Download Plate Tectonics: Note 2 T. Seno (Earthquake Res Inst, Univ of Tokyo

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Plate Tectonics: Note 2
T. Seno (Earthquake Res Inst, Univ of Tokyo)
(Revised on November 5, 2008)
2. Plate boundary processes
2.0 Plate boundaries
2.0.1 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.
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 ocean basins, fracture zones, which cut and 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 (buldge) and
deep-sea trench (depression). Another type of the convergent boundary is located along a mountain belt in the continent.
This will be explained later.
Topographic feature of transform faults is not obvious, because transcurrent or strike-slip motion itself does not
produce much 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 caused by faulting (Fig. 5), we expect normal, strike-slip and reverse fault-type
earthquakes along these boundaries.
2.0.2 Representation of faulting
Fault parameters representation
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 strike-slip faulting.
Focal mechanism representation
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
usually 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 Pwaves emanated from the white quadrants show negative first motions, and those from the black quadrants show positive
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first motions (Fig. 7). These polarities are traced back along the rays to the epicenter and are plotted in the focal sphere, in
order 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.
Letting the slip vector direction ψ, φ − ψ, λ and δ have a relation shown in Fig. 6. Then, given ψ, λ and δ, λ can
be determined.
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.
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Altered (metamorphosed) 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 (convergent boundary) and is subducted beneath an overriding plate, the
temperature and the pressure would rise, and the altered basalts dehydrate, and provide released water at the plate boundary
thrust. This would be a cause for occurrence of interplate earthquakes in subduction zones, as will be seen in 2.3. The
released water also is fed into the mantle and crust of the overriding plate. This results in the melting and arc magmas, and
also in the occurrence of intraplate earthquakes.
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 (convergent boundary). The morphological expression of such oceanic
transform faults is a fracture zone.
2.2.1 Ridge-ridge type
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 cross-sections along a fracture zone.
Strike-slip faulting
Focal mechanisms of earthquakes along a transform fault 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 fault. Then a question
arises why such an offset is produced, or, in other words, how a ridge-ridge transform fault 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 fault 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 fault, i.e., that the transform fault 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 a transform fault due to injection of seawater to the mantle through the
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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 fault 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 fault either (a) does not change, (b) decreases, or (c) increases. Case (b) means that such a transform fault is
unstable, and not seen usually.
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 transform faults in the land area. The San Andreas fault, US, the Alpine Fault, New Zealand, and the
North Anatolian fault, Turkey, are such examples. Although they are strike-slip faults in the continent, they are different
from usual intraplate strike-slip faults, because they mark a boundary between two plates. They have been developed into
the present form through some particular tectonic situations.
San Andreas fault
This strike-slip fault is elongated along the Coast Range, western US, actually consists of ~three subparallel faults.
This correctly is called the San Andreas fault system. It is peculiar that it exists within the N. American continent, but
Atwater (1970) explained its origin as follows. There was once a plate called the Farallon plate west of the Pacific plate in
the Pacific Ocean (Fig. 49). The boundary between these plates were marked by the ridge-transform fault system. Around 30
Ma, the easternmost part of the ridge-transform fault encountered the west coast, and as part of the Farallon plate has
disappeared beneath the N. American continent, part of the Pacific plate in touch with the west coast has increased. By
happen, the relative motion between the Pacific and N. American plate is parallel to the west coast (relict trench), the
boundary between these plates became a transform fault.
Although this gross scenario of the kinematic evolution of the San Andreas fault system would be correct, there are
some misterious points in the San Andreas. If the above story is correct, the Pacific and N. American plate boundary should
be located at the former trench of the Farallon plate subduction. In fact, the San Andreas are located inside the coast, i.e., in
the Coast Range (Fig. 50). Furthermore, it is not a single fault, but consists of ~3 faults. This means that the San Andreas
was generated rather inside the upper North American plate in the subduction zone.
Initiation of plate tectonics
It is noticed that a transform fault does not appear in the 2-D scheme of convection of non-temperature dependent
viscosity (Fig.5 of lec.eps.note.1). This may be due to its 2-dimensionality. Only divergent and convergent boundaries
appear in this convection. However, even in 3-D simulations of convection of non-temperature dependent viscosity,
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); divergence occurs in a center of a cell, and convergence occur around the perifery
of the cell. Toroidal velocity fields (divergence free), by which transform faults motions are characterized, do not appear.
We noted in seno.lec1.eps.note that the convection of temperature dependent viscosity becomes a stagnant-lid
convection. In 3-d, the situation is the same, and the tectonic style becomes a stagnant-lid like Venus. When the appearance
of a number of plates occurs like Earth, transform faults appear. This is because if one rigid block move w.r.t. other
neighboring blocks, it must involve a component of transcurrent motion to at least one neighboring plate (Fig. 51); this
becomes a transform fault with toroidal fields. This means that the initiation of plate tectonics is equivalent to the time of
formation of transform faults.
