For lectures on April 22, May 6, 13, and 20, 2005 (May 16, 2005; Revised Sept. 20, 2005)

Plate Tectonics by T. Seno For lectures on April 22, May 6, 13, and 20, 2005 (May 16, 2005; Revised Sept. 20, 2005) 2. Plate boundary processes Three types of the boundaries There are three types of differential motion between two plates; they are divergent, strike-slip (transcurrent), and convergent motions. Associated with these motions, three types of plate boundaries, i.e., divergent, strike-slip (transcurrent), and convergent boundaries exist (Fig. 1). The strike-slip boundary is called a transform fault. The reason why it is called transform fault would be explained in 3. Plate kinematics. Topographic features Plate boundaries are marked by the belts of shallow seismicity as seen before. Along these seismic belts, prominent topographic features are associated. They are: mid-ocean ridges (MOR), which are topographic bulges running in the mid of large oceans, fracture zones, which translate MOR, deep-sea trenches, and mountain belts (Fig. 2). If divergent and convergent boundaries correspond to the upwelling and downwelling sites of the convection cell, topographic bulges and depressions, respectively, should be the morphological features of these boundaries on the earth's surface. In fact, divergent and convergent boundaries are located along such features, i.e., the mid-oceanic ridge and deep-sea trench. Another type of the convergent boundary is located along a mountain belt in the continent. This will be explained later. It is noticed that a transform fault does not appear in the 2-D scheme of convection. This might be due to its 2-dimensionality; however, even in 3-D simulations of convection, transform faults do not appear (Fig. 3). This is because buoyancy or negative buoyancy that drives convection can produce only poloidal velocity fields (rotation free), not toroidal fields (divergence free). However, on the Earth, troidal motions with vorticity, not only poloidal motions, appear (Fig. 4). We notice here that the appearance of transform faults is closely related to the second element of the basic concepts, i.e., the appearance of a number of plates, because if a rigid plate moves w.r.t. other neighboring plates, toroidal fields must be generated at some portion of the boundaries. Topographic feature of this type of the boundary is not obvious, because transcurrent or strikeslip motion itself does not produce any uplift or depression. However, we will see that fracture zones are the morphological features of this boundary. Types of faulting at plate boundaries Types of faulting expected at each type of the boundaries are normal fault, strike-slip fault, and reverse fault, respectively. Because earthquakes are faulting (Fig. 5), we expect normal, strike-slip and reverse fault-type earthquakes in these boundaries. Fault parameters

Types of faulting are described by three fault parameters of an earthquake (Fig. 6). They are the strike φ, dip angle δ, and rake (slip) angle λ. If faulting is of mixed type, type of faulting is determined by which component of the slip is dominant. Practice 2.0.1 Determine the ranges of the rake angle representing non-pure but dominantly normal, reverse and strikeslip faulting. Focal mechanism A graphical way representing faulting is often used and called a focal mechanism. Suppose a fault plane be determined by the strike φ and the dip angle δ. Draw another plane, which is vertical to the slip direction. This is called an auxiliary plane (Fig. 7). Because this plane can define the slip direction, i.e., λ, these two orthogonal planes are enough to describe faulting except for the sense of motion. To specify the sense, we color the four spaces divided by these two planes by white and black as the white block moves toward the black one. These four sub-spaces with the two planes are embedded in a small hypothetical sphere at the hypocenter. This sphere is called a focal sphere. The two planes in this sphere are called nodal planes. Only upper or lower hemi-sphere is enough to represent the faulting, and the lower hemi-sphere is more often used. The hemi-sphere is projected on a horizontal plane by an equal area projection. The nodal planes projected into the horizontal plane are called nodal lines. The nodal lines can be determined from the P-wave first motions observed at various seismological stations. The P-waves emanated from the white quadrants show negative first motions, and those from the black quadrants show positive first motions (Fig. 7). These polarities are traced back along the rays to the epicenter and plotted in the focal sphere, to determine the nodal lines. Practice 2.0.2 Draw typical focal mechanisms representing motion between plates at each type of plate boundaries. Slip vector In plate tectonics, a slip vector is useful to describe plate motions. A slip vector is the horizontal projection of the slip on the fault plane (Fig. 8). It represents the differential motion between two plates, looked from above. Because the auxiliary plane is orthogonal to the slip on the fault plane, the line connecting two end points of the nodal line of the auxiliary plane is orthogonal to the slip vector. Therefore, it is very easy to draw a slip vector in a given focal mechanism. 2.1 Divergent boundary 2.1.1 Creation of a plate At the divergent boundary, two plates are separating from each other. There should appear vacancy between these plates. This vacancy is filled by the asthenosphere rising from below. In the scheme of convection, this corresponds to an upwelling current. The age distribution of the ocean floor (Fig. 9), which is youngest at the MOR and is getting older away from it, directly indicates that the MOR is a divergent plate boundary where an oceanic plate is created. Therefore, the most fundamental and essential phenomenon at the divergent boundary would be "plate creation.

