Mantle Convection: a Review

Fluid Dynamics Research 40 (2008) 379–398

Mantle convection: A review

Masaki Ogawa∗ Department of Earth Sciences and Astronomy, University of Tokyo at Komaba, Komaba 3-8-1, Meguro, Tokyo 153-8902, Japan

Received 10 November 2005; received in revised form 20 June 2007; accepted 8 September 2007 Available online 3 December 2007 Communicated by Y. Hayashi

Abstract The solid-state convection in the Earth mantle is characterized by plate tectonics, which shapes the tectonic activities of the Earth, and superplumes as broad hot regions chemically distinct from the surrounding regions in deep lower mantle. Recent numerical studies of mantle convection suggest that the rigidly moving plates occur on the Earth because the rupture strength of plate margins is sufﬁciently low, while that of plate interiors is high enough to inhibit spontaneous formation of new plate boundaries by the weight of the plates themselves. This implies that the activity of plate tectonics has ﬂuctuated much in the Earth history. Recent numerical studies also suggest that the superplumes develop owing to the chemical differentiation of the mantle by ridge magmatism. Superplumes have probably induced frequent and vigorous hot spot volcanism in the early Earth. It is now within reach to construct an integrated model for tectonic and structural evolution of the mantle in the Earth and other terrestrial planets. © 2007 The Japan Society of Fluid Mechanics and Elsevier B.V. All rights reserved.

PACS: 91.35.Lj; 91.45.Fj; 91.45.Jg; 91.45.Nc

Keywords: Mantle convection; Magmatism; Superplumes; Plate tectonics; Mantle evolution

1. Introduction

The solid-state convection in the Earth mantle is a complicated phenomenon that causes various tectonic activities, especially magmatism and plate tectonics, and makes the mantle evolve on geologic time (billions of years). When looked at from the viewpoint of ﬂuid dynamics, however, mantle convection is

∗ Tel.: +813 5454 6612; fax; +813 3465 3925. E-mail address: [email protected].

0169-5983/$32.00 © 2007 The Japan Society of Fluid Mechanics and Elsevier B.V. All rights reserved. doi:10.1016/j.fluiddyn.2007.09.001 380 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 a simple phenomenon. A typical velocity of this convective flow is only a few centimeters per year and its kinetic energy is comparable to the kinetic energy of a car running on a freeway. The kinetic energy is negligibly small compared with the rate of total potential energy release due to the convective flow, and the convective flow is well approximated as a Stokes flow. Furthermore, the only force that drives the convective flow is buoyancy force. Coriolis force and Lorentz force, which are important in understanding the convective flow in the atmosphere, ocean and the outer core of the Earth, do not play any significant role in the mantle. The complicated behavior of mantle convection all comes from the complicated nature of the convecting mantle materials. The rheology of mantle materials is not simple Newtonian, and mantle materials are a multi-component system that differentiates when magmatism occurs. Here in this review, I discuss how (a) the plate tectonics and (b) the thermal and chemical structure of mantle, two of the most important features of the Earth mantle convection, arise from this complicated nature of convecting mantle materials; for more general reviews of the studies on mantle convection, see Bercovici et al. (2000), Tackley (2000), and Schubert et al. (2001).

2. The origin of plate tectonics

The Earth surface is covered by several rigid plates, thousands of kilometers across and about a hundred kilometer in thickness, that move along the surface. The plate motion induces such tectonic activities as mountain building, earthquakes and magmatism mostly at plate margins, and the plate interiors are tectonically rather stable. (See standard text books of solid earth science (e.g., Moores et al., 1995) for the detail of plate tectonics.) This plate tectonics is a striking feature of the Earth mantle convection, not observed on other terrestrial planets (e.g., Schubert et al., 2001), and both numerical and laboratory modeling of plate tectonics has been carried out to understand why plate tectonics occurs on the Earth (Bercovici et al., 2000; Tackley, 2000). This issue on the origin of plate tectonics consists of three sub- issues; (a) why the lithosphere, a layer characterized by its high mechanical strength, develops in the uppermost coldest part of the mantle; (b) why the lithosphere is fragmented into several pieces or plates in spite of its high mechanical strength; and (c) why the plates rigidly move and the deformation of the lithosphere due to plate motion concentrates to narrow plate margins. I will discuss these sub-issues below:

2.1. The origin of the lithosphere

When the rheology of mantle materials is modeled by a Newtonian rheology, the effective viscosity rapidly increases with decreasing temperature, and this strong temperature-dependence of viscosity has been believed to be the cause of the high mechanical strength of the lithosphere (see Karato and Wu, 1993; Kohlstedt et al., 1995 and the references therein). To confirm this belief, numerical experiments and scaling analyses have been extensively carried out for thermal convection of Newtonian fluid with strongly temperature-dependent viscosity in a vessel with fixed temperature at the bottom and the surface boundaries (Christensen, 1984, 1985; Christensen and Harder, 1991; Ogawa et al., 1991; Solomatov, 1995; Kameyama and Ogawa, 2000). It is commonly assumed for simplicity in these studies that the viscosity exponentially depends on temperature as

! ! exp E(T Tb) , (1) = 0 [− − ] M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 381

b 8

c c

6 stagnant lid regime no convection

(r) d

10 4 sluggish lid regime e log

small viscosity contrast regime a a

0 2468

log10 (Rb)

