Accumulation of ‘anti-continent’ at the base of the mantle and its recycling in mantle plumes

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Accumulation of ‘anti-continent’ at the base of the mantle and its recycling in mantle plumes Yoshiyuki Tatsumi a,b,⇑, Toshihiro Suzuki b,c, Haruka Ozawa b, Kei Hirose b,d, Takeshi Hanyu b, Yasuo Ohishi e b

a Department of Earth and Planetary Sciences, Kobe University, Kobe 657-8501, Japan Institute for Research on Earth Evolution (IFREE), Japan Agency for Marine-Earth Science and Technology (JAMSTEC), Natsu-shima, Yokosuka 237-0061, Japan c Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Tokyo 152-8551, Japan d Earth-Life Science Institute, Tokyo Institute of Technology, Tokyo 152-8551, Japan e Japan Synchrotron Radiation Research Institute, Sayo 679-5198, Japan

Abstract The continental crust is a unique reservoir of light elements in the solid Earth; it possesses an intermediate composition and is believed to have been created principally along volcanic arcs, which are major sites of terrestrial andesitic magmatism. Mantle-derived arc magmas are, however, generally mafic or basaltic. A simple mechanism to overcome this apparent dilemma and generate andesitic melts in such a setting is through the partial remelting of an initial mafic arc crust by heat supplied from underplating basaltic magmas. An antithesis to the formation of continental crust in this way should be the production of refractory melting residue, here referred to as ‘anti-continent’. This anti-continent is likely to detach from arc crust as a result of a density inversion and descend into the upper mantle. High-pressure experiments demonstrate that sinking anti-continent is, in contrast to the subducting oceanic crust, always denser than the surrounding mantle, suggesting that it penetrates through the upper-lower mantle boundary, without stagnation, and accumulates at the base of the mantle to form a 200–400 km thick mass known as the D00 layer. Geochemical modeling provides further evidence that this accumulating anti-continent contributes to a deep-seated hotspot source. Therefore, through complementary processes, Earth creates buoyant continents and dense anti-continents at the top and the base of the mantle, respectively, and has recycled portions of anti-continent in mantle plumes. Ó 2013 Elsevier Ltd. All rights reserved.

1. INTRODUCTION The Earth’s oceanic and continental crusts are fundamentally different. Basaltic oceanic crust, which is the most abundant igneous rock at the Earth’s surface, is a direct product of mantle melting. Continental crust, on the other hand, occupies only 0.4% of the Earth’s total mass and has an intermediate (60 wt.% SiO2), ‘differentiated’ composition (e.g., Christensen and Mooney, 1995; Kelemen, 1995; ⇑ Corresponding author at: Department of Earth and Planetary

Sciences, Kobe University, Kobe 657-8501, Japan. Tel.: +81 78 803 6643; fax: + 81 78 803 5791. E-mail address: [email protected] (Y. Tatsumi).

Taylor and McLennan, 1995; Rudnick, 1995). Since such intermediate igneous rocks typify magmatism at convergent plate boundaries, arcs may thus play a central role in continental crust formation (e.g., McLennan and Taylor, 1982). There remains a significant dilemma that we are faced with and need to overcome to better understand the process of intermediate continental crust formation; that is, although mantle-derived andesitic arc magmas have been produced under unusual tectonic settings (e.g., Yogodzinski et al., 1994; Tatsumi, 2006), magmas generated in the mantle wedge in general possess basaltic compositions. If so, then the process and mechanism of differentiation from initial mafic arc crust to intermediate continental crust need to

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be understood. In order to explain the intermediate composition of the bulk continental crust, two processes are postulated to occur: (1) Differentiated magmas of intermediate composition must be generated, either by fractionation of mafic melts or by anatexis of mafic crust, and (2) a significant fraction of mafic to ultramafic crust components must be removed from the crust (e.g., Rudnick, 1995; Tatsumi, 2000). In addition to the origin of continental crust, therefore, the whereabouts and fate of these ‘complementarilycreated’ residual crustal components are intriguing in terms of both mantle dynamics and mantle evolution. This paper presents results of high-pressure experiments on the crustal residues produced complementary to the continental crust, compares the density of these components with the mantle, and discusses the fate of such crustal components in the deep mantle. 2. CONTINENT AND ANTI-CONTINENT FORMATION AT INTRA-OCEANIC ARCS

DB

20

PM

30

UMR

3. EXPERIMENTS

obs.