The morphology of the MOR is not a localized mountain belt. In fact, the ocean floor increases its depth with a gentle slope of a few km over thousands of km (Fig. 10). This gentle increase of the seafloor depth reflects the increase of the plate thickness as shown by the practice below. Practice 2.1.1 Using the isostasy, obtain a relationship between the plate thickness (a) and the ocean floor depth with respect to the MOR (hw), using the average density of a plate ρop, the density of the water ρw, and the density of the asthenosphere ρm. Using ρop and ρm values of practice 1.1.6, and ρw = 1000 kg/m^3, get a value of hw for a = 100 km. 2.1.2 Creation of an oceanic crust Volcanism The second most essential phenomenon occurring at the divergent boundary would be volcanism. Volcanism at this boundary means that a partial melting occurs beneath the MOR. The accumulated melts become magmas, which are erupted as sheets of pillow basalts and are injected as dikes (Fig. 11). The reason why melting occurs beneath the MOR is as follows (Fig. 12). The temperature of the asthenosphere is Tm, which is higher than the melting temperature of mantle rocks (solidus, Ts) at the surface. On the other hand, the solidus increases as the confining pressure (i.e., depth) increases. Therefore, the asthenosphere does not melt at depth larger than ~60 km, where the solidus equals to Tm. If the upwelling current reaches this depth (~60km), it starts to melt. Degree of melting increases as the upwelling current reaches shallower depths. This type of melting is called decompressional melting. Mantle plumes beneath hotspots have this type of melting, although their higher temperatures than Tm imply larger degrees of melting. Spreading rates The melting of mantle rocks produces chemically different materials from the mantle, i.e., basaltic rocks. The surface several km of the plate becomes oceanic crust, which has a low density (~2800 kg/ m^3) than the mantle. The creation of oceanic crust has important meaning for plate tectonics. Because the magma, which produces oceanic crust, contains tiny magnetic minerals, they acquire remnant magnetization when they solidify, proportional to the current ambient geomagnetism. We can observe this magnetization as magnetic anomaly stripes in the ocean floor (Fig. 13). Because we know a history of magnetic reversals independently from igneous rocks on land (Fig. 14), we can assign ages to the stripe boundaries, and determine the spreading velocity of ocean floor. This provides us estimates of velocities between two separating plates, which are unavoidable for plate motion determination. Altered basalts Hydrothermal circulation occurs within the crust near the MOR due to volcanic activities. Water circulates mainly in the basaltic layer, and partly in the gabbroic layer (Fig. 15). This metamorphoses basalts into greenschists including hydrous minerals such as amphiboles. The maximum water content of the altered metamorphosed basalts amounts to 6 %. When an oceanic plate reaches the trench and is subducted beneath an overriding plate, the temperature and the pressure rise, and the altered basalts

