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Earthquake Mechanisms and
Plate Tectonics
Seth Stein and Eryn Klosko
Northwestern University, Evanston, Illinois, USA
1. Introduction
Earthquake seismology has played a major role in the development of our current understanding of global plate tectonics
and in making plate tectonics the conceptual framework used
to think about most large-scale processes in the solid Earth.
During the dramatic development of plate tectonics, discussed
from the view of participants by Uyeda (1978, and this volume),
Cox (1973), and Menard (1986), the distribution of earthquakes
provided some of the strongest evidence for the geometry of
plate boundaries and the motion on them (e.g., Isacks et al.,
1968). More than thirty years later, earthquake studies retain a
central role, as summarized here.
Because earthquakes occur primarily at the boundaries
between lithospheric plates, their distribution is used to map
plate boundaries and their focal mechanisms provide information about the motion at individual boundaries.
Plate boundaries are divided into three types (Fig. 1).
Oceanic lithosphere is formed at spreading centers, or midocean ridges, and is destroyed at subduction zones, or trenches.
Ridge
Oceanic plate
Thus, at spreading centers plates move away from the
boundary, whereas at subduction zones the subducting plate
moves toward the boundary. At the third boundary type,
transform faults, plate motion is parallel to the boundary. The
slip vectors of the earthquakes on plate boundaries, which
show the motion on the fault plane, re¯ect the direction of
relative motion between the two plates.
The basic principle of plate kinematics is that the relative
motion between any two plates can be described as a rotation on
a sphere about an Euler pole (Fig. 2). Speci®cally, at any point
along the boundary between plates i and j, with latitude and
longitude , the linear velocity of plate j with respect to plate i is
v ji ˆ !ji r
…1†
the usual formulation for rigid body rotations in mechanics.
The vector r is the position vector to the point on the boundary,
Z
N
v12
Trench
Euler vector
Greenwich
Meridian
Fracture zone
Transform
fault
Lithosphere
12
r
Continental
plate
Y
Magnetic
anomalies
Euler pole
X
Asthenosphere
FIGURE 1 Plate tectonics at its simplest. Plates are formed at ridges
and subducted at trenches. At transform faults, plate motion is parallel
to the boundaries. Each boundary type has typical earthquakes.
INTERNATIONAL HANDBOOKOF EARTHQUAKE AND ENGINEERING SEISMOLOGY, VOLUME 81A
Copyright # 2002 by the Int'l Assoc. Seismol. & Phys. Earth's Interior Committee on Education.
All rights of reproduction in any form reserved.
FIGURE 2 Geometry of plate motions. At any point r along the
boundary between plate i and plate j, with geopraphic latitude and
longitude , the linear velocity of plate j with respect to plate i is
vji ˆ !ji r. The Euler pole at latitude and longitude is the
intersection of the Euler vector !ji with the Earth's surface.
ISBN: 0-12-440652-1
69
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Stein and Klosko
and !ji is the rotation vector or Euler vector. Both are de®ned
from an origin at the center of the Earth.
The direction of relative motion at any point on a plate
boundary is a small circle, a parallel of latitude about the Euler
pole (not a geographic parallel about the North Pole!). For
example, in Figure 3a the pole shown is for the motion of
plate 2 with respect to plate 1. The ®rst-named plate ( j ˆ 2)
moves counterclockwise about the pole with respect to the
second (i ˆ 1). The segments of the boundary where relative
motion is parallel to the boundary are transform faults. Thus,
transforms are small circles about the pole and earthquakes
occurring on them should have pure strike-slip mechanisms.
Other segments have relative motion away from the boundary,
Rotation pole
21
Plate 1
Plate 2
Spreading
ridge
and are thus spreading centers. Figure 3b shows an alternative
case. The pole here is for plate 1 ( j ˆ 1) with respect to plate 2
(i ˆ 2), so plate 1 moves toward some segments of the
boundary, which are subduction zones. Note that the ridge
and subduction zone boundary segments are not small circles.
The magnitude, or rate, of relative motion increases with
distance from the pole, since
jv ji j ˆ j!ji jjrj sin …2†
where is the angle between the Euler pole and the site (corresponding to a colatitude about the pole.) Thus, although all
points on a plate boundary have the same angular velocity, the
linear velocity varies.