40

Upper Mantle

Depth (km)

10

Lower Crust

0

Middle Upper Crust Crust

High-resolution active-source seismic profiling has revealed the architecture of the crust and upper mantle of the Izu-Bonin-Mariana (IBM) subduction zone where an intra-oceanic, juvenile volcanic arc has been built (Kodaira et al., 2007; Takahashi et al., 2008; Sato et al., 2008). One distinctive structural feature of this arc is widely distributed middle crust with P-wave velocity (Vp) of 6.0–6.8 km/s (Fig. 1), a value equivalent to the average Vp of continental crust and typical of plutonic rocks with intermediate compositions (Christensen and Mooney, 1995) (CC in Table 1). This characteristic middle crust layer has also been documented in other intra-oceanic arcs, such as the TongaKermadec and Kurile arcs (Crawford et al., 2003; Nakanishi et al., 2009), leading to the challenging theory that intra-oceanic arcs are primary sites of continental crust formation (e.g., Kodaira et al., 2007; Takahashi et al., 2008;

Tatsumi et al., 2008). An additional novel feature of the sub-IBM seismic structure is the occurrence of an uppermost mantle layer (Fig. 1) with an average Vp of 7.2– 7.6 km/s (Kodaira et al., 2007; Takahashi et al., 2008), significantly lower than that of typical sub-oceanic uppermost mantle (>7.6 km/s), which is bordered by a seismic reflector located at a 40-km depth in the upper mantle (Sato et al., 2008). The observed seismic structure of the sub-IBM crust and mantle has been petrologically modeled by Tatsumi et al. (2008) as a two-stage process (Fig. 2), by which a mantlederived basalt forms the initial mafic arc crust and then basaltic underplating causes partial melting of this initial crust to produce an andesitic melt. Such an andesitic melt could solidify to form the characteristic IBM middle crust with a composition similar to that of the bulk continental crust, and the melting residue could form the uppermost mantle layer. APM in Table 1 is an andesitic melt that is produced by 30% partial melting of initial basaltic arc crusts (IAB and PIAB in Table 1), and migrates upwards where it solidifies to form the IBM middle crust. AC in Table 1 is a mixture of the melting residue of APM production from IAB (RES in Table 1) and the olivine component (CM in Table 1) accumulated via IAB formation from a primary arc basaltic magma (PIAB in Table 1) and is a probable constituent of the low-velocity uppermost mantle layer (Fig. 1). The calculated Vp of the plutonic lithology of APM and AC reproduces well the entire observed Vp profile of the crust (Fig. 1). Hence, this pair of distinctive subarc seismic layers may be regarded as continent (APM) and its complementary product or anti-continent (AC) (Fig. 1). AC–P in Table 1, which represents an endmember anti-continent composition, is the melting residue of APM production from PIAB and includes no olivine cumulate component. This composition is also used as a starting material for high-pressure experiments, as the degree of olivine fractionation during differentiation from PIAB to IAB may be variable and less than 20%. Crystallization of garnet that forms 20–30% of the restitic anti-continent layer (Tatsumi et al., 2008) is noteworthy, as this causes the density of this layer to be higher than the underlying mantle (3.33 vs. 3.24 g/cm3, respectively, at 40 km depth). This results in gravitational instability and ultimate detachment of anti-continent from the arc lithosphere and its descent into the mantle. Such a process, termed lithospheric delamination, has been seismically documented at continent–continent and arc-continent collision zones (Tsumura et al., 1999).

50 5.5 6.0 6.5 7.0 7.5 8.0

8.5

Vp (km/s) Fig. 1. Observed and calculated Vp in sub-IBM crust and mantle (batched regions and red lines, respectively) after Tatsumi et al. (2008). UMR, upper mantle reflector; DB, differentiated basalt; PM, primary basalt. Petrological model used for Vp calculation is also shown.

To investigate the behavior of the anti-continent in the mantle, high-pressure experiments were conducted on a couple of anti-continent compositions (AC and AC–P in Table 1) under upper to lower mantle conditions. 3.1. Experimental methods The starting materials with “anti-continent” compositions (AC and AC–P in Table 1) were prepared as gels

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Table 1 Compositions of arc crust and mantle components. PIAB SiO2 (wt.%) TiO2 Al2O3 FeO* MgO CaO Na2O K2O Total