dehydrate, and provide water at the plate boundary thrust. This would also be a cause for occurrence of intraslab earthquakes in subduction zones as will be seen in 2.3. 2.1.3 Normal faulting Earthquakes occurring along the MOR have normal fault-type mechanisms (Fig. 16). This looks consistent with the divergent motion occurring there. However, the plate thickness is small, and it is not easy to define the plate interface (the plate boundary fault). Many of these events occur on either side of the rift valley of the MOR, and are rather regarded as intraplate earthquakes. They are thus not useful for plate motion determination. There are many normal faults near the ridge axis, consistent with the normal fault-type earthquakes (Fig. 17). 2.2 Transform fault Transform faults are divided into ones in oceans and others in continents. Typical oceanic transform faults translate axes of MORs (Fig. 18) or those of trenches. The morphology associated with oceanic transform faults is fracture zones. 2.2.1 Ridge-ridge transform Fracture zone Shallow seismicity is located between an axis and another offset axis of a MOR. This clearly shows only part between these two axes constitutes a transform fault (Fig. 19). However, a fracture zone further extends beyond the transform to both sides. Then, a question arises why a fracture zone extends beyond the transform. Practice 2.2.1 Explain the reason why a fracture zone extends beyond the transform. (A hint) In 2.1, we have learned that the seafloor depth with respect the MOR mimics the plate thickness, and this depth increases away from the ridge, i.e., as the age of the plate increases. See that ages of arbitrary opposing two points C and D across a fracture zone (Fig. 18) are different. Therefore the morphology of a fracture zone varies along its strike depending on the age difference between the two opposing plates. Fig. 20 shows schematically the seafloor topography and crosssections along a fracture zone. Strike-slip faulting Focal mechanisms of earthquakes along a transform are of strike-slip fault-type (Fig. 21). The sense of faulting is dependent on the sense of the offset between the ridge axes. If the ridge axes are offset by a right (left)-step, the faulting is left (right)-lateral. This implies that the ridge axes are not offset by the motion on the transform. Then a question arises why such an offset is produced, or, in other words, how a ridge-ridge transform is generated.

Evolution Let's consider a splitting of a continent like Pangea (Fig. 22). The line of splitting is quite different from the present ridge-transform geometry. This indicates that the present ridge-transform geometry has been developing from the original splitting line since the initial continental breakup (rifting). At present, the transform fault strike is nearly orthogonal to the ridge axis (Fig. 22). This indicates that the length of the transform faults has been increasing and that of the ridge axes has been decreasing to minimize the total energy dissipation. This implies that the energy dissipation per unit length of the MOR is much larger than that of the transform, i.e., that the transform is very weak in shear resistance compared to the MOR. This is a somewhat peculiar thing. It has been believed that this weakness might be due to serpentinization of the mantle peridotite beneath the transform due to injection of seawater to the mantle through the fracture zone; serpentine is the hydrous phase of olivine, which shows a stable sliding frictional property. Practice 2.2.2 Let the energy dissipation of a ridge and a transform fault, per unit length, be Er and Et, respectively (Fig. 23). Let the angle between the ridge strike and the normal to the transform fault be γ. Prove that when the total energy dissipaiton is minimum, sinγ = Et/Er holds. Then, when Et is much smaller than Er, the transform and the ridge-axis becomes perpendicular to each other. 2.2.2 Trench-trench transform Three basic types of trench-trench transform faults exists (Fig. 24). Others are mirror images of these. The length of the transform either (a) does not change, (b) decreases, or (c) increases. Practice 2.2.3 What type is the transform between the Tonga and Vanuatu trenches (Fig. 25) among the above three types? 2.2.3 Continental transform There are transforms in the land area. The San Andreas fault system, US, the Alpine Fault, New Zealand, and the North Anatolian fault, Turkey, are such examples. They have been developed in some tectonic situations. They also look very weak. Their origin and evolution have not been elucidated yet. 2.3 Convergent boundary The convergent boundary is divided into the subduction zone and the collision zone (Fig. 26). In the former, the oceanic plate is consumed beneath the upper plate from the deep-sea trench, and the shallow sesmicity is associated with it (Fig. 27). The latter realizes when a continent riding on an oceanic plate impinges into a subduction zone and collision occurs. This is generally believed to occur because a continental plate has a lower density than that of the asthenosphere (see later discussion in 2.3.2). 2.3.1 Subduction zone

