifune and Ringwood, 1987) and this lets subducting rocks to be denser than the surrounding mantle at all depths. This suggests that anorthosite is easily subducted to the deep mantle conditions. 10.2. Mechanism to transport TTG to deep mantle Considering the growth curve of continental crust on the Earth, we conclude that the Archean TTG was removed to the deep mantle. The TTG thus removed may reach to 6 times the mass of the total TTG on the present surface. According to the density profile of subducted TTG in relation to the mantle pyrolite (Fig. 4), at pressures below 9 GPa, TTG is the least dense rock. This was the very reason for the former consensus against the continental subduction. Nevertheless, UHP rocks demonstrated that the continental rocks were indeed subducted to 7–10 GPa and showed that the subduction is controlled by the net density of the subducting materials. In terms of zero-pressure density, upon formation of stishovite at 9 GPa, the density of TTG jumps dramatically to 3.96 g/cm3. Thereafter, TTG is always denser than the surrounding mantle down to the lower mantle. At 660 km depth, the mantle undergoes the post-spinel transition. However, at similar depth, TTG has a decomposition reaction of jadeite by Jd = NAL + St, leading to a density jump of the rock from 3.96 to 4.21 g/cm3. In addition, due to the thermal expansion effect as mentioned in the case of anorthosite subduction, TTG becomes denser than the surrounding mantle at all depths after the stishovite formation. Thus, the TTG would have accumulated on the bottom of mantle, together with Archean subducted MORB.

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1998). We, therefore, compare the density stratified structure with a seismic observation, assuming that former anorthosite and TTG were also subducted in addition to the MORB to the CMB as discussed above. Fig. 5 shows shear-wave velocity of minerals at 0 K from Stixrude (2007). Temperature dependence of the velocity for these minerals is not known. Moreover, effects of iron are not included in Fig. 5. In general, the addition of iron to an Mg–Fe phase reduces the velocity. Furthermore, the spin transition in the lower mantle Mg–Fe phases would change their seismic velocities (e.g., Crowhurst et al., 2008). The following discussion is, therefore, based on the relative relationship between each mineral in Fig. 5. High-temperature seismic velocity data should be investigated as well as the in-situ density data and the mineral chemistries in each rock system. The S-wave velocity of SiO2 phase is very fast except around the ferroelastic phase transition at 40–50 GPa (Fig. 5). In addition, CaPv also exhibits a high speed at the CMB conditions. Therefore, both anorthosite and TTG should have high velocities. In contrast, the mantle constituent phases, magnesium perovskite and ferropericlase (MgO in Fig. 5) show slower velocities. Furthermore, the MORB assemblage of MgPv + CaPv + St + CF would give a slower velocity than anorthosite and TTG due to the presence of MgPv by consumption of the SiO2-phase, although the velocity of CF is unknown. Therefore, anorthosite and TTG which are dominated by CaPv or SiO2 phase should be faster than the mantle and MORB rocks at the CMB conditions. Thomas et al. (2004) reported a paired D″ discontinuities beneath the Cocos plate. An important feature is that the shallower discontinuity shows a positive velocity jump while the deeper plane shows a negative jump. Hernlund et al. (2005) also presented paired D″ discontinuities beneath Eurasia and in the Caribbean region and proposed the so-called “double crossing” hypothesis (Fig. 6a, b). They considered that a geotherm crosses the PPv transition twice, which can explain the opposite velocity jumps at two discontinuities. However, if the compositional effect on the sharpness of the transition is considered (Fig. 6a), the transition seems too broad to match the observed seismic discontinuity. In particular the upper D″ discontinuity should be broad if the PPv transition is responsible for its formation (Fig. 6a). Nevertheless, the observations showed that the width of the discontinuity at the top of the D″ layer was up to 75 km (Wysession et al., 1998). We therefore, test an alternative solution, which is the bulk rock chemistry change. In Fig. 6c, we illustrate our concept of the density stratified layer structure which consists of the subducted materials. From the deeper level, the sequence is: MORB, TTG + anorthosite, and pyrolite mantle.