47.8 0.6 14.9 10.9 14.9 9.5 1.2 0.2 100.0

CM 40.1

13.4 46.5

100.0

IAB

APM

RES

AC

AC–P

CC

MORB

Mantle

50.0 0.8 19.1 10.2 6.0 12.1 1.6 0.2 100.0

60.0 0.6 18.1 6.8 3.1 7.6 3.3 0.5 100.0

45.7 0.9 19.5 11.7 7.2 14.0 0.9 0.1 100.0

43.5 0.5 11.7 12.4 22.9 8.4 0.5 0.0 100.0

42.6 0.6 13.5 12.7 20.0 10.3 0.3 0.1 100.0

60.2 0.7 16.1 6.7 4.5 6.5 3.3 1.9 100.0

51.1 1.7 15.7 9.9 7.7 11.2 2.4 0.2 100.0

45.8 0.2 3.3 8.2 38.6 3.6 0.3 0.0 100.0

PIAB, mantle-derived primary IBM basalt; CM, cumulate from PIAB to produce IAB; IAB, representative IBM basalt produced by 20% olivine separation from PIAB; APM, andesitic partial melt produced by 30% melting of PIAB or IAB forming the middle crust of the IBM arc; RES, melting residue after extraction of APM from IAB; AC, crustal component that is a mixture of 60% RES and 40% olivine cumulate and is transferred to the mantle during APM formation, referred to as anti-continent; AC–P, anti-continent without olivine cumulate derived from PIAB. These compositios are after Tatsumi et al. (2008); CC, continental crust (Rudnick, 1995); MORB,representative mid-ocean ridge basalt (Ono et al., 2001); Mantle, representative mantle peridotite (Ono et al., 2005).

Middle Upper Crust Crust (Continent)

Initial Basaltic Arc Crust

Partially Molten Zone

Moho Basaltic Underplating Melting Residue (Anti-Continent)

Fig. 2. A model of arc crust evolution, continental crust and anticontinent formation (Tatsumi et al., 2008). Successive arc magmatism causes anatexis of the initial basaltic arc crust to create intermediate continental crust and complementary anti-continent. Transformation of mafic crustal component or the anti-continent into the mantle causes the chemical evolution of the arc crust from mafic to intermediate compositions.

and subsequently dehydrated at 1000 °C in a H2–CO2 gasmixing furnace, in which oxygen fugacity was controlled to be slightly above the iron-wu¨stite buffer. Phase relations up to 28 GPa were determined by quenching experiments using a Kawai-type multi-anvil apparatus (MAA). The starting material was enclosed in a graphite capsule and placed at the center of the MgO pressure-transmitting medium. A cylindrical LaCrO3 tube was used as a heating element. Pressures were calibrated against load (oil pressure) of the press at 1600 °C by using the phase transitions of SiO2 (coesite–stishovite; Zhang et al., 1996) and Mg2SiO4 (a-b, b-c, c-Perovskite + MgO; Morishima et al., 1994; Fei et al., 2004; Inoue et al., 2006). Temperatures were measured by a W95Re5– W74Re26 thermocouple. Modal compositions in run products were calculated by mass balance based on the compositions of the starting material and minerals, and are listed in Table 2 and Fig. 3. Since the grain sizes of minor phases were very small (62 lm) in the recovered samples, the chemical compositions of minerals, which are listed in Supplementary Table 1, were measured by a field-emission electron probe microanalyzer (JEOL JXA-8500F), which was operated at 10 kV and 12 nA. The zero-pressure volumes of the phases in the recovered specimens were determined by an

angle-dispersive X-ray diffraction measurement system with an imaging plate at SPring-8 and a micro-focused X-ray diffractometer (Table 2). In some cases, diffraction peaks from phases were overlapped, causing difficulty in determining their unit-cell volumes. Above 30 GPa, in situ high-pressure and high-temperature angle-dispersive X-ray diffraction measurements were performed using the laser-heated diamond-anvil cell (LHDAC) techniques at BL10XU, SPring-8 (Ohishi et al., 2008). The samples were mixed with gold powder or coated with gold, which served as an internal pressure standard and a laser absorber. The mixture was loaded into a hole in rhenium gasket. An area of approximate 50 lm was heated with a focused multimode continuous wave Nd:YAG laser or fiber lasers using double-side heating technique. The temperature of the sample was measured by the spectroradiometric method. A monochromatized X-ray beam was collimated to 15 lm in diameter and exposed on the center of the heated spot. X-ray diffraction spectra were obtained on a CCD detector or an imaging plate with a typical exposure time of 20 s and 3 min, respectively. The pressure was obtained from the observed unit-cell volume of gold by applying the equation of state (Fei et al., 2007). In the first run, an amorphous sample was compressed at room temperature and then heated to 1537 °C for 3 min at 30.6 GPa. After quenching to room temperature, the sample was further compressed and reheated to 1567 °C at 42.1 GPa; such a cycle of quenching-compressing-reheating was repeated also at 77.9 and 86.4 GPa (Table 2). The second run was conducted separately at 67.0 GPa and 1577 °C for 90 min. Unit-cell volumes of MgPv, CaPv, an aluminous phase with a calcium ferrite-type structure (CF), and ferropericlase (FP) were calculated based on the observed diffraction peaks (Table 2). The volumes of CF and FP were determined only in the second run due to weak intensities and peak overlapping. The recovered sample from 67.0 GPa was thinned by Ar ion-milling method using an Ion Slicer (JEOL EM-09100IS) after Tateno et al. (2009). The chemical analyses were made using a JEOL-2010 TEM