Fig. 1. (a) The regime diagram for thermal convection of a Newtonian fluid with temperature-dependent viscosity in a two-dimensional rectangular box heated from the bottom boundary. Rb is the Rayleigh number based on the viscosity on the bottom boundary and r is the viscosity contrast between the top and the bottom boundaries. The regime boundaries are redrawn from Kameyama and Ogawa (2000) (solid lines). Also shown are the regime boundaries suggested in Solomatov (1995) (dashed lines). (b)–(d) Typical convective flow patterns for each of the regimes, all taken from Kameyama and Ogawa (2000) with modification; the contour lines show the temperature distribution, while the color indicates the regions where strong viscous dissipation occurs owing to the convective flow.

where E is a constant, T is temperature, Tb is the temperature at the bottom boundary, and !0 is the viscosity at T Tb. Fig. 1 taken= from Kameyama and Ogawa (2000) summarizes the outcome of these numerical experiments. The regime diagram in Fig. 1a classifies the numerically obtained convective flow patterns into three regimes on the plane of viscosity contrast between the surface and bottom boundaries r exp E(Tb Tsrfc) = [ −3 ] versus Rayleigh number based on the viscosity at the bottom boundary Rb " #g(Tb Tsrfc)d /! $, = 0 − 0 while Figs. 1b–d show examples of convective flow patterns for each of the three regimes. (Here, Tsrfc is the temperature at the surface boundary, "0 is reference density, # is thermal expansivity, g is gravitational acceleration, d is depth of the mantle, and $ is thermal diffusivity.) In the small viscosity contrast regime at low viscosity contrast r (Fig. 1d), the cold plumes growing from the surface boundary (see the arrows a in the figure) always cause significant deformations along the surface boundary as can be seen from the distribution of regions with strong viscous dissipation. There is no lithosphere that behaves as a rigid lid on the top of convecting mantle on this regime. On the stagnant lid regime at high r (Fig. 1b), in contrast, a stagnant lid of highly viscous fluid develops along the surface boundary (the arrow b) and cold plumes grow only at the base of the lid (the arrows c). It is reasonable to call the lid the “lithosphere”. We can define the lithosphere on thesluggish lid regime at intermediate r (Fig. 1c), too. There are cold 382 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 plumes that grow at the base of the lithosphere (the arrow d). In contrast to the stagnant lid regime, however, the lithosphere is involved in the convective flow. The lithosphere moves by its own thermally induced negative buoyancy and induces another kind of cold plumes whose roots penetrate upward into the lithosphere (the arrow e). These regimes are separated by bifurcations at the location shown by the solid lines in Fig. 1a(Ogawa et al., 1991; Kameyama and Ogawa, 2000). Among the three regimes, the sluggish lid regime is the most relevant to the Earth plate tectonics. The lithosphere develops and moves by its own negative buoyancy on this regime, as is the case for the Earth plates; the Earth plates are mostly driven by slab-pull force and ridge-push force, both occur as a result of the negative buoyancy of the plates themselves, and the main resistance to the plate motion occurs at plate margins (Forsyth and Uyeda, 1975; Conrad and Hager, 2001; Conrad and Lithgow-Bertelloni, 2002). I will show in Section 2.2 below that the sluggish lid regime is indeed the roots of the plate tectonics observed on the Earth. A comment on the convective regimes is necessary here, because there is some confusion on the definition of the three regimes shown in Fig. 1. Solomatov (1995) also classified the convective flow patterns into the three regimes, but the locations of his regime boundaries are different from the ones in Kameyama and Ogawa (2000) as illustrated by the dashed lines in Fig. 1a. The discrepancy arises because the regime diagram of Solomatov (1995) is drawn to derive a semi-empirical relationship between the efficiency of convective heat transport and the Rayleigh number and hence is not based on detailed observations of convective flow patterns. The “regime boundaries” shown by the dashed lines do not necessarily correspond to bifurcations of the convection and I will not use his regime diagram in the following discussion.

2.2. The plate like regime

Though the numerically modeled lithosphere shows some similarity to the lithosphere of the Earth on the sluggish lid regime as discussed above, there is one important difference between them. The deformation of the lithosphere spreads over the entire lithosphere on the sluggish lid regime of numerical models, while plate interiors are almost rigid and the deformation of the lithosphere strongly concentrates to narrow and mechanically weak plate margins on the Earth. The absence of plate margins in the models of Section 2.1 is a natural consequence of the simple Newtonian rheology assumed in Eq. (1). In a simple Newtonian lithosphere, ruptures cannot take place, and hence narrow and mechanically weak plate margins, to which the lithospheric deformation concentrates, cannot arise. It is necessary to more carefully model the rheology of mantle materials in numerical simulation of plate tectonics. On the Earth, plate margins are suggested to be newly developing within the Indian plates owing to the high stress induced by the collision of the Eurasian continent and the Indian continent (Coblentz et al., 1998). Plate margins also developed across the supercontinent Pangea 180–50 million years ago when hot uprising mantle plumes caused unusually vigorous magmatism called Large Igneous Provinces and induced high lithospheric stress (Burke and Dewey, 1973; Richards et al., 1989; Eldholm and Cofﬁn, 2000). Once formed, the lithosphere “remembers” the mechanically weak plate margins, even after the high stress induced by these agents dies away and the stress level in the lithosphere returns back to the normal. Namely, when considered as a function of stress, the effective viscosity of lithospheric materials is two-valued; the higher value occurs in the rigid plate interior, while the lower value occurs in the highly deforming plate margins. Which value the lithospheric materials choose is determined by whether or not M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 383 the materials have experienced a stress higher than their rupture strength in the past. The importance of this multi-valued nature of effective viscosity and the dependence of effective viscosity on the stress history has been emphasized in Zhong and Gurnis (1995) and Zhong et al. (1998). One way to realize the multi-valued nature of viscosity in numerical models is to introduce a “damage parameter” %, which indicates how severely the convecting mantle materials are damaged, and to let viscosity depend on % as well as T as

