Figure 5.1

      Earthquake activity along the San Andreas fault system (fig. 5.1) reflects brittle accommodation of the crust to the relative motion along the dextral transform boundary between the Pacific and North American plates (see chap. 3). Great (M≈8) earthquakes along the main branch of the San Andreas fault accommodate most of this relative plate motion. These plate-boundary earthquakes rupture the entire 15- to 20-km thickness of the brittle crust with right-lateral offsets of as much as 10 m extending for several hundred kilometers along the fault trace, and they recur along a given section of the San Andreas fault at intervals of several hundred years (Sieh, 1981; Sieh and Jahns, 1984). The 1857 Fort Tejon earthquake in southern California and the 1906 San Francisco earthquake in northern California are only the two most recent such great events along the San Andreas fault (see chap. 6).

      In this chapter, we focus on the smaller, more frequent earthquakes that dominate the seismicity within the San Andreas fault system between recurrences of great, plate-boundary events. Although this interevent seismicity contributes only marginally to relative plate motion, it is symptomatic of processes underlying the earthquake cycle. In particular, the wealth of data generated by frequent smaller events provides important clues to the seismotectonic fabric, kinematics, and state of stress within the brittle crust and, ultimately, to the seismogenic processes common to earthquakes of all sizes within the San Andreas fault system. This persistent, interevent seismicity also captures widespread popular interest because it includes most of the felt earthquakes in California (earthquakes of M≥3 can be locally felt), and the larger of these interevent earthquakes (M=6-7) can cause extensive damage and loss of life when they strike near major population centers.
      We examine here the detailed patterns of earthquake occurrence along the San Andreas fault system recorded by the combined northern and southern California seismograph networks for the 7-yr interval 1980-86. These networks, which had evolved to much their present configuration (fig. 5.2) by mid-1979, enable uniform detection and location of all earthquakes of M≥1.5 throughout the San Andreas fault system and of M≥2 throughout most of California. The telemetered seismic stations in the combined networks approach 550 in number. Signals from the 300 central and northern California stations are recorded and processed at the U. S. Geological Survey's (USGS) offices in Menlo Park; signals from the 250 southern California stations are recorded and processed under a joint USGS-California Institute of Technology (Caltech) effort on the Caltech campus in Pasadena. These dense telemetered networks overlie regional seismic networks operated by the University of California, Berkeley, and Caltech that provide records of M≥3 earthquakes in northern and southern California, respectively, from the early 1930's onward (table 5.1; Hileman and others, 1973; Bolt and Miller, 1975).

Table 5.1 - Number of short-period seismic stations in California

Figure 5.2 - Stations in telemetered seismic networks in California operating during 1986. Dot, single (vertical)-component station; star, multicomponent station.

      After a brief overview of the San Andreas fault system in the context of a broad transform boundary, we focus on the three-dimensional distribution of earthquakes along the San Andreas fault system itself on the basis of a series of detailed seismicity maps and cross sections for the years 1980-86. We then review selected focal mechanisms for the larger of these earthquakes as a guide to the kinematics of seismogenic deformation along the fault system. Finally, we discuss the implications of these seismicity patterns in terms of current tectonic processes along the transform boundary.
      The seismicity maps and cross sections in this chapter, which form the core of our presentation, are largely self-explanatory. The following points, however, deserve special emphasis:

Figure 5.3 - Place names and faults most commonly used in text (see front of book for more complete maps of place names and faults). Faults (dotted where concealed): B, Banning; BP, Big Pine; BS, Bartlett Springs; BZ, Brawley seismic zone; C, Calaveras; CN, Concord; CU, Cucamonga; DV, Death Valley; E, Elsinore; FC, Furnace Creek; G, Garlock; GV, Green Valley; H, Hayward; HG, Hosgri; HRC, Healdsburg-Rodgers Creek; 1M, Imperial; LVC, Long Valley caldera; M, Maacama; MC, Mission Creek; NI, Newport- Inglewood; OT, Ortigalita; PM, Pinto Mountain; PV, Palos Verdes; PV A, Panamint Valley; R, Rinconada; RC, Rose Canyon; SA, San Andreas; SC, San Clemente Island; SG, San Gregorio; SJ, San Jacinto; SN, Sierra Nevada; SNA, Sur-Nacimiento; W, Whittier; WM, White Mountains; WW, White Wolf. Arrows and numbers indicate direction and amount of motion, respectively, of Pacific and Gorda plates with respect to North American plate to the east; red lines indicate 1857 and 1906 ruptures of San Andreas fault.