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Comp. Experimental conditions

˚ 3) Unit cell volume (A

Fraction (%) of phases

Press. (GPa) Temp. Duration Olivine b-Spinel c-Spinel MgPv Garnet cpx CaPv CF FP Fe Garnet (Z = 8) CaPv (Z = 1) MgPv (Z = 4) FP (Z = 4) (°C) (min.) Multi-Anvil AC AC AC AC ACP AC AC–P AC AC–P AC AC–P

5.0 10.0 15.0 20.0 20.0 24.0 24.0 25.0 25.0 28.0 28.0

1300 1400 1500 1600 1600 1600 1600 1650 1650 1650 1650

115 300 250 240 300 240 240 240 240 240 242

25.8 24.4         

  17.8        

   13.5       

     39.1 34.3 60.5 48.7 61.7 63.7

45.7 58.2 80.0 65.5 83.7 34.5 32.5  14.1  

27.6 17.4         

   16.2 10.2 16.6 20.0 22.7 22.9 21.5 23.2

     3.9 9.3 14.7 11.5 14.2 12.3

  2.2 4.9 8.1 5.9 3.9 2.1 2.8 2.6 

          0.8

LHDAC

30.6 42.1 67.0 77.9 86.4

1537 1567 1577 1637 1617

3*

    

    

    

+ + 66.6 + +

    

    

+ + 18.4 + +

+ + 7.4 + +

+ + 5.4 + +

+ + 2.2 + +

AC AC AC AC AC

+, Present; , absent. * See text.

*

90 * *

CF (Z = 4)

1537.3 ± 0.4 1555.8 ± 0.6 1568.6 ± 1.1 1540.1 ± 0.4 1548.2 ± 0.7 1528.7 ± 0.7 1534.8 ± 1.2 1534.0 ± 1.0

42.66 ± 0.03 41.12 ± 0.04 38.79 ± 0.01 38.01 ± 0.04 37.27 ± 0.03

154.68 ± 0.14 149.28 ± 0.16 140.54 ± 0.62 60.60 ± 0.16 198.00 ± 0.34 137.89 ± 0.09 135.47 ± 0.06

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Table 2 Experimental results.

Y. Tatsumi et al. / Geochimica et Cosmochimica Acta xxx (2014) xxx–xxx

Depth (km) 200

400

5

Depth (km)

600

800

1000

1500

2000

Mineral Proportion (vol. %)

100 olivine

β-sp

γ-sp

80

FP

FP

MgPv MgPv

60 garnet

40

CaPv

20 cpx

0 5

CaPv

CF+NAL

10

15

20

30 30

25

Pressure (GPa)

Fe

CF+NAL

50

70

90

Pressure (GPa)

Depth (km) 600

Mineral Proportion (vol. %)

100

800

FP

80

MgPv

60 garnet

40 CaPv

20

CF+NALFe

0 15

20

25

30

Pressure (GPa) Fig. 3. Mineral proportions in anti-continents, AC and AC–P. At pressures below 30 GPa and 67.0 GPa (solid circles; data are listed in Table 2), mineral proportions are calculated based on mineral assemblages and compositions, whereas at 30.6, 42.1, 77.9 and 86.4 GPa (open circles) they are deduced solely from the assemblage and assumed to be constant. cpx, clinopyroxene; sp, spinel; MgPv, magnesium-iron silicate perovskite containing aluminum; CaPv, Ca-perovskite; CF, aluminous phase containing sodium with the calcium ferrite-type structure; FP, ferropericlase. Hatched region, upper mantle.