If we know the Euler vector for any plate pair, we can write
the linear velocity at any point on the boundary between the
plates in terms of the local E±W and N±S components by a
coordinate transformation. With this, the rate and azimuth of
plate motion become
q
2
EW 2
rate ˆ jv ji j ˆ …vNS
…3†
ji † ‡ …v ji †
azimuth ˆ 90
tan
1
vNS
ji
vEW
ji
!
…4†
such that azimuth is measured in degrees clockwise from North.
Given a set of Euler vectors with respect to one plate, those
with respect to others are found by vector arithmetic. For
example, the Euler vector for the reverse plate pair is the
negative of the Euler vector
Transform
(a)
!ij ˆ
Rotation pole
!ji
…5†
Euler vectors for other plate pairs are found by addition
!jk ˆ !ji ‡ !ik
12
…6†
so, given a set of vectors all with respect to plate i, any Euler
vector needed is found from
Plate 1
Plate 2
Subduction
zone
Transform
(b)
FIGURE 3 Relationship of motion on plate boundaries to the Euler
pole. Relative motion occurs along small circles about the pole;
the rate increases with distance from the pole. Note the difference the
sense of rotation makes: !ji is the Euler vector corresponding to the
rotation of plate j counterclockwise with respect to i.
!jk ˆ !ji
!ki
…7†
For further information on plate kinematics see an introductory text such as Cox and Hart (1986). As discussed there,
motions between plates can be determined by combining three
different types of data from different boundaries. The rate of
spreading at ridges is given by sea-¯oor magnetic anomalies,
and the directions of motion are found from the orientations
of transform faults and the slip vectors of earthquakes on
transforms and at subduction zones. As is evident, earthquake
slip vectors are only one of three types of plate motion
data available. Euler vectors are determined from the relative
motion data, using geometrical conditions. Since slip vectors
and transform faults lie on small circles about the pole, the pole
must lie on a line at right angles to them (Fig. 3). Similarly, the
Earthquake Mechanisms and Plate Tectonics
rates of plate motion increase with the sine of the distance from
the pole. These constraints make it possible to locate the poles.
Determination of Euler vectors for all the plates can thus
be treated as an overdetermined least-squares problem, and
the best solution found using the generalized inverse to derive
global plate motion models (Chase, 1972; Minster and Jordan,
1978; DeMets et al., 1990, 1994). Because these models use
magnetic anomaly data, they describe plate motion averaged
over the past few million years.
New data have become available in recent years due to the
rapidly evolving techniques of space-based geodesy. These
techniques (Gordon and Stein, 1992) (very long baseline radio
interferometry (VLBI), satellite laser ranging (SLR), the
global positioning system (GPS), and DORIS (similar to GPS,
but using ground transmitters)) use space-based technologies
to measure the positions of geodetic monuments to accuracies
of better than a centimeter, even for sites thousands of kilometers apart. Hence measurements of positions over time yield
relative velocities to precisions almost unimaginable during
the early days of plate tectonic studies. A series of striking
results, ®rst with VLBI and SLR (e.g., Robbins et al., 1993),
and now with GPS (Argus and He¯in, 1995; Larson et al.,
1997), show that plate motion over the past few years is
generally quite similar to that predicted by global plate motion
model NUVEL-1A. This agreement is consistent with the
prediction that episodic motion at plate boundaries, as re¯ected in occasional large earthquakes, will give rise to steady
motion in plate interiors due to damping by the underlying
viscous asthenosphere (Elsasser, 1969). As a result, the
earthquake mechanisms can be compared to the plate motions
predicted by both global plate motion models and space-based
geodesy.
2. Oceanic Spreading Center Focal
Mechanisms
Earthquake mechanisms from the mid-ocean ridge system
re¯ect the spreading process. Figure 4 schematically shows a
portion of a spreading ridge offset by transform faults. Because
new lithosphere forms at the ridges and then moves away, the
relative motion of lithosphere on either side of a transform is in
opposing directions. The direction of transform offset, not the
spreading direction, determines whether there is right or left
lateral motion on the fault. This relative motion, de®ned as
transform faulting, is not what produced the offset of the ridge
crest. In fact, if the spreading at the ridge is symmetric (equal
rates on either side), the length of the transform will not change
with time. This is a very different geometry from a transcurrent
fault, where the offset is produced by motion on the fault and
the length of the offset between ridge segments would increase
with time.
