Thermal regime of the Southwest Japan subduction zone: effects of age history of the subducting plate

TECTONOPHYSICS ELSEVIER

Tectonophysics 248 (1995) 53-69

Thermal regime of the Southwest Japan subduction zone: effects of age history of the subducting plate Kelin Wang a,., Roy D. Hyndman a, Makoto Yamano b a Pacific

Geoscience Centre, Geological Survey of Canada, Sidney, B.C. V8L 4B2, Canada b Earthquake Research Institute, University of Tokyo, Tokyo 113, Japan Received 19 August 1994; accepted 6 March 1995

Abstract

During a mid-Miocene (at about 15 Ma) tectonic reorganization, an ocean spreading ridge perpendicular to the strike of the Nankai Trough subduction zone stopped spreading. Since that time there has been subduction of the cooling fossil spreading ridge. In the present study, we have developed a time-dependent thermal subduction model for this region using the finite-element method. There are good seismological and geological constraints for the plate geometry and for the subduction history of the past 15 Ma. Boundary conditions are specified so that the oceanic plate subducting at the Nankai Trough becomes cooler as it gets older. The model results agree with the present heat-flow trend that decreases landward, and explain the paleothermal regime of high mid-Miocene temperatures inferred from land geological studies. The tectonics of the region prior to 15 Ma has some uncertainties. Assuming subduction of an active spreading ridge or an 80-Ma-old lithosphere prior to 15 Ma give the same results for the present thermal regime of the seaward portion of the forearc but different results for the most landward region (> 250 km). The thermal history of the forearc since 15 Ma can be summarized as a rapid warming period as a consequence of ridge subduction, followed by a cooling trend to the present as a result of the aging of the subducting plate. The results illustrate the thermal consequences of one type of ridge subduction. They also demonstrate that the thermal regime of a subduction zone depends critically on the age history of the subdueting oceanic lithosphere, especially if it is young, as well as such parameters as the subducting plate dip angle and thickness of insulating sediments on the incoming oceanic crust. This dependence is especially important when the thermal regime is used to constrain the seismogenic behaviour of the subduction thrust fault.

1. Introduction

At the Nankai Trough of southwestern Japan, the Philippine Sea plate is subducting beneath the Eurasia plate approximately orthogonal to the margin (Fig. 1). Although the regional tecton-

* Corresponding author

ics to the northeast -which involves the interaction of these two plates at the Pacific plate and a t r e n c h - t r e n c h - t r e n c h triple junction -is extremely complex, the present tectonic setting of the Nankai Trough is relatively simple and well constrained. A n important feature in the offshore Shikoku basin is a fossil spreading ridge striking roughly in the direction of plate convergence (Fig. 1). The spreading ridge became inactive at

0040-1951/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSDI 0 0 4 0 - 1 9 5 1 ( 9 5 ) 0 0 0 2 8 - 3

54

Kelin Wang et al./ Tectonophysics 248 (1995) 53-69

Fig. 1. Map showing the regional tectonics of the Southwest Japan subduction zone (after Hibbard and Karig, 1990b) and location of the model profile.

15-12 Ma (Chamot-Rooke et al., 1987; Okino et al., 1994). As the ridge and the adjacent oceanic crust have been subducted, their age at the trench became older and hence the subducting slab became cooler. The transient thermal state of the subduction zone associated with the aging of the subducting oceanic plate is the focus of this modelling study. Steady-state thermal models have been developed for the Nankai Trough subduction zone (Honda and Uyeda, 1983; Kinoshita and Yamano, 1995; Furukawa, 1995). In these models the age history of the subducting plate was not considered. The characteristic time constant of a typical 40-km-thick lithosphere is about 16 Ma. Since the Shikoku basin fossil ridge stopped spreading only recently, the cooling of the incoming plate makes significant contributions to the present thermal regime of the subduction zone. Our work builds on these previous models but adds the important aspect of the increasing age of the subducting plate. Several issues are to be considered about the thermal regime of the Nankai Trough subduction zone: (1) The observed heat-flow trend decreases landward from about 130 m W m -2 at the trench

axis (the deformation front of the accretionary prism) to about 60 m W m -2 at a distance of 250 km. Can this large-scale spatial variation be explained in accordance with the recent subduction history? (2) Ashi and Taira (1993) reported heat flows inferred from the depth to a bottom-simulating reflector (BSR) that decrease from 130 m W m -2 at the trench axis to 40 m W m -2 in a distance only 60 km landward. Can this steep heat-flow trend on the continental slope be explained in a thermal subduction model? (3) Since the current phase of subduction began only at about 15 Ma, its thermal effects affect only a limited region of the forearc. At large distances landward from the trench, the continental crust still shows the thermal imprints of the previous subduction history. In what part of the overriding plate does the thermal regime depend only on the subduction history of the past 15 Ma, and in what part does it depend on the previous history? (4) How important is frictional heating along the subduction thrust fault? This translates to the problem of the magnitude of average shear stress on the fault. (5) The Nankai Trough experienced great subduction earthquakes ( M > 8) averaging every 180 years [with a record for the past 1300 years (Ando, 1975)]. The most recent were the 1944 Tonankai and the 1946 Nankaido ( M = 8.2) subduction thrust events, and a number of earlier earthquakes of the same type. With extensive heat-flow data, good structural control, and relatively well constrained recent subduction history, the Nankai Trough is one of the best locations to study the relation between subduction earthquakes and the geothermal regime. Thermal constraints to the seismogenic zone of the subduction thrust fault based on our thermal models are presented in a separate paper (Hyndman et al., 1995). We give detailed description of the thermal modelling in this paper. The thermal consequences of the subduction of an active spreading ridge have been investigated elsewhere (e.g., DeLong et al., 1979; Cande et al., 1987). The Nankai Trough margin, where the subducting ridge became inactive recently,

Kefin Wang et al. / Tectonophysics 248 (1995) 53-69

represents a special type of ridge subduction. We developed a 2-D finite-element model across the margin aligned along the present position of the fossil ridge, extending landward from the deformation front or "trench" (Fig. 1). The increasing age of the incoming plate is accounted for by applying a time-dependent boundary condition at the seaward end of the model. The thermal model is compared with surface heat-flow measurements that delineate the present near-surface thermal regime and with geological evidence that indicates past thermal events in the subduction zone forearc. There is some uncertainty in the direction of the motion of the Philippine Sea plate relative to the Eurasia plate (Jolivet, 1987; Otsuki, 1990), and thus in the location of the fossil ridge along the margin, in the past 15 Ma. Therefore, the model profile should be regarded as an average along a corridor some 100 km wide. It would be interesting to construct model sections also along some profiles away from the fossil ridge, such as 150 km to the southwest where the subducting plate is now about 22 Ma old, but there are no adequate structural and thermal constraints to distinguish that region from our model profile. In order to account for an older subducting plate and some uncertainties in the time when the ridge stopped spreading, we run our model from 0 to 20 Ma, instead of 15 Ma. 0 Ma corresponds to the time when the Shikoku basin ceased spreading. In the following, we first summarize the late Cenozoic thermal history as suggested by forearc geology and the available heat-flow data, then describe the numerical model and present model results.