! ! (P ) exp E(T T0) F %/(1 %) , (2) = 0 [− − − + ] where T0 is a reference temperature, taken to be a typical temperature in the asthenosphere 1300 ◦C here, and F is a constant (Ogawa, 2003a). (In Ogawa (2003a), viscosity is also assumed to exponentially depend on pressure or depth and this dependence is included in !0(P ).) The effective viscosity of severely damaged materials with % is exp( F) times lower than that of undamaged materials with % 0. =∞ − = The damage parameter % is assumed to evolve with time as

d% &'ij (ij %/), (3) dt = ˙ − where the time derivative is Lagrangean, & is constant, ) is decay time, 'ij is stress, and (ij is strain ˙ rate (Bercovici, 1996; Auth et al., 2003). Namely, % increases in strongly deforming regions but decays with the decay time of ). When the decay time is much shorter than the turnover time of convection, the solution to Eq. (3) is well approximated by its steady solution. It can be easily shown that the viscosity calculated from the steady value of % depends on stress as illustrated in Fig. 2 (Ogawa, 2003a). (Here, the stress on the abscissa means ' √'ij 'ij .) When F 0, the rheology is simple Newtonian. As F becomes larger, the stress-dependence of≡ viscosity becomes stronger.= When F>4, the viscosity is three-valued in the stress range of ('D, 'I) illustrated in the figure; 'I and 'D correspond to the rupture strength of plate interior and plate margins, respectively, and are calculated from &, ), !0, and F (see Ogawa, 2003a). Only the solid parts of the curve “F>4” are stable and physically realizable. The regime diagram in Fig. 2b and the convective flow patterns in Fig. 2c show how the convective regimes are modified from the ones shown in Fig. 1 as F in Eq. (2) increases (Ogawa, 2003a). (The r in Fig. 2b is defined as exp E(T0 Tsrfc) and the Rayleigh number based on the viscosity at T T0 [ − ] = and % 0 is fixed.) As F increases, the sluggish lid regime shrinks and disappears at F 4. Instead, a new regime= called the plate like regime appears at F>4. On the plate like regime, the lithosphere= is fragmented into pieces or plates separated by ridges, and the plates rigidly move by their own thermally induced negative buoyancy (see the plates a and b and the curve of surface velocity in Fig. 2c). The heat flow at the surface boundary decreases with increasing distance from the ridge L as 1/√L when the plates are young as is the case for the Earth (Parker and Oldenberg, 1973). The secondary plumes grow at the base of the lithosphere as indicated in Fig. 2c (see also Richter and Parsons, 1975). Such a convective motion of plates never occurs on the stagnant lid regime at higher r, while the moving plates are not rigid at all and significantly deform by their own weight on the weak plate regime at lower r (Figs. 2b and c). As illustrated in Fig. 2b, the plate like regime occurs when

'D < 'plate (4) 384 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398

3×10-5

0.00 F = 0 surface velocity -3×10-5

F > 4 F

ln (viscosity) 0 < F < 4

1/(heat flow)2 σD σI heat flow stress 109 1 Stagnant Lid regime 0 108 -1 surface velocity

7 ridge 10 Plate Like regime plate b σD = σplate plate a secondary plume 106

3 σI = σplate 5 surface velocity

viscosity contrast (r) contrast viscosity 10 Sluggish 0 Lid Weak Plate regime 104 regime

103 0 2 4 6 8 F

Fig. 2. The assumed stress-dependence of viscosity (a), the regime diagram (b), and typical convective flow patterns in each of the regimes (c) for the numerical models of thermal convection with moving plates in Ogawa (2003a). The abscissa F in (b) indicates how strongly the viscosity depends on stress as illustrated in (a). In (c), the color shows temperature distribution in the convecting vessel, while the curves of “surface velocity”, “heat flow”, and “1/(heat flow)2” show the plots of these quantities on the surface boundary against horizontal coordinate. The unit of surface velocity is cm/yr. The figures (b) and (c) are taken from Ogawa (2003a) with modification. and

'plate < 'I, (5) where 'plate is the stress induced within the plates owing to their own weight. This stress 'plate is well approximated as " #g(T0 Tsrfc)!/2, where ! is the thickness of plates, and is about 60 MPa for old 0 − oceanic plates with ! 100 km. 'plate is certainly higher than the rupture strength at plate margins, on the order of 1 MPa, estimated≈ from the stress drop of shallow earthquakes, and inequality (4) is almost certainly satisﬁed in the Earth. Inequality (5) is also likely satisﬁed in the Earth according to the earlier estimates (Cloetingh et al., 1984; Mueller and Phillips, 1991; Erickson and Arkani-Hamed, 1993). It is most likely, therefore, that the Earth lithosphere is indeed on the plate like regime. Since the roots of the plate like regime is the sluggish lid regime (see Fig. 2b), it is safe to say that the sluggish lid regime is the roots of the plate tectonics observed on the Earth. M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 385