Morphology and tectonic elements The oceanic plate is subducting beneath the overriding plate. This corresponds to the downwelling current in a convection scheme. As the morphology of the gentle slope of the ocean floor away from the MOR represents the birth and cooling of a plate, the morphology of the subduction zone represents the destruction of a plate into the asthenosphere. They are the outer-rise, trench, accretionary prism, forearc (+basin), volcanic arc, and backarc (+basin) (Fig. 28). The aseismic front is defined by the location of the downdip limit of interplate earthquakes. The volcanic front is defined by the line marking the seaward limit of volcanoes. The subducted part of the oceanic plate is called a slab. 2.3.1.1 Earthquakes Three brothers Earthquakes that occur in the subduction zone are grouped into three types: (a) interplate earthquakes at the thrust zone, (b) earthquakes within the oceanic plate/slab, and (c) earthquakes within the overriding plate. I cal them "three brothers" [Fig. 29, (a) eldest, (b) middle, and (c) youngest brothers]. (a) Interplate earthquakes Slip vectors and plate motion The eldest brothers occur at the plate interface, between the subducting plate and the overriding plate; focal mechanisms of these interplate earthquakes are of reverse-fault type, one of the nodal planes, representing a fault plane, is dipping shallow (Fig. 30). Such a focal mechanism is often called thrust-type. Slip vectors of focal mechanisms (Fig. 8) represent the motion of the subducting plate to the overriding plate. In the western Pacific, they provide important sources of information on relative plate motions. In fact, the motions of the Philippine Sea and Okhotsk plates have been determined using slip vectors of interplate earthquakes around these plates (Fig. 31, see 3. Plate kinematics). Seismic coupling Mode of occurrence of interplate earthquakes shows a great variety from place to place and temporarily. Although these interplate earthquakes occur due to the relative plate motion and the mechanical coupling between the subducting and overriding plates, we do not understand yet why mode of occurrence, i.e., mode of release of the plate motion, varies so much. We would not be able to say we have understood the eldest brothers until we can explain why they do not occur in some places and/or why they occur only after a long period of quiescence, such as the 2004 Sumatra-Andaman earthquake. Lubrication Considering the frictional strength due to stress operating normal to the fault plane, it becomes difficult for interplate earthquakes to occur at depths greater than ~10 km because the strength becomes too large, compared with the available tectonic stress of ~100 MPa. We need some mechanisms to weaken the faults at the plate interface in subduction zones. The most plausible mechanism would be high fluid pressure in the fault plane opposing the normal stress (Fig. 32). The effective sτress σn* = σn

- Pw, where σn is the stress normal to the fault plane and Pw is the pore fluid pressure, is reduced to nearly zero if Pw is large. The pore fluid would be provided by the free water dehydrated from the altered basalts in the subducting oceanic crust or from the serpentinized slab mantle. Practice 2.3.1 Calculate the shear strength τ of the fault plane at a depth z of 30 km, using the friction of coefficient µ of 0.7, and the density of the crust-mantle ρ of 3000 kg/m^3, from τ = µσn and σn = ρgz. (b) Earthquakes within the oceanic plate/slab Trench - outer rise earthquakes The middle brothers occur firstly beneath the trench - outer rise within the oceanic plate prior to subduction. The oceanic plate bends in this place to subduct into the asthenosphere; the profile of seafloor topography in this region is explained by bending of an elastic plate (Fig. 33). The trench - outer rise earthquakes are grouped into shallow normal fault-type and deep reverse fault-type (Figs. 34 and 35). These types correspond to stresses due to bending. The maximum depth of the reverse fault earthquakes becomes large as the age of the plate becomes old (Fig. 35). Intraslab earthquakes The middle brothers mostly occur within the slab. They are divided into intermediate-depth earthquakes (30-60 km ~ 250 km depth, Fig. 36) and deep earthquakes (300-600 km)(fig. 37). At these depths, the confining pressure becomes large, and they cannot occur without some weakening mechanisms, similarly to the eldest brothers. At intermediate-depths, a very plausible mechanism of weakening is dehydration of hydrous minerals in the crust (altered/metamorphosed basalts) and/or mantle (serpentine). As dehydration occurs, fluid pressure in the pores at the grain boundaries suddenly increases. This pore pressure sustains the confining pressure, reducing the effective stress, and makes it possible for earthquakes to occur. This is called dehydration embrittlement (Fig. 38. The oceanic crust is generally metamorphosed by the hydrothermal circulation at the MOR. This provides a mechanism for intraslab earthquakes in the subducting oceanic crust. If the deeper portion of the oceanic plate is metamorphosed into serpentine, dehydration and associated earthquakes occur in the slab mantle. Along with the intraslab events in the crust, these constitute a double seismic zone (Fig. 39). Reverse fault earthquakes at the trench - outer rise occur around a depth of ~50 km. This depth is large enough to prevent earthquake occurrence. Dehydration of serpentinite is also proposed as a possible mechanism for the deep reverse faulting beneath the trench - outer rise. In this sense, a pair of the normal and reverse fault earthquakes in this area is nothing but a shallow version of the double seismic zone (see Seno and Yamanaka, 1996; Yamasaki and Seno, 2003 for further studies). Dehydration is completed around a 250 km depth, and at larger depths, phase transformation such as olivine to spinel is a likely mechanism of causing deep earthquakes. Seismicity combining both of interplate earthquakes and intraslab earthquakes is called a Wadati- Benioff zone. (c) Earthquakes within the overriding plate