10.3. Comparison of the predicted density stratified layers at the CMB with seismic velocity data Seismic discontinuity can be produced by either phase transition of mineral or bulk chemistry change. This ambiguity has led to the long debate for the mantle model, as has been discussed with respect to the 410 and 660 km discontinuities (Ringwood, 1975; Liu, 1979; Bass and Anderson, 1984; Anderson and Bass, 1986; Ito and Takahashi, 1989; Nishihara and Takahashi, 2001; Murakami et al., 2007). A key aspect in these debates is about the sharpness of the discontinuity and the phase transition. The range of the PPv transition is 110–145 km wide (Ohta et al., 2008). Stixrude (1997) discussed that the equilibrium two-phase loop would be wider than the observed seismic discontinuity width. However, these estimations seem too broad to match the observed width of the D″ discontinuity of 50–75 km (Wysession et al.,

Fig. 5. Shear-wave velocity for minerals at 0 K after Stixrude (2007). CaPv and SiO2 phase have very high speeds.

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Fig. 6. The model structure of the D″ layer: (a) PPv transition model, (b) S-wave profile (Hernlund et al., 2005), and (c) density stratified model (this study). Broken lines in (a) are the post-perovskite transition in the pyrolite system (Ohta et al., 2008). Hypothetical geotherm is put there. Number in parenthesis in (c) is the zero-pressure rock density in g/cm3. The MORB density (4.29 g/cm3) is of zero-pressure with MgPv and St. At the base of the mantle, it should be larger because of the PPv and α-PbO2–SiO2 transitions (Fig. 2). In order to explain the S-wave jumps in (b), the PPv transition occurs twice at shallower and deeper levels in (a), and the bulk rock chemistry changes in (c). In the model (a), the PPv transition occurs at the wide depth range, whereas in the model (c), the bulk chemistry change produces sharp discontinuity. See text for details.

Note that TTG and anorthosite will have a similar density and therefore, they may amalgamate together. In this model, the upper D″ discontinuity can correspond to the boundary between the mantle and the TTG + anorthosite. As mentioned above, both TTG and anorthosite will have a very high S-wave velocity due to the presence of SiO2 phase or CaPv (Fig. 5). Therefore, this change in chemistry explains the positive velocity jump (Fig. 6). The lower discontinuity can be explained by the boundary between TTG + anorthosite and the MORB crust. The observed negative velocity jump may be because of a low S-wave speed of the MORB layer as compared to TTG + anorthosite due to the presence of MgPv by consumption of the SiO2 phase (Fig. 5). Thus, the main part of the D″ layer corresponds to TTG + anorthosite in the density stratified model. The bulk chemistry change easily accounts for the sharp discontinuity compared to the phase transition in the multicomponent system. One of the features of the D″ layer is the negative correlation between the S-wave and the bulk sound velocities (Su and Dziewonski, 1997; Masters et al., 2000). Karato and Karki (2001) proposed that this negative correlation was due to the Ca content anomaly, because CaPv has a high S-wave and low bulk sound velocities compared to the MgPv (Karki and Stixrude, 1999). Afterwards, the PPv transformation in MgSiO3 was found to be able to explain this anti-correlation (e.g., Iitaka et al., 2004). The density stratified model in this study can also explain the anti-correlation because anorthosite contains abundant CaPv. Another feature of the D″ layer is the presence of S-wave polarization anisotropy (VSH N VSV) (e.g., Panning and Romanowicz, 2004). The PPv model can explain this feature by the preferred orientation of PPv (Murakami et al., 2004). For the density stratified model proposed in this study, the anisotropic nature of some minerals is not known, and therefore, should be tested in the near future. Recently, Ohta et al. (2008) showed a more detailed S-wave velocity structure for the bottom of the Pacific superplume. They detected four seismic discontinuities at different depths in the mantle and ascribed the PPv transitions in the MORB and peridotite for their origins, except the deepest one which may be due to the partial melting of MORB crust. Here we try to explain the seismic structure by our stratified layer model.