The model is illustrated by focal mechanisms. Figure 5a
shows a portion of the Mid-Atlantic Ridge composed of
71
Ridge
Strike-slip fault
(left lateral)
Normal
fault
Fracture zone
Transform
No seismicity
No seismicity
Transform
Normal
fault
Strike-slip fault
(right lateral)
Ridge
FIGURE 4 Possible tectonic settings of earthquakes at an oceanic
spreading center. Most events occur on the active segment of the
transform and have strike-slip mechanisms consistent with transform
faulting. On a slow spreading ridge, like the Mid-Atlantic, normal
fault earthquakes occur. Very few events occur on the inactive
fracture zone.
north±south trending ridge segments, offset by transform
faults, such as the Vema Transform, which trend approximately east±west. Both the ridge crest and the transforms are
seismically active. The mechanisms show that the relative
motion along the transform is right±lateral. Sea-¯oor spreading
on the ridge segments produces the observed relative motion.
For this reason, earthquakes occur almost exclusively on the
active segment of the transform fault between the two ridge
segments, rather than on the inactive extension, known as a
fracture zone. Although no relative plate motion occurs on the
fracture zone it is often marked by a distinct topographic feature, due to the contrast in lithospheric ages across it. Unfortunately, some transform faults named before this distinction
became clear, such as the Vema, are known as ``fracture zones''
along their entire length. Earthquakes also occur on the
spreading segments. Their focal mechanisms show normal
faulting, with nodal planes trending along the ridge axis.
The seismicity is different on fast spreading ridges.
Figure 5b shows a portion of the Paci®c±Antarctic boundary
along the East Paci®c Rise. Here, strike-slip earthquakes occur
on the transforms, but we do not observe the ridge crest
normal faulting events. These observations can be explained
by the thermal structure of the lithosphere, because fast
spreading produces younger and thinner lithosphere than slow
spreading. The axis of a fast ridge has a larger magma
chamber than the slow ridge, and the lithosphere moving away
from a fast spreading ridge is more easily replaced than for a
slow ridge. Thus, in contrast to the axial valley and normal
72
Stein and Klosko
faulting earthquakes on a slow ridge, a fast ridge has an axial
high and absence of earthquakes.
The mechanisms are consistent with the predictions of plate
kinematics. The area in Figure 5a is a portion of the boundary
between the South American and Nubian (West African)
plates. An Euler vector for Nubia with respect to South
America with a pole at 62 N, 37.8 W and a magnitude of
0.328 degrees My 1 predicts that at 0 N, 20 W Africa is
moving N81 E, or almost due East, at 33 mm y 1 with respect
to South America. The Vema is a boundary segment parallel to
this direction, and so is a transform fault characterized by
strike-slip earthquakes with directions of motion along the
trace of the transform. The short segments essentially at right
angles to the direction of relative motion are then spreading
ridge segments. The spreading rate determined from magnetic
anomalies, and thus the slip rate across the transform, is
described by the Euler vector.
15 20⬘
14° N
Vema
10° N
Doldrums
(a)
6° N
– 45° W
– 40° W
–50° S
Eltanin
3. Subduction Zone Focal
Mechanisms
Both the largest earthquakes and the majority of large earthquakes occur at subduction zones. Their focal mechanisms
re¯ect various aspects of the subduction process. Figure 6 is a
composite cartoon showing some of the features observed in
different subduction zones.
Most of the large, shallow, subduction zone earthquakes
indicate thrusting of the overriding plate over the subducting
lithosphere. The best such examples are the two largest ever
recorded: the 1960 Chilean (M0 2.7 1030, Ms 8.3) and 1964
Alaskan (M0 7.5 1029, Ms 8.4) earthquakes. These were impressive events; in the Chilean earthquake 24 m of slip occurred on
a fault 800 km long along-strike and 200 km long down-dip.
Smaller, but large, thrust events are characteristic. For example,
Figure 7a shows the focal mechanisms of large shallow earthquakes along a portion of the Peru±Chile Trench, where the
Nazca Plate is subducting beneath the South American Plate. The
mechanisms along the trench show thrust faulting on fault planes
with a consistent geometry; parallel to the coast, which corresponds to the trench axis, with shallow dips to the northeast.