2.3. Implications for mantle dynamics and evolution

Inequality (5) has a profound implication for the dynamics and evolution of the Earth mantle. This inequality implies that the plates are mechanically so strong that new plate boundaries would never spontaneously develop by the weight of the plates themselves. Subduction of oceanic plates would never be spontaneously initiated, for example, at the passive margins beneath the Atlantic Ocean regardless of how old and how heavy the oceanic plates become. (Notice that the thickness of plates and hence 'plate cannot indefinitely increase with the age of the plates because of the secondary plumes beneath the plates observed in Fig. 2 (Parsons and McKenzie, 1978).) As a result, the non-subducting Atlantic-type oceanic plates and buoyant continental plates would cover the entire surface of the Earth, and plate tectonics would stop working within a few hundred million years, unless some agents that reside outside the framework of plate tectonics induce new plate margins beneath the Atlantic Ocean. Namely, inequality (5) suggests that plate tectonics is not an autonomous activity and has continued for billions of years on the Earth by the help of some agents that break the lithosphere from the outside. Two suggested examples of such agents are, as discussed above, continental collision (Coblentz et al., 1998) and hot mantle plumes (Burke and Dewey, 1973; Richards et al., 1989). Continental collision is, however, a stochastic process whose frequency may have much fluctuated in the Earth history. The activity of hot mantle plumes has also been suggested to have much fluctuated in the Earth history (Prokoph et al., 2004). It is, therefore, reasonable to expect that the configuration of plates and, more generally the activity of plate tectonics, have fluctuated much in the Earth history. The Wilson cycle of continental drifts may be understood as a part of this fluctuation. Since moving plates are the most important agents that extract heat from the Earth interior and the efficiency of heat extraction by plates critically depends on plate configuration (Grigne et al., 2005), the thermal state of the mantle may have fluctuated much with time, too. The thermal history of the Earth mantle may be more stochastic than assumed in a class of models, called parameterized convection models, that are based on an empirical and deterministic relationship between the efficiency of convective heat transport and the parameters that characterize the rheology of mantle materials (see, e.g. Schubert et al., 2001 for a review). I will return back to these issues in Sections 4 and 5 below. Inequality (4) is also important to understand why, first of all, plate tectonics takes place on the Earth but does not on other terrestrial planets, especially on Venus which is similar to the Earth in both size and bulk composition (Kaula, 1995). On the Earth, inequality (4) holds probably because the water on the Earth surface softens the materials in plate margins and makes 'D < 'plate; in the absence of water, as is the case for Venus, this inequality would be violated (Kaula, 1995; Schubert et al., 2001).

3. The origin of mantle heterogeneity and superplumes

The Earth mantle is chemically heterogeneous as has been revealed from geochemical studies of crustal materials (see a review by Hofmann, 1997). The materials in oceanic crusts generated by ridge magmatism (called Mid-Oceanic Ridge Basalts or MORB) are depleted in heat producing elements U, Th and other incompatible elements compared with the crustal materials in ocean islands generated by hot spot magmatism (called Oceanic Island Basalts or OIB). (Here, “incompatible elements” are the elements that are excluded from minerals and preferentially enter magma in partially molten rocks.) This implies that the source regions of MORB, probably the uppermost mantle, are more depleted in the incompatible 386 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 elements than the source regions of OIB, suggested to be somewhere in deep mantle, and hence that there are at least two chemical reservoirs in the mantle. The presence of a chemical reservoir undepleted in heat producing elements is also inferred from the studies of thermal history of the mantle (Kellogg et al., 1999). If the entire mantle were depleted in heat producing elements, as is the case for the uppermost mantle, the cooling rate of the mantle would have been much higher than observed for the Earth mantle. Here in this section, I discuss how the chemical heterogeneity has developed in the convecting mantle of the Earth and how the conﬁguration of the heterogeneity is controlled by mantle convection.

3.1. Seismic observations

One obstacle to the study of mantle heterogeneity has been the difficulty in constraining the configuration of the chemical heterogeneity from the geochemical observations. The geochemical observations, though do show that the mantle is chemically heterogeneous, do not strongly constrain the configuration of the heterogeneity. This obstacle has been, however, overcome by the recent development of seismic tomography. The tomographic images show that there are broad regions, called superplumes, characterized by low seismic wave velocity in the lowermost mantle beneath Africa and south Pacific (e.g., Kennett et al., 1998; Bijwaard et al., 1998; Megnin and Romanowicz, 2000; Gu et al., 2001; Grand, 2002) (see Fig. 3a). The low seismic wave velocity suggests that the superplumes are hotter than the surrounding mantle. Furthermore, there are discrepancies between the images of superplumes obtained from bulk sound velocity and those from S-wave velocity, suggesting that the superplumes are both thermal and chemical origin (Kennett et al., 1998). Kellogg et al. (1999) and Davaille (1999) compared the superplumes chemically distinct from the surrounding mantle to the undepleted reservoir of OIB source and proposed a mantle heterogeneity model. According to the model, (a) the mantle is chemically layered with a global chemical boundary at depth in the lower mantle; (b) the chemical boundary is undulating and the elevated parts of the boundary are recognized as the superplumes; (c) the upper (lower) layer, which is the source of MORB (OIB), is depleted (undepleted) in the incompatible elements; and (d) hot plumes are rising up from the superplumes to induce hot spot magmatism. (See Fig. 4a, reprinted with permission from Kellogg et al., 1999; Copyright (1999) AAAS.) Davaille (1999) emphasizes that the model can explain the cluster- ing of hot spots over the topographic highs, called superswells, which develop above the superplumes (e.g., Richards et al., 1988).