2. Present and late Cenozoic thermal regime

2.1. Geologically inferred thermal history Marine magnetic anomalies and plate reconstructions indicate that the Shikoku basin, a backarc basin of the Izu-Bonin subduction zone to the east, stopped spreading at 15 Ma (Okino et al., 1994). Since that time, subduction of the Philippine Sea plate has been to the northwest

55

along the strike of the ridge, approximately perpendicular to the trench (Taira et al., 1992). The subduction of an active spreading ridge at 15 Ma is evidenced by a wide range of data (Underwood et al., 1993) including a regional cusp-like flexing of structural trends in the older (CretaceousEarly Miocene) Shimanto accretionary prism which now constitutes the seaward half of the Shikoku Island, thermal alteration of the Shimanto Belt and related forearc-basin deposits, near-trench magmatism and pervasive faulting (Hibbard and Karig, 1990a). The age of the subducting plate at the Nankai Trough has increased from near 0 to the present 15 Ma. Therefore, the thermal regime off central Shikoku Island is characterized by the cooling of the subducting Philippine Sea plate at the trench over the past 15 Ma. The present age of the oceanic plate along the trench increases from 15 to about 25 Ma with distance to the northeast and southwest away from the fossil ridge. Chamot-Rooke et al. (1987) proposed that there might be active N-S-directed spreading of the Shikoku basin at 14-12 Ma. Their main argument is that in the northern central part of the basin there are NNW-trending basement faults formed in this period that were thought to be transform faults. Okino et al. (1994) do not think that these faults are transform faults because there are no distinguishable offsets of magnetic lineations. Okino et al. (1994) also pointed out that new SeaBeam and seismic reflection surveys revealed intense basement deformation including many basement faults and NNW-trending grabens caused by NNW-SSE-directed compression. They attributed this compression to the collision of southwestern Japan and the Shikoku basin. Hibbard and Karig (1990a,b) also considered the spreading to have stopped at 15 Ma. Study of the oldest sediment at ODP site 808 suggests that the basement rock is 15 Ma old (Olafsson, 1993). In our modelling, we accepted the more popular view of Okino et al. (1994) and Hibbard and Karig (1990a,b) that spreading stopped at 15 Ma. However, we also ran a test case in which spreading stopped at 12 Ma, so that the thermal effects of a later phase of the spreading or similar thermal events can be evaluated.

Kefin Wang et aL / Tectonophysics 248 (1995) 53-69

56

The velocity of the Philippine Sea plate relative to the Eurasia plate is not well constrained because of the lack of an active spreading ridge boundary for the former. Published estimates of the rate range from 17 to 57 mma -1, with the most recent estimates being 43 to 46 m m a (summary by Seno et al., 1993). We take the most often cited value of 40 mm a-] in our modelling but study the effects of assuming other values,

active Shikoku basin spreading ridge. At 15 Ma, the same time as the ridge stopped spreading, the ridge collided with the Shimanto accretionary prism and initiated the present phase of subduction at the Nankai Trough. Before the collision of the ridge with the forearc, the age of the subducting plate would be about 80 Ma. Byrne and Ditullio (1992) has argued that the EoceneOligocene history of the Shimanto Belt was affected by subduction of young lithosphere generated by spreading along the Pacific Kula ridge system, which died out at 42 Ma. Other tectonic models for this region, however, assume that the active ridge had been subducted for some time prior to 15 Ma (e.g., Seno and Maruyama, 1984). The ambiguity of the situation causes difficulties for a thermal model. However, we examine the effects of different scenarios on the present thermal regime by using different temperature fields at the beginning of the model run, that is, at

tOO.

Whereas there is general agreement on the plate regime since 15 Ma before present, different models exist for the earlier period. Hibbard and Karig (1990b) pointed out that the structural and magmatic effects of the ridge subduction at 15 Ma were very short-lived. They proposed that an extension of the Pacific plate was subducting beneath the Eurasia plate at the location of the present southwestern Japan margin prior to 15 Ma, separating the Eurasian plate and the then

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Kelin Wang et aL / Tectonophysics 248 (1995) 53-69

heat flow is about 120-130 mWm -2. Using the theory of Hutch•son (1985) and the Shikoku basin sedimentation history inferred from ocean drilling cores (Taira et al., 1993), the model of a cooling lithosphere of a 15-Ma age with sedimentation predicts a heat flow of 90-100 mWm -2, and that of a 12-Ma age predicts a value of about 120 mWm -2. The observed high heat flows were cited by Chamot-Rooke et al. (1987) as supporting evidence for a later phase of spreading. However, the high values are generally considered to be related to hydrothermal circulation in the oceanic crust and the accretionary prism (Yamano et al., 1992). Detailed definition of the heat-flow distribution in the marine area is given by the depth to the thermally controlled gas hydrate bottom-

about 15 Ma. The model results may be helpful in the search for thermal data that can constrain the tectonics prior to 15 Ma. 2.2. Thermal data

Extensive heat flux data are available across the continental margin of southwestern Japan, both offshore and onshore. The offshore marine probe data have been summarized by Yamano et al. (1984), Kinoshita and Yamano (1986) and Yamano et al. (1992). The considerable variability is attributed to crustal hydrothermal circulation. There are some scattered probe values on the upper continental slope that are probably unreliable because of bottom water temperature transients. At the deformation front, the observed

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58

Kelin Wang et al. / Tectonophysics 248 (199.5) 53-09

simulating reflector (BSR) that is widespread on the continental slope (Yamano et al., 1982). A study employing 47 seismic lines along 150 km of the margin off Shikoku Island (Ashi, 1991; Ashi and Taira, 1993) shows a steady decrease in heat flow landward (Fig. 2), similar to that observed on other large accretionary prisms where young oceanic lithosphere is being subducted beneath (e.g., Davis et al., 1990; Hyndman et al., 1993). However, the particularly rapid landward decrease to about 45 m W m -2 in a distance as small as 60 km from the base of the continental slope is puzzling and will be discussed later. The BSR data are distributed parallel to the margin over subducting oceanic crust ranging in age from about 23 to 15 Ma. An important feature of the BSR heat-flow pattern parallel to the margin is that for the middle to upper continental slope (sediment thicknesses greater than 3000 m) the heat flow decreases away from the location where the fossil spreading ridge meets the margin off Shikoku Island (Fig. 3). The heat flow from land boreholes has been summarized by Li et al. (1989) and Yamano (1995). One group of sites in the eastern part of the Shikoku Island which has particularly low heat flow has a suspected sedimentation or groundwater effect, so has been excluded. There is considerable scatter, from 32 to 89 m W m -" with a mean of 57 + 16 m W m 2. The scatter is undoubtedly, in part, due to local effects such as groundwater flow and, in part, due to real larger-scale variations in conductive heat flow. The larger-scale variations are probably mainly from the distribution of crustal radioactive heat generation. Radioactive heat generation is an important factor in subduction zone thermal models, especially for the landward portion of the profiles. Neglect of continental-crust or island-arc radioactive heat generation has resulted in an incorrect inference of frictional heating in some previous subduction-zone thermal models (e.g., discussion by McCaffrey, 1993). For the seaward parts of the profiles the small heat generation in the oceanic crust may be disregarded. The heat generation of the incoming sediment section at the toe of the prism as determined from Ocean Drilling Pro-