operating at 200 kV with Oxford Inca energy-dispersive analytical system. Compositions of minerals in run products are listed in Supplementary Table 1. A typical TEM image of an AC run product recovered from 67 GPa and 1577 °C and backscattered electron images of AC and AC–P quenched from 24 GPa and 1600 °C are presented in Fig. 4. 3.2. Phase relations Garnet is a major constituent phase in AC up to 24.0 GPa, whereas by 25 GPa it is completely replaced by Mg- and Ca-perovskites (MgPv and CaPv) and aluminous phases containing sodium with the calcium ferritetype structure (CF) or the hexagonal structure (NAL) (Fig. 3). This phase transition thus takes place at 24 GPa, a pressure corresponding to the upper–lower mantle boundary at 670 km (ULMB). Garnet in AC–P is also completely replaced by MgPv and CaPv at 26 GPa (Table 2 and Fig. 3). On the other hand, in mid-ocean ridge basalt (MORB in Table 1), i.e., subducting oceanic crust, garnet is stable up to a lower-mantle pressure of 30 GPa (Irifune and Ringwood, 1993; Hirose et al., 1999; Ono et al., 2001). The lower Al2O3 contents of anti-continents (Table 1) result in crystalliza-

tion of garnet with higher majorite/pyrope component, which causes the garnet–perovskite transition to occur at a lower pressure than in MORB. Because garnet is less dense than MgPv and CaPv, major constituents of the lower mantle (Ono et al., 2005; Ohta et al., 2008), the sinking anti-continent should behave differently than oceanic crust at depths approximating the ULMB. The phase relation of AC does not change at pressures >25GPa (Fig. 3). Presence of Fe-metal >30 GPa in LHDAC is a consequence of high Fe3+ in MgPv, whereas few metal phases crystallized at 28 GPa in MAA (Fig. 3). This may be due to the difference in Al2O3 in MgPv (Frost et al., 2004; Sinmyo et al., 2011): higher in MgPv at high pressures, 12 and 9 wt.% at >30 and 28 GPa, respectively (Supplementary Table 1). However, this causes little difference in density estimates for the anti-continent composition. 3.3. Density relations Densities of the constituent minerals were calculated from their chemical compositions and volumes. The volumes of phases at high pressures were calculated based on the observed zero-pressure volumes and third-order Birch–Murnaghan equation of state:

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(a)

  @K T ðT  T 0 Þ @T P Z T  ¼ V 0 exp adT

KT ¼ KT0 þ

FP

V 0T

T0

metallic iron

MgPv Void 200 nm

(b)

CaPv

CF

FP

MgPv Gt 5 μm

(c) FP MgPv

CaPv

Gt 5 μm

a ¼ a0 þ a1 T þ a2 T 2 @K  T and a are temperature derivative of bulk modulus @T P and volumetric thermal expansion coefficient at atmospheric pressure, respectively. For calculation of metallic iron, P–V–T relation reported by Dubrovinsky et al. (2000) was used. Zero-pressure volumes listed in Table 2 were used in the present density calculations. Since CaPv cannot be quenched to ambient condition, the zero-pressure volume calculated from LHDAC results described below was adopted. As shown in Table 2, zero-pressure volumes of phases quenched from different pressure conditions were almost identical except for garnet. Hence, when zero-pressure volume cannot be determined, the result of the other experiments with the same starting material was used in the calculations. The compositions of quenched specimens of multi-anvil experiments at each pressure conditions were adopted for density calculations below 30 GPa. The compositions above 30 GPa were assumed to be the same as those determined at 67.0 GPa. The P–V–T data for MgPv and CaPv acquired by LHDAC (Table 2) were fitted to the Birch-Murnaghan equation of state, by fixing K0 = 4, a1 = a2 = 0. The ob˚ 3, tained zero-pressure volume of MgPv was 165.0 ± 0.4 A which was in good agreement with the observed zero-pressure volumes of MgPv in Multi-Anvil experiments (Table 2). The zero-pressure volume of CaPv was calculated to be ˚ 3, which is also consistent with the previously 45.2 ± 0.2 A ˚ 3 (e.g., Wang et al., 1996; Shim reported values, 45.6 A et al., 2000). The obtained values for other parameters are listed in Table 3, and are used for density calculations. Thermoelastic parameters adopted for other phases are also listed in Table 3. Calculated densities of AC and AC–P along an inferred mantle geotherm (Ono, 2008) and at a temperature 500 °C below the geotherm are listed in Table 4. 4. DISCUSSIONS 4.1. Accumulation of anti-continent at D layer

Fig. 4. TEM and BSE images of run products. (a) AC recovered from 67GPa and 1577 °C; (b) AC from 24 GPa and 1600 °C; (c) AC–P from 24 GPa and 1600 °C. Abbreviations are the same as Fig. 3.