Fig. 7 is an S-wave velocity structure from Ohta et al. (2008). We introduce the density stratified structure made of former subducted materials in the figure. The discontinuity (a) in Fig. 7 may correspond to the boundary between pyrolite (surrounding mantle) and slabperidotite, harzburgite. From the comparison of zero-pressure density at 300 K, harzburgite (4.10 g/cm3) is less dense than the surrounding pyrolite mantle (4.19 g/cm3). However, harzburgite could reach the bottom of the mantle because of cold subduction (Irifune and Ringwood, 1987). Once it reaches there, harzburgite can be gravitationally stabilized even after thermal equilibrium because it may have PPv which the pyrolite mantle just above may not have. The discontinuity (a) shows a negative S-wave jump. This is reasonable since harzburgite does not contain Ca-Pv with a high velocity (Fig. 5). The discontinuity (b) can be the boundary between harzburgite and TTG + anorthosite. The positive seismic speed jump is explainable by the feature that both TTG and anorthosite are dominated by SiO2 phase or CaPv which has a very high speed (Fig. 5). The discontinuity (c) may be the boundary between TTG + anorthosite and the MORB crust. As discussed above, the MORB layer has a low S-wave speed compared to TTG + anorthosite (Fig. 5). Finally, the discontinuity (d) may correspond to the boundary between the MORB crust and the partially melted layer. The partial melting can explain the negative S-wave velocity jump at the discontinuity. The partial melting of the MORB layer produces anorthositic restite and FeO-rich melt as discussed above. Anorthosite should be less dense than the MORB layer and they move upward to amalgamate with the anorthosite layer just above the MORB layer. Therefore, the anorthosite layer should grow with time. In summary, in our density stratified model, the paired D″ discontinuities is a consequence of TTG + anorthosite and MORB subduction. The single D″ discontinuity with only positive velocity jump may indicate that only TTG + anorthosite were subducted. Absence of discontinuity suggests no subducted materials, although the PPv transition in pyrolite with a wide transition interval may show some seismic anomaly. For the multiple (more than two) discontinuities, we should consider another subducted material (harzburgite) or the partial melting of the MORB crust. In order to understand the distributions of these subducted materials at the CMB, detail seismic

Fig. 7. Shear-wave structure of the base of the mantle at the Pacific low-S velocity zone (Ohta et al., 2008). The density stratification of the subducted materials is shown. The density for each rock (g/cm3) is denoted in the parenthesis. Note that harzburgite (slabperidotite) is less dense than pyrolite in zero-pressure density. However, due to its low temperature, it can be denser during the subduction. Once it reaches the bottom of the mantle, the formation of PPv may make it denser than the above pyrolite without PPv. See text for details. In our interpretations, each seismic discontinuity is due to the bulk chemistry change, instead of the PPv transitions proposed by Ohta et al. (2008). See text for details.