These thrust events directly re¯ect the plate motion. At a point
on the trench (17 S, 75 W), global plate motion model
NUVEL-1A (DeMets et al., 1994) predicts motion of the Nazca
plate with respect to South America at a rate of 68 mm y 1 and
an azimuth of N76 E. The direction of motion is toward the
trench, as expected at a subduction zone. The major thrust
earthquakes at the interface between subducting and overriding
plates thus directly re¯ect the subduction, and slip vectors from
their focal mechanisms can be used to determine the direction
of plate motion. The rate of subduction is harder to assess.
Although the rate can be computed from global plate motion
models or space geodesy, not all of the plate motion is always
Bending earthquakes
Small earthquakes
- Few, small
–54° S
Great thrust
earthquakes
Normal fault earthquakes
Udintsev
–58° S
(b)
– 145° W
– 140° W
- Few, large
- e.g., 1933 Sanriku, 1965 Rat Island,
1977 Indonesia
- Not observed everywhere
–135° W
Deep seismic zone
FIGURE 5 Maps contrasting faulting on slow and fast spreading
centers. (a) The slow Mid-Atlantic ridge has earthquakes both on the
active transform and ridge segment. Strike-slip faulting on a plane
parallel to the transform azimuth is characteristic. On the ridge segments, normal faulting with nodal planes parallel to the ridge trend is
seen. (b) The fast East Paci®c Rise has only strike-slip earthquakes
on the transform segments. Mechanisms from Engeln et al. (1986),
Huang et al. (1986), and Stewart and Okal (1983).
- Often, but not always
- e.g., 1960 Chile,
1964 Alaska
Intermediate
earthquakes
- Near slab top
660
km
- Either single or double
- Either downdip compression
or downdip extension
- Dip may vary considerably
- Depth may vary considerably
“Composite” subduction zone
FIGURE 6 Schematic of some of the features observed at subduction zones. Not all features are seen at all subduction zones.
Earthquake Mechanisms and Plate Tectonics
73
Nazca–South America Plate Boundary Zone
290°
.
300°
.
.
.
.
South
American
Plate
.
.
–10° S
.
.
.
.
.
.
For
e
n
la
.
r
Th
us
.
.
.
d
Andes
t
Be
.
lt
.
.
.
–20° S
(a)
.
.
.
.
(b)
le
Stab erica
m
th A
30 –40 mm/y locked
18 –33 mm/y stable sliding
Sou
land
Fore Belt
st
Thru
lano
m/y
0m
5–1
Altip
arc
Fore
ch
Tren
late
ca P
Motion
Naz
Plate
GPS Site
Motion
NUVEL-1A
77 mm/y
.
.
.
Nazca
Plate
10– 15 mm/y
shortening
68 –77 mm/yr net convergence
FIGURE 7 (a) GPS site velocities relative to stable South America
(Norabuena et al., 1998), and selected earthquake mechanisms in
the boundary zone. Rate scale is given by the NUVEL-1A vector.
(b) Cross-section across Andean orogenic system showing velocity
distribution inferred from GPS data.
released seismically in earthquakes (Kanamori, 1977). In this
case, the seismic slip rate estimated from seismic moments can
be only a fraction of the real plate motion. Nonetheless, it is
useful to determine the seismic slip rate to assess the fraction
of seismic slip, as it re¯ects the mechanics of the subduction
process. It is also interesting to know how this seismic slip varies
as a function of time and position along a subduction zone.
Figure 6 also shows other types of shallow subduction zone
earthquakes. An interesting class of subduction zone earthquakes result from the ¯exural bending of the downgoing
plate as it enters the trench. Precise focal depth studies show a
pattern of normal faulting in the upper part of the plate to a
depth of 25 km and thrusting in its lower part, between 40
and 50 km. These observations constrain the position of the
neutral surface separating the upper extensional zone from
the lower ¯exural zone, and thus provide information on the
mechanical state of the lithosphere. Occasionally, trenches are
the sites of large normal fault earthquakes (e.g., Sanriku 1933
and Indonesia 1977). There has been some controversy
whether to interpret these earthquakes as bending events in
the upper ¯exural sheet or as ``decoupling'' events showing
rupture of the entire downgoing plate due to ``slab pull.''
The deeper earthquakes, which form the Wadati±Benioff
zone, go down to depths of 700 km within the downgoing slab.