3 P ¼ KT 2 ( 

"

 5 # V 0T 3  V " 21 #) 3 0 V 0T 3 1 þ ðK T  4Þ 4 V V 0T V

73

where KT, V0T, V and K0 T are the isothermal bulk modulus, zero-pressure volume, high-pressure volume, and pressure derivative of KT, respectively. The temperature effects on KT and V0T were applied by using following equations:

Comparison of calculated anti-continent densities with seismically inferred mantle density (Dziewonski and Anderson, 1981) (PREM in Fig. 5) demonstrates that anti-continent is always denser than the surrounding mantle, except at 24 GPa, the ULMB, where buoyancy is neutral. Previous experiments for MORB crust (Hirose et al., 1999; Ono et al., 2001) (MORB in Fig. 5) indicate that it has a significant positive buoyancy immediately below the ULMB, at depths between 670 and 750 km. Furthermore, seismic tomography (Fukao et al., 2009) has reinforced the concept that oceanic crust stagnates at the ULMB. It is uncertain whether the subducting oceanic crust is entirely trapped at the ULMB or whether it can accumulate to a thickness >100 km, at which point it becomes denser than the

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Table 3 Thermoelastic parameters of minerals used for density calculations. aT = a0 + a1T + a2T2 (K1)

Mineral

5

Olivine b-Spinel c-Spinel cpx Garnet MgPv CaPv CF NAL FP a b c d e f g h i j k l m n

9

a0  10

a1  10

a2  10

2.73a 2.28c 1.9d 2.56e 2f 3.14g 3.44g 3.21h 3.36k 1.69l

22.2 0 12 2.6 10 0 0 2.3 7 2.04

0 0 0 0 0 0 0 0.224 0 0.19

K0 T

(dKT/dT)P (GPa/K)

129b 173c 187d 125e 173f 246g 257g 169i 176k 153m

4.61b 4.8c 4.41d 5e 4f 4g 4g 6.3i 4.9k 4m

0.038a 0.027c 0.028d 0.016e 0.022f -0.020g 0.030g 0.030j 0.030k 0.027n

Liu and Li (2006). Liu et al. (2005). Meng et al. (1993). Nishihara et al. (2004). Zhao et al. (1997). Nishihara et al. (2005). This study. Assumed to be identical to CaFe2O4 (Skinner, 1966). Imada et al. (2012). Assumed to be identical to NAL. Shinmei et al. (2005). Saxena et al. (1993). Fei et al. (2007). Anderson et al. (1992).

Table 4 Results of density calculations. Pressure

Temperature*

AC

(GPa)

(°C)

GT0 °C

GT500 °C

5.0 10.0 15.0 20.0 24.0 25.0 28.0 30.6 42.1 67.0 77.9 86.4 100.0 120.0

1373 1439 1503 1566 1582 1590 1615 1637 1730 1923 2004 2066 2163 2307

3.544 3.689 3.919 4.012 4.353 4.493 4.544 4.549 4.700 4.990 5.105 5.191 5.323 5.504

3.598 3.735 3.960 4.055 4.399 4.543 4.592 4.595 4.741 5.022 5.134 5.218 5.345 5.522

*

KT,300 (GPa)

AC–P GT0 °C

GT500 °C

4.082 4.357 4.453 4.569

4.125 4.404 4.502 4.620

Along an inferred mantle geotherm (GT) after Ono (2008).

surrounding mantle, and penetrates into the lower mantle. In marked contrast, the higher density of anti-continent would enable it to founder continuously through the ULMB, and descend toward the core–mantle boundary. This characteristic behavior of the anti-continent is due to sharp transition between less dense garnet and denser MgPv/CaPv at a 24 GPa pressure, i.e., at the ULMB (Fig. 3), which may be caused by the garnet–perovskite transition at a lower pressure in the anti-continent with less

Al2O3 than in MORB (e.g., Ono et al., 2005; Ohta et al., 2008). The continental crust, and hence by implication to anticontinent, is believed to have been present from at least 4.4 Ga, the oldest known age for detrital zircons (Wilde et al., 2001). In this time how much anti-continent has accumulated at the base of the mantle? A simple calculation, based on the existing volume of continental crust (7.3  109 km3; Anderson, 1989), an average degree of melting for creating intermediate crust from a initial basaltic arc crust (20–30%) (Tatsumi et al., 2008; Tatsumi and Suzuki, 2009) and the fraction of olivine fractionated from a primary magma (20%; Tatsumi et al., 2008), provides a minimum estimate for the volume of accumulated anticontinent (2.3–3.8  1010 km3). Assuming current global rates of conveyance of continental crust materials into the mantle, via terrigenous sediment subduction and tectonic erosion (Scholl and von Huene, 2009), then this volume may be doubled and is almost equal to the total input of oceanic crust into the mantle over 4 Ga (7.2  1010 km3 based on oceanic crust production of 18 km3/y (Cogne´ and Humler, 2004)). We therefore suggest that massive foundering and accumulation of anti-continent is a major dynamic process in the Earth’s interior. More interestingly, the calculated volume of accumulated anti-continent would form a chemically distinct layer with a thickness of 200–400 km immediately above the core–mantle boundary. This is comparable to the thickness of the D00 layer, a seismically defined layer above the core–mantle boundary, which exhibits significant seismic anisotropy and a laterally

