Earthquake Mechanisms and Plate Tectonics

7 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 (eg, 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 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, N Z v 12 Fracture zone Oceanic plate Transform fault Ridge Trench Continental plate Greenwich Meridian Euler vector r 12 Y Lithosphere 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 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 v ji ˆ! ji r The Euler pole at latitude and longitude is the intersection of the Euler vector! ji with the Earth's surface INTERNATIONAL HANDBOOK OF EARTHQUAKE AND ENGINEERING SEISMOLOGY, VOLUME 81A ISBN: 0-12-440652-1 Copyright # 2002 by the Int'l Assoc Seismol & Phys Earth's Interior Committee on Education All rights of reproduction in any form reserved 69

70 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, Plate 1 Spreading ridge Rotation pole 21 Plate 2 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 1moves 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 jˆj! ji jjrj sin 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 rate ˆjv ji jˆ v NS ji 2 v EW ji 2 3 azimuth ˆ 90 tan 1! v NS ji v EW ji 2 4 (a) Transform Rotation pole 12 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! ij ˆ! ji 5 Euler vectors for other plate pairs are found by addition! jk ˆ! ji! ik 6 Plate 1 Plate 2 so, given a set of vectors all with respect to plate i, any Euler vector needed is found from! jk ˆ! ji! ki 7 (b) Transform Subduction zone 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 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 71 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 (eg, 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 Strike-slip fault (left lateral) No seismicity Transform Strike-slip fault (right lateral) Ridge Transform Fracture zone No seismicity Ridge Normal fault Normal fault 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, 378 W and a magnitude of 0328 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 (a) 14 N 10 N 6 N 50 S 15 20 45 W Vema 40 W Doldrums 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 (M 0 27 10 30, M s 83) and 1964 Alaskan(M 0 75 10 29,M s 84)earthquakesThesewereimpressive 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 54 S Eltanin Bending earthquakes - Few, small Small earthquakes (b) 58 S Udintsev 145 W 140 W 135 W 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) Normal fault earthquakes - Few, large - eg, 1933 Sanriku, 1965 Rat Island, 1977 Indonesia - Not observed everywhere Deep seismic zone - Either single or double - Either downdip compression or downdip extension - Dip may vary considerably - Depth may vary considerably Composite subduction zone Great thrust earthquakes - Often, but not always - eg, 1960 Chile, 1964 Alaska Intermediate earthquakes - Near slab top 660 km 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 10 S 20 S (a) Plate Motion (b) Nazca Plate GPS Site Motion NUVEL-1A 77 mm/y Nazca Plate Trench Nazca South America Plate Boundary Zone Andes Forearc 5 10 mm/y 30 40 mm/y locked 18 33 mm/y stable sliding 290 300 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 South American Plate Foreland Thrust Belt Altiplano Foreland Thrust Belt 68 77 mm/yr net convergence Stable South America 10 15 mm/y shortening 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 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 (eg, 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 JF NA PA CO CA AF AR IN PH 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-1global 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 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 (eg, Prescott et al, 1981), a smooth distribution of motion (eg, England and Jackson, 1989), or motion taken up by a few relatively large microplates or blocks (eg, 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 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 Plate boundary zone slip distribution Displacement relative to Plate A Rigid plate interior Time Long-term (MY) steady (?) plate motion Minor episodic seismic slip and creep Major fault Plate boundary zone of deformation Major episodic seismic slip and creep Rigid plate interior Minor episodic seismic slip and creep Time Plate B Seismic Aseismic Plate A FIGURE 9 Schematic illustration of the distribution of motion in space and time for a strikeslip boundary zone between two major plates (Stein, 1993)