gram K, Th and U data (Taira et al., 1993; Hill et al., 1993), is 1 . 9 / , W m -3 for an average density of 2.4 gem ~. Assuming that this sediment represents an average for the nearby continental area from where it was eroded, the heat generation for a crustal density of 2.7 g cm 3 is 2.2 # W i n 2. This heat generation is in good agreement with that of 2.4 / , W m -~ from the average K, Th and U values for plutonic rocks from the " O u t e r Z o n e " of southwestern Japan (Kanaya and Ishihara, 1975). The land sample distribution is limited, but there is no clear areal pattern.

3. Model description 3. 1. Numerical method

We solve the heat transfer equation to obtain temperature (T) at distances (x) from the deformation front and depths (z) below sealevel as a function of time (t): pc-- = V' (kVf) -pcv" VT+ Q i~l

(1)

where k is thermal conductivity, pc is volumetric heat capacity and v is velocity. The continental plate is fixed and the velocity of the subducting slab is the plate convergence rate. Q is a heat source term that arises from the radiogenic heat production of rocks and from frictional heating. In the case of frictional heating, Q is generally temperature dependent as is discussed later, which makes the equation nonlinear. A 2-D isoparametric finite-element model is used to solve the equation with prescribed boundary and initial conditions. Each element has six to eight nodes on its boundaries. A uniform thermal conductivity, heat capacity, heat source and motion velocity is assigned to each element, with the value evaluated at the element centre. Temperature is allowed to vary quadratically within each element. The origin of time (t = 0 ) in the model is fixed at 15 Ma before present, so that t = 15 Ma corresponds to the present. An unconditionally stable fully implicit finite difference scheme is employed for time stepping. The time step lengths vary from 0.01

Kelin Wang et aL / Tectonophysics 248 (1995) 53-69

m.y. at the beginning of the calculation to 0.5 m.y. at the end (20 Ma). The modelling technique is a transient version of the numerical method used by Hyndman and Wang (1993) and Wang et al. (1995). Previous 2-D thermal models for the Nankai Trough subduction zone assumed a simple straight plate interface. With the much increased knowledge of the structure of this subduction zone over the past decade, and the flexibility of finite-element discretization, a more realistic plate geometry can be used. Within 40 km landward of the deformation front, the top of the oceanic crust is well controlled by seismic reflection data (e.g., Ashi, 1991; Moore et al., 1990). Further landward, the control on the top of the subducting oceanic plate is from Benioff-Wadati seismicity within the downgoing plate (Hirahara, 1981; Mizoue et al., 1983; Okano et al., 1985; Ito, 1990). A number of authors have contoured the depths to the Philippine Sea plate as inferred from earthquake data (e.g., Yoshika, 1991; Sato and Matsu'ura, 1992). Based on the assumption that the plate interface is about 5 km above the top of the Benioff-Wadati earthquakes and with reference to the contour maps provided by other authors, we define the plate interface geometry for our model profile. The finite-element grid (Fig. 4) extends from 5 km seaward of the deformation front to 600 km landward. However, igneous upwelling and asthenospheric flow processes are probably important beyond 300 km (e.g., Honda, 1985) (although there is no well defined volcanic arc), so only the model thermal regime to about this distance is meaningful. The subducting slab enters the model domain from the seaward boundary and exits from the land-

0

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Distance (km) Fig. 4. Part of the finite-element grid used for the calculation. Boundary conditions are described in the text.

ward boundary. The bottom of the grid is 200 km below, and follows the geometry of the upper surface of the subducting plate. The measured values of thermal properties of near-surface rocks are quite variable, depending on the local rock types. We take a uniform thermal conductivity of 2.5 W m-1 K-~ and thermal capacity of 2.5 MJm - 3 K -1 for the continental side of the model, and 2.9 W m 1K-1 and 3.3 M J m - 3 K -t, respectively, for the oceanic lithosphere. Detailed modelling using fine conductivity structures is not yet warranted for the scale of our study. The thermal conductivity of the frontal part of the accretionary prism increases landward systematically because of decreasing water contents; therefore, we let the conductivity of the wedge sediment in the model increase from 1.5 W m -~ K -1 at the deformation front to 2.5 W m - t K - l at 60 kin. For radioactive heat production of rocks, we employ average values of 1.9 p~Wm-3 for the upper 10 km of the prism and continental crust and 0.4 /zWm -3 for the lower crust from 10 to 20 km. This roughly corresponds to an exponential decrease in heat production.

Table 1 Thermal properties for the model cross section Geological unit

Thermal conductivity (Wm 1 K - l )

Heat production (/zWm -3)

Thermal capacity (MJm 3 K - I )

Accretionary prism Continental crust (0-10 km) Continental crust (10-20 km) Continental mantle Oceanic lithosphere

1.5-2.5 a 2.5 2.5 2.5 2.9

1.9

2.5 2.5 2.5 2.5 3.3

a Increase linearly with distance from the deformation front (see text).

1.9 0.4 0.0 0.0

60

Kefin Wang et a L / Tectonophysics 248 (1995) 53-69

These values are averages of the observed values (summarized by Hyndman et al., 1995). Similar values have been used by Furukawa (1995). Since radiogenic heating contributes only to the steady-state component of the solution and its distribution is largely one-dimensional, the effect on the surface heat flow of varying these values can be readily evaluated without numerical modelling. Frictional or strain heating along the plate interface involves nonlinearity, and will be dealt with separately in a later section. The thermal properties used for the modelling are summarized in Table 1.