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Depth (km) 0

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6.0 Anti-continent (AC) along mantle geotherm

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Anti-continent (AC)along geotherm 500°C colder than mantle geotherm

Depth (km) 600

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4.6

4.5 4.0

Density (g/cm3 )

Density (g/cm3)

5.5

MORB PREM

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700

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AC

4.4

-P

PREM

AC

RB

4.2

MO

4.0 20

22

24

26

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Pressure (GPa)

3.0 0

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20

30

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Pressure (GPa) Fig. 5. Densities of anti-continents, MORB (Irifune and Ringwood, 1993; Hirose et al., 1999; Ono et al., 2001, 2005; Ohta et al., 2008) and seismically inferred mantle (PREM; Dziewonski and Anderson, 1981). Anti-continent densities are calculated along a mantle geotherm (Ono, 2008) based on observed mineral proportions (solid circles) and their extrapolation to 100 and 120 GPa (open circles) and at temperatures 500 °C lower than the geotherm (broken lines). PREM and anti-continent densities are extrapolated to the uppermost mantle pressure by using densities calculated for IBM (Tatsumi et al., 2008).

variable discontinuity along its upper surface (e.g., Garnero, 2000). A phase transition of perovskite, the dominant mantle silicate, to post-perovskite (Murakami et al., 2004; Oganov and Ono, 2004) has enabled the observed seismic characteristics of the D00 layer to be better understood (Iitaka et al., 2004; Tsuchiya et al., 2004; Ohta et al., 2008; Catalli et al., 2009), and has resulted in a general consensus that this layer consists of a heterogeneous mixture of MORB and peridotite components (MORB and Mantle in Table 1). Since the major phases of anti-continent in the lower mantle (MgPv, CaPv, and CF-type Al-phase) are identical to those of MORB (Ono et al., 2005; Hirose et al., 2005), the postperovskite transition may take place in anti-continent under conditions similar to those at which it takes place in MORB. Therefore accumulation of anti-continent, together with or instead of MORB, can also account for the seismic characteristics of the D00 layer. 4.2. Recycle of the anti-continent Accumulation of anti-continent should cause chemical heterogeneities in the D00 layer, not only in terms of major elements (Table 1), but also in terms of trace element compositions. Anti-continent is produced initially as a melting residue and coexists with an andesitic partial melt that migrates upwards to form the continental crust. Given trace element concentrations in the bulk continental crust (Rudnick and Gao, 2003), the concentration of an element, i, in anti-continent, C ianti-continent , can be calculated using the following equation: C ianti-continent ¼ p  C irestite ¼ p  C icontinent  Di where p, C irestite , and Di indicate the fraction of restite in anti-continent, concentration of element i in the restite,

and the bulk distribution coefficient of element i, respectively. Calculated compositions of restite and anti-continent at p = 0.6 (Tatsumi et al., 2008) are listed in Supplementary Table 2, together with continental crust, MORB and mantle compositions. Fig. 6a indicates that anti-continent is characteristically more enriched in Ba, K, Pb and Sr than MORB and mantle peridotite. Is the entire anti-continent that has been generated over Earth’s history amassing at the base of the mantle, or has part of it been recycled back towards the Earth’s surface by mantle convection? To address this important geodynamic question, isotopic characteristics of anti-continent are compared with those of the deep mantle geochemical reservoirs (HIMU, EMI and EMII in Fig. 6b) that have been proposed to explain the isotopic diversity observed for oceanic hotspot basalts of deep mantle origin (Zindler and Hart, 1986). It has been repeatedly emphasized that accumulation of sinking oceanic crust and sediments with experiencing compositional modification during subduction processes may be principal causes of the formation of HIMU and EMII mantle components (Devey et al., 1990; Chauvel et al., 1992; Hauri and Hart, 1993; Kogiso et al., 1997; Aizawa et al., 1999; Hanyu et al., 2011). In contrast, the origin of EMI is still controversial. Some workers favor the involvement of pelagic sediments (Weaver, 1991; Woodhead et al., 1993), whereas others emphasize the recycling of the continental lower crust or the metasomatized subcontinental upper mantle (Zindler and Hart, 1986; Tatsumoto and Nakamura, 1991; Tatsumi, 2000). The contribution of CO2-fluxed lower mantle melting to the EM1–HIMU reservoirs is also proposed (Collerson et al., 2010). Present-day Sr–Nd–Pb isotopic compositions of an inferred anti-continent (Fig. 6b) that has been created and deposited constantly since a given age can be calculated