76 Stein and Klosko 60 80 220 240 260 280 Trench 80 N 200 Alaska 1964 40 N 220 San Francisco 1906 Loma Prieta JdF Plate Borah Peak PA-NA Parkfield Basin and range Landers Northridge SAF Transform V JdF-NA 42 mm/y JdF-NA PA - NA POLE 60 N V PA-NA 59mm/y San Fernando 40 N 20 N Ridge GPS Site Motion 240 260 280 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 NUVEL- 1A 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 77 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 References Acton, GD, S Stein, and JF Engeln (1991) Block rotation and continental extension in Afar: a comparison to oceanic microplate systems Tectonics 10, 501±526 Argus, DF and MB He in (1995) Plate motion and crustal deformation estimated with geodetic data from the Global Positioning System Geophys Res Lett 22, 1973±1976 Bennett, RA, JL Davis, and BP Wernicke (1999) Present-day pattern of Cordilleran deformation in the Western United States Geology 27, 371±374 Bina, CR (1996) Phase transition buoyancy contributions to stresses in subducting lithosphere Geophys Res Lett 23, 3563±3566 Chase, CG (1972) The n-plate problem of plate tectonics Geophys J R Astron Soc 29, 117±122 Cox, A (1973) ``Plate Tectonics and Geomagnetic Reversals'' WH Freeman, San Francisco Cox, A and RB Hart (1986) ``Plate Tectonics: How it Works'' Blackwell Scienti c, Palo Alto DeMets, C, RG Gordon, DF Argus, and S Stein (1990) Current plate motions Geophys J Int 101, 425±478 DeMets, C, RG Gordon, DF Argus, and Stein, S (1994) Effect of recent revisions to the geomagnetic reversal time scale on estimates of current plate motion Geophys Res Lett 21, 2191±2194 Dixon, TH, A Mao, and S Stein (1996) How rigid is the stable interior of the North American plate? Geophys Res Lett 23, 3035±3038 Elsasser, WM (1969) Convection and stress propagation in the upper mantle In: ``The Application of Modern Physics to the Earth and Planetary Interiors'' (SK Runcorn, Ed), pp 223±246 Wiley, New York Engeln, JF, DA Wiens, and S Stein (1986) Mechanisms and depths of Atlantic transform earthquakes J Geophys Res 91, 548±577 England, P and J Jackson (1989) Active deformation of the continents Annu Rev Earth Planet Sci 17, 197±226 Gordon, RG and S Stein (1992) Global tectonics and space geodesy Science 256, 333±342 Huang, PY, SC Solomon, EA Bergman, and JL Nabelek (1986) Focal depths and mechanisms of Mid-Atlantic Ridge earthquakes from body waveform inversion J Geophys Res 91, 579±598 Isacks, B, J Oliver, and LR Sykes (1968) Seismology and the new global tectonics J Geophys Res 73, 5855±5899 Kanamori, H (1977) Seismic and aseismic slip along subduction zones and their tectonic implications In: ``Island Arcs, Deepsea Trenches and Back-arc Basins, Maurice Ewing Series'' (M Talwani and WC Pitman, Eds), Vol III, pp 163±174 American Geophysical Union, Washington, DC Larson, KM, JT Freymueller, and S Philipsen (1997) Global plate velocities from the Global Positioning System J Geophys Res 102, 9961±9981 McCaffrey, R (1992) Oblique plate convergence, slip vectors, and forearc deformation J Geophys Res 97, 8905±8915 Menard, HW (1986) ``The Ocean of Truth: A Personal History of Global Tectonics,'' Princeton Series in Geology and Paleontology (AG Fischer, Ed), Princeton University Press, Princeton Minster, JB and TH Jordan (1978) Present-day plate motions J Geophys Res 83, 5331±5354 Newman, A, S Stein, J Weber, J Engeln, A Mao, and T Dixon (1999) Slow deformation and lower seismic hazard at the New Madrid Seismic Zone Science 284, 619±621 Norabuena, E, L Lef er-grif n, A Mao, T Dixon, S Stein, IS Sacks, L Ocala, and M Ellis (1998) Space geodetic observations of Nazca±South America convergence along the Central Andes Science 279, 358±362 Prescott, WH, M Lisowski, and JC Savage (1981) Geodetic measurements of crustal deformation on the San Andreas,

78 Stein and Klosko Hayward, and Calaveras faults, near San Francisco, California J Geophys Res 86, 10853±10869 Robbins, JW, DE Smith, and C Ma (1993) Horizontal crustal deformation and large scale plate motions inferred from space geodetic techniques In: ``Contributions of Space Geodesy to Geodynamics: Crustal Dynamics, Geodynamics Series'' (DE Smith and DL Turcotte, Eds), Vol 23, pp 21±36 American Geophysical Union, Washington, DC Stein, S (1993) Space geodesy and plate motions In: ``Space Geodesy and Geodynamics, Geodynamics Series'' (DE Smith and DL Turcotte, Eds), Vol 23, pp 5±20 American Geophysical Union, Washington, DC Stewart, LM and EA Okal (1983) Seismicity and aseismic slip along the Eltanin Fracture Zone J Geophys Res 88, 10495± 10507 Thatcher, W (1995) Microplate versus continuum descriptions of active tectonic deformation J Geophys Res 100, 3885±3894 Vassiliou, MS (1984) The state of stress in subducting slabs as revealed by earthquakes analyzed by moment tensor inversion Earth Planet Sci Lett 69, 195±202 Uyeda, S (1978) ``The New View of the Earth'' WH Freeman, San Francisco Editor's Note Please see also Chapter 6, Continental Drift, Sea-Floor Spreading, and Plate/Plume Tectonics, by Uyeda