3.2. Boundary conditions The landward boundary is assigned zero horizontal heat flux. Because of this simple boundary condition and the uncertainties in the slab geometry as well as in the physical processes at great depths, solutions beyond 300 km distance from the deformation front are not physically meaningful. The upper and lower boundaries are maintained at constant temperatures of 0 and 1400 ° C, respectively. As will be seen later, the temperature in the lower part of the model from about 5 Ma after the beginning of the calculation depends mainly on the boundary condition of the seaward vertical boundary; therefore, the temperature used for the lower boundary is not important. A critical factor in this model is the time-dependent boundary condition at the seaward boundary. We assume that the temperature-depth profile in the oceanic plate beneath the offshore Shikoku basin can be described by the 1-D half-space lithospheric cooling model applicable to young oceanic lithosphere (Davis and Lister, 1974). This t e m p e r a t u r e - d e p t h profile as a function of time is used to prescribe the boundary condition for the seaward boundary. At the beginning of the model run, the temperature at the boundary is constant with depth at 1400 ° C, representing the subduction of an active ridge or zero age oceanic lithosphere. For each successive time step, the vertical temperature distribution on this boundary is calculated from the half-space model for that particular plate age (Fig. 5). The

Temperature (°C) 0

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advective heat brought into the subduction zone by the oceanic plate thus decreases with increasing time. The simple analytical solution for a cooling half-space assumes a uniform thermal conductivity (Davis and Lister, 1974). In order to account for the much lower conductivity ( ~ 1.5 W m -1 K 1) of the 1.2-km-thick sediment layer, we modify the half-space solution by scaling the thermal gradient of this section obtained from the half-space model inversely proportional to the conductivity. The calculated surface heat flow remains unchanged. This practice does not take into account the transient effects of sedimentation on the thermal regime properly which we have ignored in this study. As stated previously, sedimentation on the seafloor off Shikoku is expected to reduce the heat flow by about 20%, but the observed heat flow is consistent with values without sedimentation. The present heat flow predicted by our simply modified half-space solution is 123 m W m -2, close to the observed value. The lack of evidence for a sedimentation effect requires further study. In any case, we expect that ignoring the sediment will have an effect only in the thermal regime near the frontal area of the accretionary prism.

Kefin Wang et aL / Tectonophysics 248 (1995) 53-69

4. M o d e l results

4.1. The preferred model A preferred model is calculated with the thermal properties and the boundary conditions described in the preceding section. The initial condition (temperature distribution at t = 0) of this model is the temperature field of a half-space cooled from 1400°C for 80 Ma (i.e., no subduction). This simple initial condition for the preferred model is chosen considering the lack of conclusive information on the pre-15 Ma tectonic regime. This initial condition leads to heat-flow values about 60 mWm -2 at t = 15 Ma (present) in the region 200 to 300 km landward of the deformation front, consistent with present heatflow observations available in that region. Initial conditions based on different tectonic models prior to 15 Ma will be tested with the results discussed in a later section. The modern accretionary wedge (the most seaward 100 km of the overriding plate) has formed since 15 Ma (Taira, 1985, 1988). To account grossly for the effects of the growth of the accretionary wedge, we set the initial temperatures to 1400°C for all grid nodes from x = 100 km seaward. However, as will be shown later, the position of the landward end of the ridge at t = 0 affects the model results significantly only for a short time ( < 5 Ma). The oceanic lithosphere and all the material below it subducts together at the plate convergence velocity of 40 mma-1. The effect of allowing different thickness of the subducting material will also be discussed later. There is no frictional heating in the preferred model. Conclusions on the width of the seismogenic zone of the subduction thrust fault by Hyndman et al. (1995) are based on this model. The calculated surface heat fluxes at four time steps together with the present heat-flow data for the Shikoku (Nankaido) margin are shown in Fig. 6a. The temperature fields at the same time steps are shown in Fig. 7. The 15-Ma solution represents the present thermal regime. The transient state of the thermal regime is well characterized by the temperature distribution as a function of distance along the plate interface (Fig. 6b). Along the plate interface, the thermal pulse induced by

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Fig. 6. (a) Surface h e a t flow c a l c u l a t e d at four t i m e steps for the p r e f e r r e d model. 15 M a a p p r o x i m a t e l y c o r r e s p o n d s to the present. Symbols are m e a s u r e d heat-flow v a l u e s p r o j e c t e d to the m o d e l profile. O are m a r i n e p r o b e m e a s u r e m e n t s , • are gas h y d r a t e B S R values, • and [] w i t h e r r o r bars are land b o r e h o l e m e a s u r e m e n t s from N a n k a i d o a n d T o n a n k a i regions, respectively. (b) T e m p e r a t u r e s a l o n g the p l a t e i n t e r f a c e for the s a m e four t i m e steps. In b o t h (a) a n d (b), the v a l u e s a s s o c i a t e d with the s t e a d y - s t a t e s u b d u c t i o n of a 15-Ma-old slab are s h o w n as d o t t e d lines.

the subduction of the active ridge at t = 0 disperses landward. At 5 Ma, the point on the subducting plate that was initially at distance x = 0 when subduction started has travelled nearly 200 km and has cooled significantly, but the effect on surface heat flow is felt mostly offshore (Fig. 6a). The high heat flows predicted for the coastal region at this time readily explain the high geothermal gradients qualitatively estimated for the mid-Miocene thermal pulse in the Shimanto Belt (Hibbard and Karig, 1990a; Underwood et al., 1993). The model surface temperature gradient at 5 Ma is about 50 K km-1. The distribution of surface heat flow for 15 Ma (present) and 20

Kelin Wang et aL / Tectonophysics 248 (1995) 53-69

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300

Distance (km) Fig. 7. Temperature contours of the preferred model at four time steps. 15 Ma approximately corresponds to the present. The dashed line shows the plate interface.

mation front, which agreed well with the seismologically and geodetically determined one. Based on the steady-state solution, however, the seismogenic zone would be about 50 km wider downdip. This shows the importance of considering the tectonic history of individual subduction zones when relating their seismogenic behaviour with their geothermal regimes. The most important result of the model is that the landward decreasing trend of the observed surface heat flow can be explained by the subduction model at t = 15 Ma. Small-scale variations of measured heat flows are related to local processes, as discussed in section 2.2. These include the effects of advective heat transfer by magma, groundwater, or seafloor hydrothermal circulation. Because of the presence of these effects, a point-by-point comparison of calculated and observed heat flow is not meaningful. There remain three major issues. First, the rapid landward decrease of BSR heat flow within 60 km of the deformation front is not predicted by the model. Second, different scenarios for the thermal condition prior to 15 Ma need to be investigated. Third, the potentially important heat source, frictional heat along the plate interface, has not been considered. We examine each of these three issues below.

4.2. On the rapid landward decrease of BSR heat

flow Ma are similar, with the latter about 10 m W m 2 lower in the first 200 km. Also shown in Fig. 6 are values associated with the steady-state subduction of a 15-Ma-old oceanic plate. These are the values that would be obtained if the subduction of the fossil ridge and the aging of the subducting slab were ignored. The difference in plate interface temperatures between the transtent and steady-state solutions becomes significant when the seismogenic behaviour of the fault is considered. Hyndman and Wang (1993) summarized the evidence that the landward end of the selsmogenic zone is limited by a temperature of 350 ° C. Hyndman et al. (1995) found that the downdip limit of the seismogenic zone inferred from the results of our preferred model (Fig. 6b) was about 150 km from the defor-

The BSR heat flows reported by Ashi and Taira (1993) (Fig. 2) decrease systematically from 130 m W m -z at the deformation front to one third of this value in a distance of 60 kin. Ashi and Taira (1993) found that the heat-flow values were negatively correlated with the thickness of the accretionary wedge sediments and, according to a steady-state model, that they required the base of the sediment to be at a nearly uniform temperature of 120 ° C. This rapid heat flow decrease landward is very puzzling. For a conductive model, the temperature on the plate interface always increases landward (Fig. 7) because the isotherms have a tendency to stay parallel with the upper surface which is maintained at a constant temperature. An isothermal