Their mechanisms provide important information about the
physics of the subduction process. The essence of the process
is the penetration and slow heating of a cold slab of lithosphere
in the warmer mantle. This temperature contrast has important
consequences. The subducting plate is identi®ed by the locations of earthquakes in the Wadati±Benioff zone below the
zone of thrust faulting at the interface between the two plates.
Earthquakes occur to greater depths than elsewhere because
the slab is colder than the surrounding mantle. The mechanisms of earthquakes within the slab similarly re¯ect this phenomenon. The thermal evolution of the downgoing plate and
its surroundings is controlled by the relation between the rate
at which cold slab material is subducted and that at which it
heats up, primarily by conduction as it equilibrates with the
surrounding mantle. In addition, adiabatic heating due to the
increasing pressure with depth and phase changes contribute.
Numerical temperature calculations show that the downgoing plate remains much colder than the surrounding mantle
until considerable depths, where the downgoing slab heats up
to the ambient temperature. Comparison of calculated temperatures, the observed locations of seismicity, and images
from seismic tomography shows that the earthquakes occur in
the cold regions of the slab. The thermal structure also helps
explain their focal mechanisms. The force driving the subduction is the integral over the slab of the force due to the
density contrast between the denser subducting material and
the density of ``normal'' mantle material outside. This force,
known as ``slab pull,'' is the plate driving force due to subduction. Its signi®cance for stresses in the downgoing plate
and for driving plate motions depends on its size relative to the
resisting forces at the subduction zone. There are several such
forces. As the slab sinks into the viscous mantle, material must
be displaced. The resulting force depends on the viscosity of
the mantle and the subduction rate. The slab is also subject to
drag forces on its sides and resistance at the interface between
the overriding and downgoing plates. The latter, of course, is
often manifest as the shallow thrust earthquakes.
One way to study the relative size of the negative buoyancy
and resistive forces is to use focal mechanisms to examine
the state of stress in the downgoing slab. Earthquakes
above 300 km generally show stress axes corresponding to
extension directed down the slab dip, whereas those below
74
Stein and Klosko
300 km generally show downdip compression. A proposed
explanation is that there are two basic processes operating:
near the surface the slab is being extended by its own weight;
at depth the slab begins to ``run into'' stronger material and
downdip compression occurs. Another crucial effect may be
buoyancy due to mineral phase changes that occur at different
depths in the cold slab and in the surrounding mantle.
Numerical models of stress in downgoing slabs, using these
assumptions, can reproduce the shallow down-dip tension and
deep downdip compression (Vassiliou, 1984; Bina, 1996).
Finally, it is worth noting that not all features shown in the
schematic (Fig. 6) have been observed at all places. For
example, the dips and shapes of subduction zones vary substantially. Some show double planes of deep seismicity; some
do not. Even the very large thrust earthquakes, considered
characteristic of subduction zone events, are not observed in
all subduction zones. In recent years, considerable effort has
been made to understand such variations.
4. Diffuse Plate Boundary
Earthquake Focal Mechanisms
Although the basic relationships between plate boundaries and
earthquakes apply to continental as well as oceanic lithosphere,
the continents are more complicated. The continental crust is
much thicker, less dense, and has very different mechanical
properties from the oceanic crust. Because continental crust and
lithosphere are not subducted, the continental lithosphere
records a long, involved tectonic history. In contrast, the oceans
record only the past 200 million years. One major result of these
factors is that plate boundaries in continents are often diffuse,
rather than the idealized narrow boundaries assumed in the rigid
plate model, which are a good approximation to what we see in
the oceans. The initial evidence for this notion comes from the
distribution of seismicity and the topography, which often
imply a broad zone of deformation between the plate interiors.
EU
NA
JF
AR
CO
CA
IN
AF
PH
PA
SA
NZ
AU
SC
AN
FIGURE 8 Comparison of the idealized rigid plate geometry to the broad boundary zones implied by seismicity,
topography, or other evidence of faulting. Fine stipple shows mainly subaerial regions where the deformation has
been inferred from seismicity, topography, other evidence of faulting, or some combination of these. Medium
stipple shows mainly submarine regions where the nonclosure of plate circuits indicates measurable deformation; in
most cases these zones are also marked by earthquakes. Coarse stipple shows mainly submarine regions where the
deformation is inferred mainly from the presence of earthquakes. These deforming regions form wide plate
boundary zones, which cover about 15% of the Earth's surface. The precise geometry of these zones, and in some
cases their existence, is under investigation. Plate motions shown are for the NUVEL-1 global relative plate motion
model. Arrow lengths are proportional to the displacement if plates maintain their present relative velocity for
25 My. Divergence across mid-ocean ridges is shown by diverging arrows. Convergence is shown by single arrows
on the underthrust plate. (After Gordon and Stein, 1992.)