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observations as was made by Ohta et al. (2008) beneath the Pacific should be made in other regions as well. A seismological observation reported a low velocity structure which was never explained by the PPv transition. Wen (2001) found a low velocity zone with steeply dipping edges and rapidly varying thicknesses and geometries beneath Indian Ocean. The S-wave velocity was anomalously decreased by −2% at 200 km above the CMB to −9 to −12% at the CMB, relative to the preliminary reference Earth model. This seismic anomaly zone looked like a high land with a steep edge on the CMB. This structure cannot be explained by the temperature change along the PPv transition. Our interpretation is that this high land is a pile of subducted slabs without TTG and anorthosite. As discussed above, harzburgite would have a slower S-wave velocity than the surrounding mantle. Moreover, the partial melting of the MORB crust can explain the large S-wave decrease to −12%. Thus, the density stratified layer model can reasonably explain the multiple D″ discontinuities (Hernlund et al., 2005; Ohta et al., 2008), negative correlation between S- and bulk sound velocities in the D″ layer (Su and Dziewonski, 1997; Masters et al., 2000), and the lowvelocity high land structure (Wen, 2001). Contrary to the PPv transition, the stratified layers can produce a sharp seismic discontinuity. Considering the global material circulation through geologic time, the density structure of the subducted materials, and the sharpness of the PPv transition in the multicomponent systems, we propose that the density stratified layer model is more plausible for the D″ discontinuity rather than the phase transition model, concluding that the D″ layer is a chemically distinct region. 10.4. The anti-crust at the CMB The foregoing discussions about the global material circulation, growth history of the continental crust, high-pressure petrology, and mineral physics, all suggested that the D″ layer is a chemically distinct layer, which is the anti-crust. The seismologically observed D″ discontinuity, which shows positive velocity increases of about 3% in both P- and S-waves, may correspond to the presence of the former TTG+anorthosite rocks. The paired discontinuities may denote the top and bottom of the TTG+ anorthosite unit (Fig. 6). Therefore, the presence of the D″ discontinuity indicates the presence of TTG+anorthosite. The lower plane of the paired discontinuities indicates the presence of the MORB crust beneath the TTG+anorthosite unit. The thickness of the D″ layer represents the thickness of the subducted TTG+anorthosite. Wysession et al. (1998) mapped the D″ discontinuity by reviewing the previous studies. Although it is not conclusive, the discontinuity seems to exist ubiquitously, suggesting that the TTG + anorthosite unit is homogeneously distributed on the CMB. On the other hand, at the bottom of the D″ layer, the presence of ULVZ may not be ubiquitous. Garnero et al. (1998) mapped the presence and absence of the ULVZ at the CMB although the observation was not globally made. The largest ULVZ occurs on the bottom of the Pacific superplume. A similar zone is also recognized in the African superplume. At the bottoms of both superplumes, the existence of a denser material was suggested from seismology (Ishii and Tromp, 2004; Trampert et al., 2004). This heavy material may indicate the former subducted MORB crusts. Hence, the magma generation from the MORB crust is expected at those places. More recently, Idehara et al. (2007) suggested the existence of the ULVZs beneath Philippine–Kalimantan and East of Australia. The presence of the ULVZ may represent on-going magmatism due to the partial melting of the subducted MORB crust. Similar to the subduction zones at the surface, the ULVZ is a place where chemical differentiations are operated at the CMB. 10.5. Comparison of the D″ layer with continental crust on the Earth The continental crust at the surface is classified into the upper (felsic) and lower (mafic) units in the general sense. It grows at its