Kelin Wang et al. / Tectonophysics 248 (1995) 53-69

wedge basement would induce large lateral heat transfer and the temperature on the plate interface would very soon increase landward again. One mechanism Ashi and Taira (1993) proposed is fluid flow in the oceanic crust or along the subduction thrust fault. Water flow from the deformation front downdip along the plate interface is unlikely. The high trench heat flow, if anything, indicates f u id discharge. If water flows updip from a deeper region, the 120°C temperature is much too low. For the f u i d flow mechanism, the observed heat flow seems to require that water recharge at the landward end of the 60-km region to suppress heat flow, circulate through the sediment and the basement, then discharge at the seaward end to enhance heat flow, a mechanism believed to prevail in a number of continental sedimentary basins (Wang, 1992). However, there is no known mechanism to drive the fluid flow this way. Landward thickening of the wedge sediment also causes the decrease of heat flow. However, applying the model of Wang et al. (1993) to this area, we find that the perturbation to surface heat flow by this source is only a few percent. Small-scale hydrothermal circulations through the seafloor sediment must exist, judging from the scattering of heat-flow data, but they can not produce a systematic trend of landward decrease. In summary, we have not been able to find a satisfactory explanation for the rapid landward decrease in BSR heat flow. The uncertainty in BSR heat flow increases landward to about 30% because on the upper continental slope the subseafloor depth to the BSR decreases and the uncertainties in BSR and seafloor temperatures become significant. The possibility of a systematic error in BSR heat-flow determination cannot yet be excluded. This is the subject for further study. 4.3. Thermal regime prior to 15 M a

We have employed a very simple initial condition, the temperature field of a half-space cooled from 1400°C for 80 Ma in the preferred model. The poor knowledge of the thermal state of the subduction zone prior to 15 Ma (t = 0 in the model) does not warrant a more complex initial condition. Low temperatures before the subduc-

63

tion of the active ridge are supported by the inferred short time interval near 15 Ma for the thermal imprint of the subducted ridge on the coastal Shimanto Belt (Hibbard and Karig, 1990a; Underwood et al., 1993). However, we have investigated how and to what extent different tectonic scenarios prior to 15 Ma affect the results. As summarized in section 2.2, models for the tectonics prior to 15 Ma include either subduction of active Shikoku spreading ridge (or young lithosphere) or subduction of an 80-Ma-old oceanic (Pacific) plate. We simulate these two situations by using as initial conditions the steady-state (continuing for infinite time) temperature fields resulted from the subduction of a 1-Ma-old and an 80-Ma-old oceanic lithosphere, respectively. An intermediate initial condition is also tested, which involves the steady-state temperature field associated with the subduction of a 15-Ma-old oceanic lithosphere. The calculated surface heat flows at t = 15 Ma using all three of the initial conditions are shown in Fig. 8a, and the calculated temperatures along the plate interface at the same time in Fig. 8b. The values for the preferred model and the steady-state solution with a 15-Ma-old subducting slab are also shown for comparison. As is evident from Fig. 8, there is no difference in surface heat flows and plate interface temperatures between the four transient models within 150 km of the deformation front. This justifies the use of the results of the preferred model by Hyndman et al. (1995) in the study of the downdip width of the seismogenic zone the subduction thrust fault, where the region of interest is small (within 150 km). Significant difference between the models occur beyond 250 km landward of the deformation front where the present surface heat flow is little affected by the subduction over the past 15 Ma. The thermal effects of the slab subduction during that period has not yet reached the surface in that region. The presently observed values thus potentially provide some constraints on the thermal regime prior to 15 Ma. The available land heat flow and heat generation data contain large uncertainties and show significant scattering, but they favour a relatively young subducted plate prior to 15 Ma. However, to

Kelin Wang et al. / Tectonophysics 248 (1995) 53-69

64

I

,.-... 150

E

~

•

b

energy by viscous flow of material in the plate interface shear zone. In the brittle regime, the Byerlee's law of friction is used (Byerlee, 1978):

I

o

,~o

r=O.85~r.(1-a)

100

[¢.(1-A)

r = 50 + 0.6cr.(1 - A ) 0

(2a)

[O'n(1- A ) > 200MPa] (2b)

i-i- 50 coast

[

o

"1-

<200MPa]

0

I

I

I

800

(a)

~

I

0 ~ " 600

3 ............~

~400 O

"'"""~

...... .-

Q.

E 200 O

I

I-

0 0

(b)

I

I

100

200

300

Distance (km) Fig. 8. (a) Heat flows at 15 Ma (present) calculated with different initial conditions, and (b) the associated plate interface temperatures. The initial conditions are the steady-state temperature fields associated with the subduction of (1) 1 Ma. (2) 15 Ma and (3) 80 Ma slabs. In both (a) and (b), values of the preferred model are shown as dashed lines, and values of the steady state model involving the subduction of a 15-Ma plate as dotted lines. See Fig. 6 for definition of heat-flow data symbols.

constrain the thermal regime in detail prior to 15 Ma, it is necessary to obtain further high-quality thermal data from that region. In summary, thermal regime within 200 km of the deformation front is little affected by, but that beyond 200 km depends mainly on, the thermal state prior to 15 Ma. All three models with subduction initial conditions have heat-flow values lower than observed at distances of 200-300 km. This suggests that other mechanisms such as back flow in the mantle (e.g., Honda, 1985) may control the thermal regime there.

4.4. Frictional heating By frictional heating, we include both heating by friction of brittle rocks and dissipation of heat

where ~- is the shear stress on the fault plane in MPa, ~r,1 is the normal stress, and A is the pore pressure ratio defined as A = (pf - P d ) / ( P l -- Pd), with pf, p~ and Pa being the pore fluid pressure, the lithostatic pressure and the pressure at the earth's surface (nonzero for the seafloor), respectively. Because of the gentle dip of the plate interface, we can approximate (r, by the load of the rock column above, pgz, where p is the average density (2750 kgm-3), and g is gravity. By using (2), we assume that the frictional coefficients of the thrust fault are the same as those of the country rocks. In the plastic regime of the present model, the following flow law for diabase (Caristan, 1982) is employed: = A(2z)" exp(-Q/RT)

(3)

where k is strain rate, R is the universal gas constant, and A, Q and n are experimentally determined parameters, with values given by Caristan (1982). The shape of the strength envelopes thus defined is schematically illustrated in Fig. 9. The amount of frictional heat per unit fault width downdip for the brittle regime for a sliding velocity u is given by: Qf = ~-u

(4)

and that per unit volume in the shear zone along the plate interface for the ductile regime by: Qr = z~

(5)

In the numerical implementation, the frictional heat is formulated as body heat sources in a thin element layer with thickness w (500 m in the present model) along the plate interface. In the brittle domain, the numerical solution employs (5) instead of (4) because of the finite thickness of the layer, with ~ = v / w ( ~ 10-12). In the plastic regime, the heat generated due to shear heating depends on temperature through