The youngest brothers occur due to stress accumulation within overriding plates. Interaction of plates at the plate boundaries and variation of the vertical stress due to change in the crust/mantle structure determine the stress levels of the overriding plate. They are then modified by heterogeneities of structure, volatiles, temperature, and material properties in the upper plate. At present, it is difficult to know where stresses are concentrated to the critical values from observations or by what mechanism such stresses have been concentrated. Active faults are the places where the youngest brothers have happened in the Holocene - Quaternary (Fig. 40) and their fault-types are controled by the stresses stated above (Fig. 41); however, they sometimes occur where no active faults have appeared on the surface (e.g., the 2000 Western Tottori and 2004 Chuetsu earthquakes). GPS measures only displacements and strains. Even if strains are large, stress might not be necessarily large. In-situ continuous measurements of stresses at crustal depths would be the most promising method to detect dangerous places. However, they have not been realized yet (see Seno, Syntheses of regional stress fields of the Japanese islands, 1999 for further studies). 2.3.1.2 Volcanism Volcanism occurs in subduction zones (Fig. 42). The volcanic front is situated above a subducting slab at a depth of ~100 km and larger. The asthenosphere circulates in the mantle wedge beneath the overriding plate in association with the subduction. Upwelling an downwelling currents occur, but the existence of the overriding plate prohibits reaching the shallow depth. The temperature of the asthenosphere might be also slightly smaller than that beneath the MOR due to the cooling effect of a cold slab. Therefore, the asthenosphere does not cross the (dry) solidus in the subduction zone. However, dehydration from the subducting slab provide water to the mantle wedge and reduce the solidus of the mantle rocks significantly (wet solidus). In association of the circulation, the asthenosphere can cross the wet solidus. In this case, not only the upwelling current, but also the downwelling current may cut the solidus. This is called compressional melting. 2.3.1.3 Accretion Another important phenomenon, which occurs in subduction zones, is accretion of sediments beneath the inner trench wall. This is one of the processes that form continents. There are two modes of accretion: one is offscaping (Fig. 43) and another is underplating. The sediments above the decollement are offscraped, and those below the decollement are underplated or subducted. Accretion is presently occurring along the Nankai Trough, for example (Fig. 44). This place has attracted many geologists who study land areas of the Japanese islands, because the islands have been made of accretionary prisms since Paleozoic time (Fig. 45). 2.3.2 Collision zone Suture zone and ophiolites A belt of seismicity runs along the mountain belt within a continent. The seismicity along the Himalayan mountains and Zagros mountains are typical examples. Geology of these mountain belts shows that different continents are juxtaposed along these belts. This can occur if a continent is resting on an oceanic plate and impinges into the subduction zone (Fig. 46). The place of the former trench is

called a "suture zone". In the suture zone, a set of rocks composed of oceanic sediments, igneous rocks, and ultramafic rocks, which might be fragments of an oceanic plate, is often seen; they are called "ophiolites". Fontal thrust and accretion The location of active thrusting in the collision zone is however located further seaward within the colliding continent. This means a large-scale accretion has been occurring and the plate boundary has migrated toward the colliding plate. For example, in the Himalayas, the presently active thrust boundary is located at the Himalayan Frontal Thrust, about three hundreds of km south of the Indus- Zangbo Suture (Fig. 47). Hinterland Collision also induces intense deformation in the wide areas behind the suture zone like in the Tibet and SE Asia (Fig. 48). This indicates the stress level at the collision boundary is larger than that in the thrust zone of the subduction zone.