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
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(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
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 (intraslab earthquakes), and (c) earthquakes within the overriding plate
(intraplate earthquakes).
(a) Interplate earthquakes
Slip vectors and plate motion
These earthquakes 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 Plate kinematics lecture by Okino).
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.
Further on this subject, consult Seno (2003), Fractal asperities, invasion of barriers...., a pdf in the reference list of Seno's
web site.
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
operated normal to the fault plane in this case. This 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, which is formed at the MOR,
or dehydrated from the serpentinized slab mantle.
Practice 2.3.1
Calculate the shear strength τ of the fault plane at a depth z of 20 km, using the friction of coefficient µ of 0.6, and the
density of the overriding crust-mantle ρ of 3000 kg/m^3, from τ = µσn and σn = ρgz.
(b) Intrasalab earthquakes
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Trench - outer rise earthquakes
The first type of intraslab earthquakes occur 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).
Intermediate-depth and deep earthquakes
The second type earthquakes occur deeper than the trench and are divided into intermediate-depth earthquakes (a
depth range of 30-60 km ~ 325 km, Fig. 36) and deep earthquakes (a depth range of 325-600 km)(Fig. 37). At these depths,
the confining pressure becomes very large, and they cannot occur without some weakening mechanisms, similarly to the
interplate earthquakes. 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.
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 last type of earthquakes occur due to stresses 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 (consult Seno, 1999 in the web site). They are then modified by strength heterogeneities, affected by
volatiles, temperature, pressure and material properties. At present, it is difficult to know where stresses are close to the
critical values from observations or by numerical methods. GPS measures only displacements and strains. In-situ stress
measurements would neither be usefull to detect dangerous places, because these stresses are superficial, not representing
stresses at crustal depths. Active faults are the places where surface ruptures have happened repeatedly in the Holocene Quaternary (Fig. 40). They may be related to faulting in the basement but still only remaining as surface expressions (e.g.,
no active faults have appeared on the surface when the 2000 Western Tottori, the 2004 Chuetsu, and other recent
earthquakes occurred in Honshu). It is however interesting to note that active faults have characteristic fault-types in each
regions of Honshu (Fig. 41).
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. The asthenosphere circulates in the mantle wedge beneath the overriding plate in association with the
subduction. Upwelling and downwelling currents occur, but the existence of the overriding plate prohibits them from
reaching the very shallow depth. The temperature of the asthenosphere might be also slightly smaller than that of the mantle
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 (it becomes a wet solidus). In association with the circulation, the asthenosphere can cross the
wet solidus. In this case, not only the upwelling current, but also the downwelling current may happen to cut the solidus.
The latter case is called compressional melting.
2.3.1.3 Accretion
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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 is attached to the upper plate as thrust
sheets. Those below the decollement are underplated; some of them might be subducted into deeper part of the subduction
zones but others are attached at the base of the upper plate. This is called underplating.
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 mostly of accretionary prisms
since Paleozoic time (Fig. 45).
2.3.2 Collision zone
Suture zone and ophiolites
A belt of seismicity runs along a 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 trailed by an oceanic plate and impinges into the subduction zone after the
subduction of the oceanic plate (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 within the colliding
plate. For example, in the Himalayas, the presently active thrust boundary is located at the Himalayan Frontal Thrust, ~300
km south of the Indus-Zangbo Suture, which is a relict trench (Fig. 47).
Focal mechanisms of earthquakes occurring in the frontal thrust are of thrust-type, very similar to that of interplate
earthquakes in subduction zones. In fact, they are representing relative motions of two coverging plates, such as the Indian
and Eurasian plates. Sometimes, bending earthquakes occur similarly to subducition zones. Significant difference is that
there is no intermediate or deep earthquakes in collision zones.
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.
What causes collision?
We calculated the average density of a continental plate and see that it is lower than that of the asthenosphere. I
mentioned that it is the reason why the continental plate does not recycle. It was, but, based on fake reasoning. It often
occurs that an oceanic plate is trailing a continent, and the oceanic plate has an excess density than the asthenosphere. If the
continental plate trailed by the oceanic plate has a length L, and the oceanic plate has a length larger than 3L, total density
anomaly is positive to the asthenosphere. It occurs that such an oceanic plate can drag down the continental plate into the
asthenosphere. Once the continental plate is dragged deep into the mantle, phase changes occur in the crust, and basaltic
material turns into higher density materials such eclogite.
It suggests us that a true cause of collision is not excess density of the continental plate. Please consult Seno
(2008), Condisions for a crustal block to be sheared off, a pdf file in the reference list of Seno's web site) for further
reference on this subject.
--------------------------------------------------------------------------------------------References for figures
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