Kelin Wang et aL / Tectonophysics 248 (1995) 53-69

(3), and the problem is nonlinear. For eacn ume step, the fixed point iteration scheme is used. During each iteration, both (2) and (3) are used to calculate shear stress, and the smaller one is used to update the shear heat; the resulted shear heat, in turn, is used to update the temperatures. Two models are generated with h = 0.8 and 0.95, respectively. All other parameters used are the same as in the preferred model. The calculated heat flows and temperatures along the plate interface, at t = 15 Ma (present) in comparison with that of the preferred model, are shown in Fig. 10a and b, respectively. The average shear stress on the fault is about 40 MPa for A = 0.8 and 20 MPa for h = 0.95. The temperature field for the h = 0.8 case is shown in Fig. 10c. Even with the relatively high pore fluid pressures and hence low shear stress (see Eq. 2), there is significant increase of surface heat flow in both models associated with frictional heating. The amount of frictional heating decreases rapidly landward in the plastic domain because of the decrease of stress at high temperatures (see Eq. 3). Unfortunately, the most diagnostic region where surface heat flow can be used to constrain the amount of frictional heating, the continental shelf, does not have reliable data. However, with the addition of

..-..150

" o

'

65

I

o

'

r

,

,

c°ist .

,

,

'

'

=

'

.

,

.

,

.

E ~E~ 100

k == 0 . 8 o

IT 50 (a) -1-

0

..--.. 800

600 ~400

E200

'~)

I-

0

N

50

c°ist

121 lOO 0

1 O0

200

300

Distance (km)

Shear Stress

Fig. 10. (a) Surface heat flows calculated at 15 Ma (present) with frictional heating (solid lines), with two different pore pressure ratio values. A diabase rheology (Caristan, 1982) is assumed for the ductile regime. Heat flow of the preferred model (no frictional heating) is shown as a dashed line. (b) Temperatures along the plate interface associated with the heat flows in (a). (c) Temperature field at 15 Ma for a model with frictional heating (A = 0.8). The dashed line shows the plate interface. See Fig. 6 for definition of heat-flow data symbols•

a

Fig. 9. Schematic illustration of the yield stress of the subduction stress fault, used for the calculation of frictional heating. A straight lines represents the brittle regime and a curved line represents the plastic regime. The point of intersection is the brittle-plastic transition.

frictional heating, the predicted heat flow on the upper continental slope (distance 30-60 km), which is already too high compared to the BSR values even without frictional heating, is now much higher. Wang et al. (1995) have shown using better resolved thermal data and earthquake data that shear stress and hence frictional heating on the Cascadia subduction thrust fault is

Kelin Wang et aL / Tectonophysics 248 (1995) 53-69

66

very low ( < 20 MPa), and argued that it is likely the case for most subduction boundaries. It is possible that the shear stress on the Nankai Trough subduction thrust fault is also low. In summary, definite conclusion on the amount of frictional heating for the Nankai Trough subduction zone cannot be drawn given the available thermal data, although we suspect that the value would be very low.

...... 1 5 0

0l

100

I

t

0

J

,

t~

coast

o

i

i

100

200

300

Distance (km)

In all the models presented so far, we have let all the material below the subducting slab move together with the slab. The thickness of the material beneath the fault that subducts should be controlled by the temperature-sensitive mantle rheology and the dynamics, but this is beyond the ability of the present kinematic thermal model. To evaluate the effect of varying the slab thickness, we try another extreme case, namely, one for which the thickness of the subducting slab is only 15 km and the material below it is fixed. This is like a dipping flow channel in the mantle. All the other parameters are the same as in the preferred model. The surface heat flows (not shown) calculated at all the four times are nearly identical to those of the preferred model (Fig. 6a). The calculated temperatures below the slab, shown in Fig. 11 only for t = 15 Ma (present), however, are quite different from the preferred model. The temperatures below the slab change less from the initial condition compared to the preferred model.

a

,

_° 0

Slab thickness

-'-'-

i

. .

"1-

4.5. Effects of some other parameters

,

~

i

1O0

J

~.

i

200

,

~

~.1

I

300

Fig. 12. Surface heat flow calculated at four time steps in a model similar to the preferred model (Fig. 6a) but with the growth of the accretionary prism over the past 15 Ma ignored. See Fig. 6 for definition of heat-flow data symbols.

Neglecting wedge growth If we assume that most of the modern accretionary wedge was part of the overriding plate at 15 Ma, that is, the initial temperatures of all nodes at x = 100 km and seaward are assigned the same way as for the rest of the preferred model, we obtain the calculated surface heat flow shown in Fig. 12. The thermal pulse induced by the subduction of the ridge is not yet felt at the coast at t = 5 Ma (10 Ma ago). This is in conflict with the observed high Miocene geothermal gradient in the coastal Shimanto belt (Hibbard and Karig, 1990a). However, the recent thermal regime, such as heat flow at 15 Ma since the models begins, is essentially the same as in the preferred model.

Com, ergence L,elocity We have used a plate convergence velocity of 40 m m a -1, but recognize that there are some uncertainties in this value. In Fig. 13, surface heat flows and temperatures along the plate interface at t = 15 Ma (present) calculated with convergence velocities 10 m m a -1 higher and lower than this value are compared with those of the preferred model. The results are not significantly affected.

Distance (km) Fig. 11. Temperature contours at 15 Ma (present) for a model in which a plate of 15 km thickness is subducted. All other parameters are the same as in the preferred model.

Presence of a later phase of spreading According to Chamot-Rooke et al. (1987), the Shikoku basin experienced another phase of

Kelin Wang et aL / Tectonophysics 248 (1995) 53-69 150

,

E

~o

~EIO0

•

i

,

"

i

67

values are higher, with the difference decreasing with time.

,

o

°o 5. Conclusions

o

~. _

5o

",,.;

...... .1.

tl:l

coast

"1-

o 0

I

I

1O0

200

300

D i s t a n c e (krn) Fig. 13. Surface heat flow calculated at 15 Ma (present) with different plate convergence velocities. The dashed line is the preferred model. All other parameters are the same in different models. See Fig. 6 for definition of heat-flow data symbols.

spreading at 14-12 Ma. Although this interpretation is not conclusive, there is indication that the Shikoku basin might remain thermally active shortly after the cessation of spreading at 15 Ma (Okino et al., 1994). In Fig. 14, we present the calculated heat flow assuming spreading stops at t = 3 Ma (at 12 Ma). The heat flow results are compared with those of the preferred model. The

150 j \ .

~E,~100 ~ .

,

-

'\

0

'

o~--...~.'k~_

1

ot=

'

T

'

t ql-

.................