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Primitive Mantle Normalized Element Concentration

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(a) Continental Crust

10

Oceanic Crust

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0.1 Rb Ba Th U Nb K La Pb Sr Nd Zr Sm Y Yb

0.5132

(b)

DMM HIMU

143

BSE

Mixing v

nt asce ia

Nd/144Nd

0.5128 0.5124

EMII

EMI

0.5116

Continental Crust

3.0

Accum

ulated A

0.5112 0.700 22

2.0

1.0 Ga

nti-Con

0.704

tinent

0.708

0.712

0.716

HIMU

18

206

Pb/204Pb

20 EMII

Continental Crust

BSE Mixing via ascent DMM 1.0 Ga

EMI

16

14 0.700

o

3.0 4.0

t

en ntin

2.0

ted

C nti-

A

ula

um Acc

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0.708 87

0.712

not need to assume the mantle compositions and complex behavior of element transport during hydration and dehydration at ridges and subduction zones, respectively. A mantle plume initially composed of anti-continent having these isotopic ratios varies compositionally in its ascent as it entrains and mixes with ambient mantle materials. Results from modeling this mixing process for a representative mantle composition are shown in Fig. 6b. Although strongly enriched Sr–Nd isotopic signatures of deep mantle reservoirs EMI and EMII can be explained by the contribution of anti-continent, a rather high Pb isotopic ratio for EMII is inconsistent with the involvement of anti-continent in its genesis (Fig. 6b). On the other hand, anti-continent that has been stored and isotopically evolved for 3.0– 3.5 Ga can be an endmember component of EMI (Fig. 6b), indicating that anti-continent has been recycled in mantle plumes. 5. CONCLUSIONS

0.5120 4.0

9

0.716

86

Sr/ Sr

Fig. 6. Geochemical characteristics of anti-continent. a. Primitive mantle-normalized (Sun and McDonough, 1989) trace element compositions of continental crust (Rudnick and Gao, 2003), MORB (Sun and McDonough, 1989) and anti-continent inferred from continental crust compositions, proportions of residual phases after 20–30% batch melting (Tatsumi et al., 2008; Tatsumi and Suzuki, 2009) and the mineral-melt partition coefficients (Tatsumi, 2000) are listed in Supplementary Table 2. The fraction of olivine cumulate in the anti-continent ranges from 0 (restite only) to 0.4 (Tatsumi et al., 2008); b. Sr–Nd–Pb isotopic compositions of anti-continent stored for a given duration and deep mantle geochemical reservoirs (HIMU, high-l; EMI, enriched mantle I; EMII, enriched mantle II; BSE, Bulk Silicate Earth, a representative mantle).

on the basis of representative continental crust compositions (Rb/Sr = 0.153, Sm/Nd = 0.195, U/Pb = 0.118, 87 Sr/86Sr = 0.7155, 143Nd/144Nd = 0.51155 and 206Pb/204Pb = 18.5; Alle`gre and Lewin, 1989; Rudnick and Gao, 2003) and parent-daughter element concentrations in anticontinent (Supplementary Table 2). This method is better than the previous approach (Tatsumi, 2000) in that it does

Subduction zones have been called ‘factories’, where raw materials such as oceanic crust and sediments, together with mantle wedge materials, are processed to produce arc magmas and their final solidification product, continental crust (Tatsumi, 2005). As a consequence of this processing, the ‘factory’ emits waste products, such as chemically modified oceanic crust and sediments that are residues after certain elements have been extracted by fluids and/or melts during magma production. These waste products are then recycled, appearing as HIMU and EMII components in hotspot sources (Tatsumi, 2005; Hanyu et al., 2011). In addition to this, subduction factories have produced huge volumes of anti-continent in a process complementary to continental crust formation. This waste product sinks from the subduction factory and has accumulated at the base of the mantle for much of Earth’s history, and traces of it reappear in mantle plumes bearing the EMI signature. ACKNOWLEDGEMENTS We thank R. Fiske and A. Nichols for discussion and reading of the manuscript and three reviewers for constructive comments. Y.T. and co-workers are funded by Japan Society for the Promotion of Science (19GS0211).

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