Earthquake Mechanisms and Plate Tectonics
75
belt, and into the stable interior of the South American continent. The GPS site velocities are relative to stable South
America, so if the South American plate were rigid and all
motion occurred at the boundary, they would be zero. Instead,
they are highest near the coast and decrease relatively
smoothly from the interior of the Nazca plate to the interior of
South America. Figure 7b shows an interpretation of these
data. In this, about half of the plate convergence (30±
40 mm y 1) is locked at the plate boundary thrust interface,
causing elastic strain that is released in large interplate trench
thrust earthquakes. Another 18±30 mm y 1 of the plate motion
occurs aseismically by smooth stable sliding at the trench. The
rest occurs across the sub-Andean fold-and-thrust belt, causing
permanent shortening and mountain building, as shown by the
inland thrust fault mechanisms. Comparison of strain tensors
derived from GPS and earthquake data shows that the shortening rate inferred from earthquakes is signi®cantly less than
indicated by the GPS, implying that much of the shortening
occurs aseismically. The focal mechanisms also indicate some
deformation within the high Andes themselves. There may be
some (at most 5±10 mm y 1) motion of a forearc sliver distinct
from the overriding plate, a phenomenon observed in some
areas where plate convergence is oblique to the trench, making
earthquake slip vectors at the trench trend between the trenchnormal direction and the predicted convergence direction
(McCaffrey, 1992).
Another broad plate boundary zone is the Paci®c±North
America boundary in western North America. Figure 10 shows
the boundary zone, in a projection about the Euler pole. The
relative motion is parallel to the small circle shown. Thus the
This effect is especially evident in continental interiors, such
as the India±Eurasia collision zone in the Himalayas or the
Paci®c±North America boundary zone in the Western US. Plate
boundary zones (Fig. 8), indicated by earthquakes, volcanism,
and other deformation, appear to cover about 15% of the
Earth's surface (Gordon and Stein, 1992; Stein, 1993).
Insight into plate boundary zones is being obtained by
combining focal mechanisms with geodetic, topographic, and
geological data. Although plate motion models predict only
the integrated motion across the boundary, GPS, geological,
and earthquake data can show how this deformation varies in
space and time. Both variations are of interest. Possible spatial
variations include a single fault system taking up most of
the motion (e.g., Prescott et al., 1981), a smooth distribution
of motion (e.g., England and Jackson, 1989), or motion taken
up by a few relatively large microplates or blocks (e.g., Acton
et al., 1991; Thatcher, 1995). Each of these possibilities
appears to occur, sometimes within the same boundary zone.
The distribution of the motion in time is of special interest
because steady motion between plate interiors gives rise to
episodic motion at plate boundaries, as re¯ected in occasional
large earthquakes, and in some cases steady creep (Fig. 9). The
detailed relation between plate motions and earthquakes is
complicated and poorly understood and hence forms a prime
target of present studies.
For example, Figure 7a shows focal mechanisms and vectors derived from GPS illustrating the distribution of motion
within the boundary zone extending from the stable interior of
the oceanic Nazca plate, across the Peru±Chile trench to the
coastal forearc, across the high Altiplano and foreland thrust
Displacement
relative to Plate A
Plate boundary zone slip distribution
an
dc
cs
lip
mi
mi
eis
eis
re
pis
no
iso
od
ic s
dic
s
Mi
Mi
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ep
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ic s
re
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eis
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lip
lip
an
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ree
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te
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pla
ad
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?)
ste
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)
(M
erm
g-t
Lo
n
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plate
interior
ree
p
Time
ree
p
Time
Plate boundary
zone of deformation
Rigid
plate
interior
Seismic
Plate B
Plate A
Aseismic
FIGURE 9 Schematic illustration of the distribution of motion in space and time for a strikeslip boundary zone between two major plates (Stein, 1993).
Stein and Klosko
60°
80
°
0°
22
24
0°
260
°
280°
76
.
Tr
80° N
en
.
°
ch
200
.