1oo

r:t

f

20o

30o

D i s t a n c e (km) Fig. 14. Surface heat flow (solid lines) calculated at four time steps in a model similar to the preferred model (Fig. 6a) in which the Shikoku basin stopped spreading at 12 Ma, i.e., 3 Ma after the beginning of the model. The boundary conditions shown in Fig. 5 were all deferred by 3 Ma. The heat flows for the preferred model (Fig. 6a) at the same time steps are shown as dotted lines and labelled with time since the model begins. See Fig. 6 for definition of heat-flow data symbols.

We developed a 2-D finite-element model for the thermal regime of the Southwest Japan subduction zone. A critical feature of the recent subduction history is the Shikoku basin fossil spreading ridge perpendicular to the margin, which stopped spreading at 15 Ma, and has subsequently been subducted at the Nankai Trough. The effects of the subduction of the fossil ridge and the increasing age of the subducting plate is the focus of the present thermal model. Based on the model results we draw the following main conclusions: (1) The observed landward decrease of surface heat flow is consistent with the recent subduction history of the margin. The Philippine Sea plate brought high temperatures to the forearc in the early history of subduction but relatively low temperatures now. (2) We have not been able to explain the rapid landward decrease of heat flow within 60 km of the deformation front as inferred from BSR data. If fluid flow is to be used as an explanation, rather extreme conditions must occur. (3) Within about 150 km of the deformation front, the forearc thermal regime depends only on the subduction history of the past 15 Ma, not the previous thermal state. Beyond 250 km, the present near-surface thermal regime is not affected by the subduction of the past 15 Ma if heat transfer in the overriding plate is by conduction only. (4) The amount of frictional heating for the Southwest Japan subduction zone is probably low, but it cannot yet be well resolved given the available thermal data. Reliable heat-flow data are required from the uppermost continental slope, the continental shelf, or the immediate coastal region. (5) The thermal model that takes into account the effects of the subduction of the fossil spreading ridge and the aging of the subducting slab is significantly different from the steady-state model

68

Kelin Wang et a L / Tectonophysics 248 (1995) 53-69

for the subduction of a 15-Ma-old lithosphere. Van den Beukel and Wortel (1988), Molnar and England (1990), Dumitru (1991) and Wang et al. (1995) have emphasized that some or all of the following factors must be considered carefully in thermal subduction zone models: plate age, convergence rate, seafloor sediment thickness and slab geometry. The results presented in this paper further demonstrate that the detailed local plate tectonic history also has to be taken into account. For subduction zones such the Cascadia in the northeast Pacific, a steady-state solution is a good approximation; but for subduction zones like the Nankai Trough, transient models constrained by known subduction histories are required. This is especially important when the seismogenic behaviour of the subduction thrust fault is considered. Generic subduction models, especially the steady-state type, are in most cases inadequate for such purpose.

Acknowledgements We thank A. Taira for fruitful discussions on the geology of the Outer Zone of Japan. Reviews by M. Underwood, C. Lowe, T. Dumitru and an anonymous referee improved the manuscript. Geological Survey of Canada contribution 52194.

References Ando, M., 1975. Source mechanisms and tectonic significance of historical earthquakes along the Nankai Trough. Tectonophysics, 27:119-140. Ashi, J., 1991. Structure and hydrogeology of the Nankai accretionary prism, Ph.D. Thesis, Univ. Tokyo, Tokyo, 120 PP. Ashi, J. and Taira, A., 1993. Thermal structure of the Nankai accretionary prism as inferred from the distribution of gas hydrate BSRs. In: M.B. Underwood (Editor), Thermal Evolution of the Tertiary Shimanto Belt, Southwest Japan: An Example of Ridge-Trench Interaction. Geol. Soc. Am., Spec. Pap., 273: 137-149. Byerlee, J.D., 1978. Friction of rocks. Pure Appl. Geophys., 116: 615-626. Byrne, J. and Ditullio, L., 1992. Evidence for changing plate motions in southwest Japan and reconstructions of the Philippine Sea plate. Isl. Arc, l: 148-165.

Cande, S.C., Leslie, R.B., Farra, J.C. and Hobart, M., 1987. Interaction between the Chile ridge and the Chile trench: Geophysical and geothermal evidence. J. Geophys. Res., 92: 495-520. Caristan, Y., 1982. The transition from high temperature creep to fracture in Maryland diabase. J. Geophys. Res., 87: 6781-6790. Chamot-Rooke, N., Renard, V. and Le Pichon, X., 1987. Magnetic anomalies in the Shikoku Basin: A new interpretation. Earth Planet. Sci. Lett., 83: 214-228. Davis, E.E. and Lister, CR.B., 1974. Fundamentals of ridge crest topography. Earth Planet. Sci. Lett., 21: 405-413. Davis, E.E., Hyndman, R.D. and Villinger, H., 1990. Rates of fluid expulsion across the northern Cascadia accretionary prism: Constraints from new heat flow and multichannel seismic reflection data. J. Geophys. Res., 95: 8869-8889. DeLong, S.E., Schwarz, W.M. and Anderson, R., 1979, Thermal effects of ridge subduction. Earth Planet. Sci. Lett., 44: 239-246. Dumitru, T.A., 1991. Effects of subduction parameters on geothermal gradients in forearcs, with an application to Franciscan subduction in California. J. Geophys. Res., 96: 621-641. Furukawa, Y., 1995. Temperature structure in the crust of the Japan arc and the thermal effects of subduction. In: M.L. Gupta and M. Yamano (Editors), Terrestrial Heat Flow and Geothermal Energy in Asia. Oxford and IBH Publ. Co., New Delhi, pp. 203-219. Hibbard, J.P. and Karig, D.E., 1990a. Structural and magmatic responses to spreading ridge subduction: An example from southwest Japan. Tectonics, 9: 207-230. Hibbard, J.P. and Karig, D.E., 1990b. Alternative plate model for the early Miocene evolution of the Southwest Japan margin. Geology, 18: 170-174. Hill, I.A. et al., 1993. Proc. ODP, Sci. Results 131. Hirahara, K., 1981. Three-dimensional seismic structure beneath Southwest Japan: The subducting Philippine Sea plate. Tectonophysics, 79: 1-44. Honda, S., 1985. Thermal structure beneath Tohoko, Northeast Japan - - a case study for understanding the detailed thermal structure of the subduction zone. Tectonophysics, 112: 69-102. Honda, S. and Uyeda, S., 1983. Thermal process in subduction zones - - a review and preliminary approach on the origin of arc volcanism. In: D. Shimozuru and I. Yokoyama (Editors), Arc Volcanism: Physics and Tectonics. TERRAPUB, Tokyo, pp. 117-140. Hyndman, R.D. and Wang, K., 1993. Thermal constraints on the zone of major thrust earthquake failure: The Cascadia subduction zone. J. Geophys. Res., 98: 2039-2060. Hyndman, R.D., Wang, K., Yuan, T. and Spence, G.D., 1993. Tectonic sediment thickening, fluid expulsion, and the thermal regime of subduction zone accretionary prisms: The Cascadia margin off Vancouver Island. J. Geophys. Res., 98: 21,865-21,876. Hyndman, R.D., Wang, K. and Yamano, M., 1995. Thermal constraints to the seismogenic portion of the Southwest Japan subduction thrust. J. Geophys. Res. (in press).