3.2. Convective stirring and mantle heterogeneity

In spite of the model proposed by Kellogg et al. (1999) and Davaille (1999), the nature of the superplumes is still an open problem. It is not clear at all how such chemically distinct superplumes have developed and survived in a convecting mantle stirred by moving plates (Gurnis and Davies, 1986). In the Earth mantle, subducted oceanic plates penetrate deep into the lower mantle and induce a mantle-wide convective circulation (Grand, 1994; van der Hilst et al., 1997). (See the images of subducted slabs in the lower mantle indicated by blue bands beneath American continents and the Alps–Himalayan regions in Fig. 3b.) The mantle-wide circulation may have efﬁciently smeared out any chemical heterogeneity in the mantle on geologic time (Gurnis and Davies, 1986). This problem has spurred detailed studies on how efﬁciently mantle convection stirs the chemically heterogeneous mantle. M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 387

Fig. 3. The tomographic images of the Earth mantle obtained from (a) S-wave analysis (taken from Gu et al., 2001) and (b) P-wave analysis (taken from Bijwaard et al., 1998). Blue (red) implies high (low) seismic wave velocity. In (a), the vertical section of tomographic image is shown along the great circle indicated by the map inside the image. In (b), the horizontal section of tomographic image is presented for the depth of 1325 km.

One outcome of the detailed studies is that such a global chemical layering as the one suggested by Kellogg et al. (1999) and Davaille (1999) survives the stirring by mantle convection for geologic time only when the lower layer is chemically denser than the upper layer by the amount comparable to or greater than the thermal density contrast that drives the convection (Richter and McKenzie, 1981; Montague and Kellogg, 2000; McNamara and Zhong, 2004, 2005). Otherwise, the chemical layering would be smeared out by mantle convection within a few billion years even when there is a signiﬁcant viscosity contrast between the two layers (Christensen, 1989; van Keken and Ballentine, 1998; Becker et al., 1999; van Keken et al., 2002). Though the chemical density contrast necessary for the preservation of chemical layering can be easily attained in the Earth mantle (e.g., Hirose et al., 2005), a problem still remains for this simple model of chemically layered mantle (Vidale et al., 2001; Tackley, 2002; van Keken et al., 2002). Since the lower layer is undepleted in heat producing elements and the only way to transport heat across the chemical boundary is thermal conduction, heat would build up in the lower layer and a temperature contrast large and sharp enough to be seismically detectable would arise across the chemical boundary (e.g., see Fig. 5 of Zhong, 2006). Such a large and sharp temperature contrast has, however, never been reported from tomographic studies of the mantle. 388 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398

Fig. 4. (a) An illustration of mantle heterogeneity model taken from Kellogg et al. (1999). The mantle is chemically layered with a chemical boundary in the lower mantle. The upper (lower) layer is depleted (enriched) in heat producing elements and other incompatible elements. (b) An illustration of mantle heterogeneity model proposed by Hofmann and White (1982) and Davies (1990). The subducted oceanic crusts are drawn by dots and short segments of lines. The ﬁgure is taken from Davies (1990).

3.3. Magmatism and mantle heterogeneity

The absence of global chemical boundary with a large temperature contrast discussed above suggests that the observed superplumes (Fig. 3) are not statically maintained by their chemical buoyancy. Instead, the superplumes are likely to be maintained by a dynamic balance between (a) some agents that generate the superplumes by chemically differentiating the mantle and (b) convective stirring that smears out the chemically distinct superplumes; such a balance allows mass exchange and the resulting convective heat exchange between the superplumes and the surrounding mantle. The only possible agent that can chemically differentiate the mantle on the global scale is magmatism and ridge magmatism is the most active among the magmatism observed on the present Earth. Hofmann and White (1982) and Davies (1990) speculate that the subducted oceanic crusts formed by ridge magmatism are chemically denser M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 389

composition temperature & magma 1800 Chemically Stratified regime basaltic crust ridge magmatism residual layer PLs8 1600 1.1 Gyr PL10 PL14

1400 upper mantle whole mantle 1.7 Gyr

average temperature [K] Thermal Convection regime 1200 4 8 12 16 20 superplume average internal heating rate H [pW/kg] 2.5 Gyr

basalt harzburgite 273 2373 [K] Case PL10; H = 10 pW/kg

Fig. 5. Numerical models of magmatism in a convecting mantle calculated at various internal heating rates, taken from Ogawa (2007) with modification; the mantle is internally heated by incompatible radioactive elements. (a) The average temperature in the upper mantle and that in the whole mantle are plotted against the assumed spatially averaged internal heating rate H. Also shown are snapshots of (b) compositional distribution and (c) temperature (color) and magma (contour lines) distributions at the elapsed time shown in the figures for Case PL10 calculated at H 10 pW/kg. In (b), the blue (red) color stands for basaltic materials (residual materials of harzburgite composition) generated by= magmatism. The contour interval in (c) is 5%. than the surrounding mantle, segregate from the convecting mantle, and stay at depth in the lower mantle for a long time, making the mantle chemically stratified as a whole in spite of the stirring caused by moving plates (Fig. 4b).1 Recent high pressure experiments on mantle materials confirm that subducted oceanic crusts of basaltic composition are always chemically denser than the surrounding mantle materials except at depths less than about 40 km and just beneath the 660 km seismic discontinuity (Irifune and Ringwood, 1993; Ono et al., 2005; Hirose et al., 1999, 2005). A series of numerical models of magmatism in convecting mantle (Ogawa and Nakamura, 1998; Ogawa, 2003b, 2007) have been developed to confirm that hot and chemically distinct superplumes can indeed develop in the way speculated in Hofmann and White (1982) and Davies (1990) (Fig. 5 taken from Ogawa, 2007). In these models, magmatism is modeled as a magma-generation by pressure release melting in the upwelling part of convecting mantle and upward migration of the generated basaltic magma: The mantle material is regarded as a mixture of basaltic component and its residue of harzburgite composition, and the magma migrates through matrix as a permeable flow driven by the buoyancy of magma itself (McKenzie, 1984). The mantle is internally heated by incompatible radioactive elements, and mantle convection occurs on the plate like regime discussed in Section 2.2, in the models of Ogawa (2007). The coupled magmatism-mantle convection system experiences a bifurcation as the spatially averaged internal heating rate H exceeds a threshold around 9 pW/kg as shown in Fig. 5 (Ogawa, 2007). On the thermal convection regime below the threshold, the mantle is rather cold, magmatism does not occur