Alaska 1964
40
°N
San Francisco 1906
F
Jd te
Pla
42 mm/y
Borah Peak
PA - NA
POLE
Basin and
range
Landers
9 mm
SAF
A5
Parkfield
40° N
rm
°N
nsfo
Tra
/y
Northridge
San Fernando
20
60° N
F-NA
NA
Loma Prieta
V PPAA-N-NA
°
220
VJd
JdF-
.
ge
Rid
.
280°
260°
24
0°
GPS Site
Motion
FIGURE 10 Geometry and focal mechanisms for a portion of the North America±Paci®c
boundary zone. Dot-dash line shows small circle, and thus direction of plate motion, about the
Paci®c±North America Euler pole. The variation in the boundary type along its length from
extension, to transform, to convergence, is shown by the focal mechanisms. The diffuse nature
of the boundary zone is shown by seismicity (small dots), focal mechanisms, topography
(1000 m contour shown shaded), and vectors showing the motion of GPS and VLBI sites with
respect to stable North America (Bennett et al., 1999; Newman et al., 1999).
boundary is extensional in the Gulf of California, essentially a
transform along the San Andreas fault system, and convergent
in the eastern Aleutians. The focal mechanisms re¯ect these
changes. For example, in the Gulf of California we see strikeslip along oceanic transforms and normal faulting on a ridge
segment. The San Andreas has both pure strike-slip earthquakes (Park®eld) and earthquakes with some dip-slip motion
(Northridge, San Fernando, and Loma Prieta) when it deviates
from pure transform behavior. The plate boundary zone is also
broad, as shown by the distribution of seismicity. Although the
San Andreas fault system is the locus of most of the plate
motion and hence large earthquakes, seismicity extends as far
eastward as the Rocky Mountains. For example, the Landers
earthquake shows strike-slip east of the San Andreas, and the
Borah Peak earthquake illustrates Basin and Range faulting.
The diffuse nature of the boundary is also illustrated by vectors showing the motion of GPS and VLBI sites with respect to
stable North America. Net motion across the zone is essentially that predicted by global plate motion model NUVEL1A. The site motions show that most of the strike-slip occurs
along the San Andreas fault system, but signi®cant motions
occur for some distance eastward.
Earthquake Mechanisms and Plate Tectonics
5. Intraplate Deformation and
Intraplate Earthquakes
A ®nal important use of earthquake mechanisms is to study the
internal deformation of major plates. Although idealized plates
would be purely rigid, the existence of intraplate earthquakes
re¯ect the important and poorly understood tectonic processes
of intraplate deformation. One such example is the New Madrid
area in the central United States, which had very large earthquakes in 1811±1812. The seismicity of such regions is generally thought to be due to the reactivation of preexisting faults
or weak zones in response to intraplate stresses. Because
motion in these zones are at most a few mm y 1, compared to
the generally much more rapid plate boundary motions, seismicity is much lower (Fig. 10). Similarly, major intracontinental earthquakes occur substantially less frequently than
plate boundary events; recurrence estimates for 1811±1812
type earthquakes average 500±1000 years. Efforts are being
made to combine geodetic data, which indicate deviations
from rigidity, to the earthquake data. For example, comparison
of the velocities for permanent GPS sites in North America
east of the Rocky Mountains to velocities predicted by modeling these sites as being on a single rigid plate shows that the
interior of the North American plate is rigid at least to the level
of the average velocity residual, less than 2 mm y 1 (Dixon
et al., 1996; Newman et al., 1999). Similar results emerge from
geodetic studies of other major plates, showing that plates
thought to have been rigid on geological time scales are quite
rigid on decadal scales. Moreover, geological data suggest that
such intraplate seismic zones may be active for only a few
thousands of years, even though plate motions have been steady
for millions of years. As a result, understanding how these
intraplate seismic zones operate is a major challenge. A special
case of this phenomenon occurs at passive margins, where
continental and oceanic lithosphere join. Although these areas
are in general tectonically inactive, magnitude 7 earthquakes
can occur, as on the eastern coast of North America. Such
earthquakes are thought to be associated with stresses at the
continental margin, including those due to the removal of
glacial loads, which reactivate the faults remaining along the
continental margin from the original rifting.
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Editor's Note
Please see also Chapter 6, Continental Drift, Sea-Floor
Spreading, and Plate/Plume Tectonics, by Uyeda.