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margins by accreted arcs and material supplied by island arc magmas. It is also subjected to metasomatism by fluid phases from dehydrations of subducted slabs. Thus, the surface crust is chemically diverse and density-stratified. On the other hand, at the bottom of the mantle, there may be up to five layers. From the bottom to top these are; (1) FeO-rich silicate melt, (2) former MORB crust, (3) CaPv-enriched anorthosite, and (4) TTG crust dominated by the dense SiO2 phase, and (5) slab or mantle peridotite. Through geologic time, the subducted materials were accreted to the anticrust to form the D″ layer. The magma (1) is generated by the partial melting of subducted MORB (2). In addition to these structures, a metasomatic layer might be present, formed by the reaction between the liquid outer core and the solid mantle, even though it could be as thin as 10–100 m (Takafuji et al., 2005; Ozawa et al., 2008). At both the surface and CMB, chemically different layers are expected to be stratified according to the rock density. In addition, accretion, magma generation, and metasomatism, all are operated at the CMB similar to the tectonics on the surface. The above comparison strongly suggests that the evolution of the anti-crust (D″ layer) occurred in the same manner as that of the surface continental crust, indicating that the top and bottom of the mantle are very similar in terms of tectonics. Acknowledgement We thank two anonymous reviewers and M. Santosh for their critical comments. References Akaogi, M., Tanaka, A., Kobayashi, M., Fukushima, N., Suzuki, T., 2002. High-pressure transformations in NaAlSiO4 and thermodynamic properties of jadeite, nepheline, and calcium ferrite-type phase. Physics of the Earth and Planetary Interiors 130, 49–58. Akber-Knutson, S., Steinle-Neumann, G., Asimow, P.D., 2005. Effect of Al on the sharpness of the MgSiO3 perovskite to post-perovskite phase transition. Geophysical Research Letters 32 (L14303). doi:10.1029/2005GL023192. Anderson, D.L., Bass, J.D., 1986. Transition region of the Earth's upper mantle. Nature 320, 321–328. Armstrong, R.L., 1981. Radiogenic isotopes: the case for crustal recycling on a nearsteady-state no-continental-growth Earth. Philosophical Transactions of the Royal Society of London Series A 301, 443–472. Bass, J.D., Anderson, D.L., 1984. Composition of the upper mantle: geophysical tests of two petrological models. Geophysical Research Letters 11 (3), 237–240. Bose, K., Ganguly, J., 1995. Quartz-coesite transition revisited: reversed experimental determination at 500–1200 C and retrieved thermochemical properties. American Mineralogist 80, 231–238. Bowen, N.L., 1928. The evolution of the igneous rocks. Princeton University Press. 334 pp. Brown, G.C., 1979. The changing pattern of batholith emplacement during earth history. In: Atherton, M.P., Tarney, J. (Eds.), Origin of granite batholiths: geochemical evidence: based on a meeting of the Geochemistry Group of the Mineralogical Society. Shiva Publishing Nantwich, UK, pp. 106–115. Bullen, K.E., 1949. Compressibility-pressure hypothesis and the Earth's interior. Monthly Notices of the Royal Astronomical Society. Geophysical Supplement 5, 355–368. Chopin, C., 1984. Coesite and pure pyrope in high-grade blueschists of the western Alps: a first record and some consequences. Contributions to Mineralogy and Petrology 86 (2), 107–118. Crowhurst, J.C., Brown, J.M., Goncharov, A.F., Jacobsen, S.D., 2008. Elasticity of (Mg,Fe)O through the spin transition of iron in the lower mantle. Science 319, 451–453. Dewey, J.F., Windley, B.F., 1981. Growth and differentiation of continental crust. Philosophical Transactions of the Royal Society of London Series A 301, 189–206. Drummond, M.S., Defant, M.J., 1990. A model for trondhjemite–tonalite–dacite genesis and crustal growth via slab melting: Archean to modern comparisons. Journal of Geophysical Research 95 (B13), 21503–21521. England, P.C., Thompson, A.B., 1984. Pressure–temperature–time paths of regional metamorphism. I. Heat transfer during the evolution of regions of thickened continental crust. Journal of Petrology 25, 894–928. Fyfe, W.S., 1978. The evolution of the Earth's crust: modern plate tectonics to ancient hot spot tectonics. Chemical Geology 23, 89–114. Garnero, E.J., Helmberger, D.V., 1996. Seismic detection of a thin laterally varying boundary layer at the base of the mantle beneath the central-Pacific. Geophysical Research Letters 23 (9), 977–980. Garnero, E.J., McNamara, A.K., 2008. Structure and dynamics of Earth's lower mantle. Science 320, 626–628. Garnero, E.J., Revenaugh, J., Williams, Q., Lay, T., Kellogg, L.H., 1998. Ultralow velocity zone at the core–mantle boundary. In: Gurnis, M., Wysession, M., Knittle, E., Buffett, B. (Eds.), The core–mantle boundary region. AGU, Washington, D.C., pp. 319–334.

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