1 Reprinted from Davies, 1990, Copyright (1990) with permission from Elsevier. 390 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 at all or is negligible, and a convective flow pattern of the plate like regime similar to the one shown in Fig. 2c takes place. On the chemically stratified regime above the threshold, in contrast, the mantle becomes hot enough to induce active ridge magmatism (Fig. 5c). The ridge magmatism chemically differentiate the mantle and induces a superplume at depth in the lower mantle: The ridge magmatism induces a thin basaltic crust and a layer of residual harzburgite indicated by blue and red, respectively, in Fig. 5b. After subduction, the basaltic crusts sediment on the bottom boundary of mantle owing to their chemically induced high density. Since the subducted basaltic crusts are enriched in incompatible radioactive elements, heat builds up in the pile of basaltic materials (Fig. 5c), and the pile becomes a superplume characterized by high temperature and high content of basaltic component. Such superplumes well develop in all of the cases with H higher than the threshold (Fig. 5a). A striking feature of the coupled magmatism-mantle convection system is that the chemically stratified regime can take place, i.e. superplumes can well develop, even below the threshold; I started a calculation from a thermo-chemical state above the threshold and reduced H to 8 pW/kg, the value at the points labeled as “PLs8” in Fig. 5a. This can happen because the rate of heat extraction from the mantle by moving plates strongly depends on the size and configuration of plates, which are not uniquely determined from such parameters as H and the Rayleigh number that characterize the mantle as discussed in Section 2.3 above. I will return to this issue in Section 4 below. Further numerical experiments in Ogawa (2007) show that superplumes well develop only on the plate like regime and do not develop well on the weak plate regime identified in Fig. 2b. This occurs because the plate motion on the weak plate regime is more chaotic than that on the plate like regime and stirs the mantle more efficiently. The strong influence of the dynamic behavior of the lithosphere may be one reason that superplumes similar to the one observed in Figs. 5b and c do not occur in some of the earlier models of magmatism in convecting mantle (e.g., Gurnis, 1986; Christensen and Hofmann, 1994; Davies, 2002; Nakagawa and Tackley, 2005): The efficiency of convective stirring in these models may be higher than that in the models shown in Fig. 5. (Notice, however, that the way mantle magmatism is modeled is also different in these earlier models. The bifurcation observed in Fig. 5a takes place only when magmatism becomes more vigorous as the mantle becomes hotter. This feature is not well captured in some of the earlier models.)

4. Superplumes in the Earth history

The numerical experiments on the plate like regime also show that the dynamics of superplumes signiﬁcantly depends on the internal heating rate and hence on the age of the Earth (Ogawa, 2007), and I present Fig. 6 to discuss this dependence. In the early Earth where the spatially averaged internal heating rate H is high (Case PL14), a part of the superplume is frequently detached and rises up to the surface as a hot plume (Fig. 6f). The hot

−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−−→ Fig. 6. The numerical models of Case PL14 (a–f) and Case PLs8 (g–l), both indicated in Fig. 5, taken from Ogawa (2007) with modiﬁcation. The average temperatures in the upper mantle TUM and whole mantle TWM (a and g), surface heat ﬂow (b and h), plate velocity (c and i), and horizontally averaged magma-eruption rate (d and j) are plotted against time. The snapshots in (e), (f), (k), and (l) are the same as Figs. 5b and c but for the indicated cases. The horizontal bars above (a) and (g) indicate the period during which plates move; the single (double) bars imply that there is one (are two) moving plate(s). The red marks beneath (b)–(d) show the period when the plume shown in (f) for 4.8 and 4.83 Gyr induces hot spot magmatism. M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 391

Case PL14; H = 14pW/kg Case PLs8; H = 8pW/kg

2000 2000 TUM T TUM

T [K] WM T [K] TWM 1000 1000

200 ] 200 ] 2 2 h [mW/m 0 h [mW/m 0

5 5 [cm/yr] [cm/yr] srfc srfc U 0 U 0

0.2 0.2 m [cm/yr] m [cm/yr] 0 0.0 0 1 2 3 4 5 6 7 0 5 10 15 20 time [Gyr] time [Gyr]

plate ridge magmatism superplume superplume 4.77 Gyr 7.83 Gyr

residual layer hot spot magmatism basaltic crust hot plume 4.80 Gyr 8.18 Gyr

plate b plate a hot spot magmatism ridge magmatism

4.83 Gyr 8.51 Gyr

plate c

5.20 Gyr 10.3 Gyr

basaltic harzburgite 273 2373 [K]