Kelin Wang et aL / Tectonophysics 248 (1995) 53-69 Hutchison, I., 1985. The effects of sedimentation and compaction on oceanic heat flow. Geophys. J. R. Astron. Soc., 82: 439-459. Ito, K., 1990. Seismic activity and tectonics of southwestern Japan. Zisin (J. Seismol. Soc. Jpn.), 4 3 : 5 5 5 - 5 6 9 (in Japanese). Jolivet, L., 1987. America-Eurasia plate boundary in eastern Asia and the opening of marginal basins. Earth Planet. Sci. Lett., 81: 282-288. Kanaya, H. and Ishihara, S., 1975. Uranium, thorium and potassium contents of Japanese granitic rocks: a summary up to 1972. In: J.A.S. Adams, W.M. Lowder and T.F. Gesell (Editors), The Natural Radiation Environment II. U.S. Energ. Res. Dev. Adm., Washington, DC, pp. 517533. Kinoshita, M. and Yamano, M., 1986. The heat flow anomaly in the Nankai Trough area. Init. Rep. DSDP, 87: 737-743. Kinoshita, M. and Yamano, M., 1995. Heat flow distribution in the Nankai Trough region. In: S.A. Shcheka et al. (Editors), Geology and Geophysics of the Philippine Sea Floor. TERRAPUB, Tokyo (in press). Li, X., Furukawa, Y., Nagao, T., Uyeda, S. and Suzuki, H., 1989. Heat flow in central Japan and its relations to geological and geophysical features. Bull. Earthquake Res. Inst., 64: 1-36. McCaffrey, R., 1993. On the role of the upper plate in great subduction zone earthquakes. J. Geophys. Res., 98: 11,953-11,966. Mizoue, M., Nakamura, M., Seto, N. and Ishiketa, Y., 1983. Three layered distribution of microearthquakes in relation to focal mechanism variation in the Kii peninsula, Southwest Japan. Bull. Earthquake Res. Inst., 58: 287-310. Molnar, P. and England, P., 1990. Temperatures, heat flux, and frictional stress near major thrust faults. J. Geophys. Res., 95: 4833-4856. Moore, G.F., Shipley, T.H., Stoffa, P.L., Karig, D.E., Taira, A., Kuramoto, S., Tokuyama, H. and Suyehiro, K., 1990. Structure of the Nankai Trough accretionary zone from multichannel seismic reflection data. J. Geophys. Res., 95: 8753-8765. Okano, K., Kimura, S., Konomi, T. and Nakamura, M., 1985. The focal distribution of earthquakes in Shikoku and its surrounding regions. Zisin (J. Seismol. Soc. Jpn.), 38: 93-103 (in Japanese). Okino, K., Shimakawa, Y. and Nagaoka, S., 1994. Evolution of the Shikoku basin. J. Geomagn. Geoelectr., 46: 463-479. Olafsson, G., 1993. Calcareous nonnofossil biostratigraphy of the Nankai Trough. Proc. ODP, Sci. Results 131. Otsuki, K., 1990. Westward migration of the Izu-Bonin Trench, northward motion of the Philippine Sea Plate, and their relationships to the Cenozoic tectonics of Japanese island arcs. Tectonophysics, 190: 351-367. Sato, T. and Matsu'ura, M., 1992. Cyclic crustal movement, steady uplift of marine terraces, and evolution of the island arc-trench system in southwest Japan. Geophys. J. Int., 111: 617-629. Seno, T. and Maruyama, S., 1984. Palaeographic reconstruc-

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tion and origin of the Philippine Sea. Tectonophysics, 102: 53-84. Seno, T., Stein, S. and Gripp, A.E., 1993. A model for the motion of the Philippine Sea plate consistent with Nuvel-1 and geological data. J. Geophys. Res., 98: 17,941-17,948. Taira, A., 1985. Sediment evolution of Shikoku subduction zone: The Shimanto belt and Nankai Trough. In: N. Nasu et al. (Editors), Formation of Active Ocean Margins. TERRAPUB, Tokyo, pp. 835-851. Taira, A., 1988. The Shimanto belt in Shikoku--Evolution of Cretaceous to Miocene accretionary prism. Mod. Geol., 12: 5-46. Taira, A., Hill, I., Firth, J. et al., 1992. Sediment deformation and hydrogeology of the Nankai Trough accretionary prism: Synthesis of shipboard results of ODP Leg 131. Earth Planet. Sci. Lett., 109: 431-450. Taira, A., Hill, I., Firth, J. et al., 1993. Proc. ODP, Sci. Results 131. Underwood, M.B., Byrne, T., Hibbard, J.P., DiTullio, L. and Laughland, M.M., 1993. The effects of ridge subduction on the thermal structure of accretionary prisms: A Tertiary example from the Shimanto Belt of Japan. In: M.B. Underwood (Editor), Thermal Evolution of the Tertiary Shimanto Belt, Southwest Japan: An Example of Ridgetrench Interaction. Geol. Soc. Am. Spec. Pap., 273: 151168. Wang, K., 1992. Groundwater flow and geotemperature pattern. In: W.A. Nierenberg (Editor), Encyclopedia of Earth System Science. Academic Press, San Diego, CA, pp. 441-453. Wang, K., Hyndman, R.D. and Davis, E.E., 1993. Thermal effects of sediment thickening and fluid expulsion in accretionary prisms: Model and parameter analysis. J. Geophys. Res., 98: 9975-9984. Wang, K., Mulder, T., Rogers, G.C. and Hyndman, R.D., 1995. Case for very low coupling stress on the Cascadia subduction fault. J. Geophys. Res. (in press). Van den Beukel, J. and Wortel, R., 1988. Thermo-mechanical modelling of arc-trench regions. Tectonophysics, 154: 177-193. Yamano, M., 1995. Recent heat flow studies in and around Japan. In: M.L. Gupta and M. Yamano (Editors), Terrestrial Heat Flow and Geothermal Energy in Asia. Oxford and IBH Publishing Co., New Delhi, pp. 173-201. Yamano, M., Uyeda, S., Aoki, Y. and Shipley, T.H., 1982. Estimates of heat flow derived from gas hydrates. Geology, 10: 339-343. Yamano, M., Honda, S. and Uyeda, S., 1984. Nankai Trough: A hot trench? Mar. Geophys. Res., 6: 187-203. Yamano, M., Foucher J.-P., Kinoshita, M., Fisher, A. and Hyndman, R.D., 1992. Heat flow and fluid regime in the western Nankai accretionary prism. Earth Planet. Sci. Lett., 109: 451-462. Yoshika, S., 1991. The interplate coupling and stress accumulation process of large earthquakes along the Nankai Trough, Southwest Japan, derived from geodetic and seismic data. Phys. Earth Planet. Inter., 66: 214-243.