11.5 Gyr 392 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 plume induces magma at depths as great as the top of the lower mantle (Fig. 6f) and also induces vigorous magma-eruption (Fig. 6d). The plume shown in Fig. 6f for 4.8 Gyr, for example, induces the spikes in Fig. 6d during the period indicated by the red bar in the figure. The “hot spot magmatism” induces new ridges just above the superplume. Since the hot spot magmatism is so frequent (see the spikes in Figs. 6b–d), there is always a ridge above the superplume (Fig. 6f). As a consequence, plate tectonics continuously occurs throughout the calculated period of time (see the horizontal bar above Figs. 6a and c). The rather steady plate tectonics keeps the thermal and chemical state of the mantle statistically steady (Figs. 6a, e, and f). When I reduced the average internal heating rate H to the value appropriate for the present Earth (Schubert et al., 2001) from H for Case PL14, I reached a thermo-chemical state of the mantle qualitatively identical to that of Case PLs8 shown in Figs. 6g–l; the mantle is still on the chemically stratified regime, though both superplume activity and plate tectonics become more intermittent. (See Ogawa (2007) for the detail of initial condition of Case PLs8.) Heat builds up in the mantle including the superplume when the lithosphere is stagnant (Figs. 6g and i). The mantle materials just above the superplume are basally heated by the superplume and eventually rise up to the surface as plumes (Fig. 6l for 7.83–8.51 Gyr). The plumes are not so hot as the ones observed in Case PL14 and hence does not induce deep magma generation. The plumes, however, do induce hot spot magmatism and let plate tectonics start (Fig. 6l for 8.18 Gyr). The resulting ridge magmatism chemically differentiates the mantle and lets a superplume grow (Fig. 6k for 8.51 and 10.3 Gyr). The moving plates, however, efficiently cools the mantle and superplume (Figs. 6g and l for 8.51 and 10.3 Gyr), and both hot spot magmatism and ridge magmatism decline (Fig. 6j). Eventually, plate tectonics stops when the ridge is swallowed by a subduction zone and annihilates. (See Fig. 6l for 10.3 Gyr. The ridge in the figure is moving to the right because the “plate c” is stagnant. Eventually, the ridge is swallowed by the subduction zone at the top right corner of the figure.) Heat builds up in the mantle again during the subsequent time when the lithosphere is stagnant (Fig. 6g). Plate tectonics does not continue long in Case PLs8, because the activity of superplume is so low that hot spot magmatism does not induce new ridges before the existing ones annihilate. The intermittency of plate tectonics and superplume activity is the reason why the mantle stays on the chemically stratified regime in Case PLs8, despite that the assumed H is lower than the threshold discussed in Section 3.4.

5. Toward the integrated model of mantle evolution The progress in the studies of mantle dynamics reviewed above has made it within reach of numerical efforts to construct an integrated model of tectonic evolution and mantle evolution for the Earth. Plate tectonics that shapes the overall features of tectonic activity on the present Earth has been already well modeled as discussed in Section 2. The similarity between the tomographic images of the mantle (Fig. 3) and the mantle heterogeneity patterns obtained in numerical models (Figs. 5 and 6) suggests that the major agents that shape the mantle structure have been already captured in the model, too. It makes sense now to ask, on one hand, how the thermal and chemical state of the mantle has evolved under the influence of tectonic activities of the Earth and to ask, on another hand, how the mantle evolution has influenced the tectonic evolution of the Earth based on such numerical models as the ones shown in Figs. 5 and 6. Figs. 6a–f drop hints on how the Earth looked like on the early stage of its evolutionary history, when the mantle was strongly heated by radioactive elements (Ogawa, 2007). In the mantle of early Earth, hot M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 393 plumes are likely to have frequently risen up from superplumes to cause active hot spot magmatism. The active hot spot magmatism would have split the lithosphere more frequently than the plumes do in the present mantle (Burke and Dewey, 1973; Richards et al., 1989). The early Earth is, therefore, likely to have been covered with a large number of small plates in contrast to the present Earth that is covered by a limited number of large plates, thousands of kilometers across. The speculated small size and large number of plates in the early Earth are exactly what are suggested from the detailed studies of the structure of continental crusts formed in the Archean (an era from 3.8 to 2.5 Gyrs ago) and the Proterozoic (2.5–0.56 Gyrs ago). There were only fragments of continents, a few hundred kilometers in size, till middle Archean, but these fragments were assembled to form large and stable continents, thousands kilometers in size, during a period from the late Archean to the early Proterozoic (e.g., Hoffman, 1988; de Wit et al., 1992; Myers, 1993). It has been suggested from these observations that, in the Archean, (a) the entire Earth has been covered by a large number of small oceanic plates as is the case for the present west Pacific region, and (b) the island arc volcanism due to the subduction of small oceanic plates formed the fragments of continental crusts (e.g., Myers, 1993). The suggested small size and large number of plates in the Archean is consistent with the high superplume activity in a strongly heated mantle shown in Figs. 6a–f (Ogawa, 2007). The deep magma-generation observed in Fig. 6, another striking feature of plumes in a strongly heated mantle, is also consistent with the frequent komatiite magma-generation in the Archean (Arndt, 2003). The episodic superplume activity at low internal heating rate (Figs. 6g–l) is, in contrast, important in understanding the nature of tectonic activity of the Earth since the Proterozoic (Ogawa, 2007). A striking feature of the tectonic activity during that period is the occasional formation of supercontinents by continental collision and their splitting, i.e. Wilson cycle (e.g., Windley, 1995). A simple kinetic model of continental drift suggests that Wilson cycle can occur on the Earth, only when the activity of plumes that split supercontinents largely fluctuates with time (Duncan and Turcotte, 1994). Indeed, superplume activity has been suggested to be episodic (e.g., Ernst and Buchan, 2002), and the episodic nature of superplume has been suggested to be the cause of the fluctuation in tectonic activity in the Earth history (e.g., Stein and Hofmann, 1994; Davies, 1995; Condie, 1995; Yale and Carpenter, 1998). The episodic superplume activity suggested from the Wilson cycle can be well understood from the episodic superplume activity at low internal heating rate shown in Figs. 6g–l. A corollary to the episodic superplume activity is a large fluctuation in the thermal and chemical state of the Earth mantle: Figs. 6g, k, and l suggest that the thermochemical evolution of the Earth mantle has been more fluctuating than that expected from classical deterministic models of thermal history of the Earth based on the parameterized convection technique (e.g., Schubert et al., 2001) as has been already discussed in Section 2.3. A comment is necessary here on the cause of episodic activity of superplumes. The flushing event caused by the high pressure induced phase transition at the depth around 660 km (Christensen and Yuen, 1985; Machetel and Weber, 1991; Tackley et al., 1994) has been invoked as the cause of the episodic activity of superplumes (e.g., Stein and Hofmann, 1994; Davies, 1995; Condie, 1995). The phase boundary has a negative Clausius–Clapeyron slope (Ito and Takahashi, 1989), and the phase boundary is at a depth slightly greater than the average depth of the 660 km seismic boundary in cold subducting slabs. The cold subducting slabs are, therefore, less dense than the surrounding mantle just beneath the average 660 km boundary. The resulting positive buoyancy impedes slab penetration into the lower mantle. When the impedance to slab penetration is moderate, subducting slabs intermittently penetrate or “flush” into the lower mantle. This flushing event has been suggested to episodically activate superplumes in the Earth. It is, however, hard to cause flushing of mechanically strong slabs (Christensen, 1996) at the 394 M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 given Clausius–Clapeyron slope of the phase boundary (e.g., Fei et al., 2004). The tomographic images of subducting slabs indicate that the slabs do stagnate beneath several subduction zones, but that the stagnation takes place at various depths ranging from 550 to 800 km (Fukao et al., 2001), often away from the 660 km seismic discontinuity. It is unlikely that the 660 km phase boundary is relevant to the episodic activation of superplumes. It is easy to extend the mantle evolution model developed for the Earth to other terrestrial planets where the lithosphere is stagnant and plate tectonics does not take place (Venus, Mars, the moon, etc., see the review by Schubert et al., 2001). We can make the lithosphere stagnant by raising the rupture strength at plate margins 'D in Fig. 2 above the stress level induced by the weight of the lithosphere itself. The resulting model may be useful in understanding the tectonic and volcanic activities of Venus: 'D is expected to be high because of the lack of water on Venus (Kaula, 1995) (see Section 2.3 above). We may also obtain mantle evolution models for Mars and the moon, whose mantles are shallower and gravitational acceleration is smaller, by making the Rayleigh number smaller. Following the pioneering modeling of mantle convection in Venus and Mars (see the reviews of Schubert et al., 2001 for Venus and Nimmo and Tanaka, 2005 for Mars), such a systematic study may shed light on the cause of, say, the resufacing of Venus (e.g., Phillips and Hansen, 1998; Stofan et al., 2005) and the localized volcanic activity that has continued for billions of years at Tharsis and Elysium on Mars (Solomon et al., 2005) (but see the reservation below). To further proceed in the study of mantle evolution, however, the models must be extended in three directions:

(a) First, more attention must be paid to the influence of planetary formation processes on mantle evolution, especially in the study of small planets like Mars and the moon where a large portion of the surfaces are formed as early as 4 billion years ago or more (e.g., Parmentier et al., 2002; Elkins-Tanton et al., 2005). Magma-ocean that most likely develops at the time of planetary formation (e.g., Kaula, 1979; Abe and Matsui, 1986; Benz and Cameron, 1990) strongly influences the thermal and chemical state of the mantle on the earliest stage of the history of terrestrial planets. (b) Second, the formation of continental crusts, a major component of the Earth tectonic activities, must be modeled and must be implemented in the above discussed models of the Earth mantle evolution. This extension is crucial to understand the history of continental growth (e.g., Condie, 1989; Windley, 1995; Rino et al., 2004). This extension is also important to understand the Wilson cycle of continental drift, which is known to have occurred at least since Proterozoic (Windley, 1995), within the context of mantle evolution. In modeling the formation of continental crusts, much attention must be paid to the role that water plays; water strongly affects the properties of mantle and crustal materials, and this effect is crucial in the formation of continental crusts (Campbell and Taylor, 1983). (c) Third, the models must be extended to the mantle in a three-dimensional spherical shell. As discussed in Section 2.3, plate behavior is expected to be rather stochastic and fluctuating because of inequality (5), especially when the internal heating is mild (see Figs. 6g–l). Such stochastic behavior would strongly depend on the geometry of convecting vessel. The issue of the Wilson cycle discussed above, for example, can be ultimately resolved only by calculations in a three-dimensional spherical shell (see, e.g., Phillips and Bunge, 2005). The influence of plate motion on the configuration of thermochemical heterogeneity in deep mantle, including superplumes, also strongly depends on the geometry (Bunge et al., 1996; Van Keken and Zhong, 1999; Xie and Tackley, 2004; McNamara and Zhong, 2004, 2005; Huang and Davies, 2007). In spite of the enormous computational load, the pioneering M. Ogawa / Fluid Dynamics Research 40 (2008) 379–398 395

three-dimensional calculations (e.g., Baumgardner, 1985; Bercovici et al., 1989; Tackley et al., 1994) encourage us to extend the models in this direction.

These three extensions are all challenging and are highly rewarding.

Acknowledgment

This work is ﬁnancially supported by KAKENHI 16540379 by MEXT of Japan.

References

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