      Dickinson (1981), among others, emphasized that the San Andreas fault system and the San Andreas transform boundary are not strictly equivalent structures. The latter is more general, incorporating, for example, the concept of temporal evolution of Pacific-North American plate interaction and the recognition that the faults accommodating most of the plate motion have changed over time. In this section, we emphasize that, although great earthquakes along the San Andreas fault system currently accommodate most of the relative plate motion, plate interactions along the transform boundary influence deformation of the brittle crust over a much broader region.
      The breadth of the seismicity pattern in California and western Nevada (fig. 5.4) suggests the lateral extent of deformation associated with the San Andreas transform boundary. Indeed, it corresponds closely to the zone of distributed shear between those plates as interpreted by Ward (1988) from more than 5 years of very long baseline interferometry (VLBI) observations at 20 Western United States stations from 1982 through 1987. Figure 5.4 includes all M≥1.5 events recorded by the telemetered seismic networks in figure 5.2 during the 7-yr interval 1980-86, as well as events recorded by adjacent telemetered networks in Nevada (see Rogers and others, in press). Although details within this seismicity pattern fluctuate from year to year, the broad aspects of the pattern have remained stable for the entire historical record of earthquake occurrence in California (see chap. 6; Ellsworth and others, 1981; Hill and others, in press; Hutton and others, in press).

Figure 5.4 - Locations of 64,000 M≥1.5 earthquakes in California and western Nevada during 1980-86 and mapped Holocene faults (dotted where concealed; major branches of the San Andreas fault system marked in red).

      In outline, the seismicity pattern for California and western Nevada forms a hollow ellipse with its long axis nearly coincident with the transform boundary. This pattern is widest across central California, where it approaches nearly a third of the 1,100-km length of the transform boundary, from the Mendocino triple junction in the north to the head of the Gulf of California at the south. Extended alignments of epicenters within this pattern suggest a coarse structural fabric linking the broad distribution of earthquakes to the transform boundary. Seismicity along the San Andreas fault system itself stands out as a series of subparallel, northwest-trending lineations extending the length of coastal California. The alignment of epicenter clusters along the east side of the Sierra Nevada in eastern California branches northward from the south end of the San Andreas fault system in the Salton Trough only to bend back toward the north terminus of the San Andreas fault system at the Mendocino triple junction in northern California. The Sierra N evada-Great Valley and western Mojave Desert blocks form a broad quiescent region between the San Andreas and eastern California seismicity bands. In southern California, pronounced transverse seismicity belts coincident with the southern margin of the Sierra Nevada and Transverse Ranges, respectively, span this otherwise-quiescent region. A weaker, more diffuse seismicity belt near lat 37° N. spans the Sierra Nevada-Great Valley block in central California, forming a visual, if not structural, link between the San Andreas fault system and the dense cluster of epicenters in eastern California. This major cluster in the eastern Sierra Nevada represents an episode of intense earthquake activity in Long Valley caldera and vicinity that began in 1978 and has persisted to the present (Van Wormer and Ryall, 1980; Hill and others, 1985b).
      The displacement pattern associated with earthquakes throughout California and western Nevada is broadly consistent with deformation under a coherent, regional stress field dominated by plate-boundary interaction along a northwest-striking, dextral transform boundary (Hill, 1982). Strike-slip focal mechanisms with right-lateral slip on northwest- to north-north west-striking planes, for example, are common through most of the region. Regional variations within this pattern include a tendency toward normal slip on northerly-striking planes along the western margin of the extensional Basin and Range province in eastern California, and toward reverse slip on easterly-striking planes in the Transverse Ranges of southern California. Compressional deformation perpendicular to the San Andreas fault within the Coast Ranges, however, represents an important deviation from this regional pattern.

      Sections of the San Andreas fault system stand out on seismicity maps as a network of northward-branching alignments of epicenters through the central and northern Coast Ranges and as subparallel lineations of clusters of epicenters that branch northward from the south end of the Imperial fault toward the Transverse Ranges in southern California (fig. 5.4). One of the most remarkable aspects of the seismicity pattern associated with the fault system, however, is the nearly complete absence of earthquake activity down to even the smallest magnitudes (M≈1.5) along those sections of the fault that have ruptured with the largest historical earthquakes, the great (M≈8) 1857 and 1906 earthquakes (see figs. 5.3, 5.4). The southernmost section of the San Andreas fault, from Indio to the Salton Sea, also lacks microseismicity, although no large earthquake has ruptured this section in the past 200 yr. These quiescent ("locked") segments of the fault stand in sharp contrast to the segments marked by persistent linear concentrations of small to moderate earthquakes.
      This dual expression of the fault system evident on current seismicity maps apparently reflects fundamental differences in the long-term behavior of the respective segments. In particular, seismic activity along the "locked" segments of the main trace of the San Andreas fault may be limited to the recurrence of major earthquakes at intervals of 100 to 300 yr accompanied by immediate fore shock and aftershock sequences, and these segments may remain quiescent for most of the interevent time associated with the cycle between great earthquakes (Ellsworth and others, 1981). In contrast, those segments with persistent microearthquake activity probably seldom, if ever, rupture with great earthquakes, although they may be capable of generating earthquakes as large as M≈6.
      Aseismic creep also characterizes and is largely confined to those fault segments along the San Andreas fault system that show persistent micro earthquake activity (Wesson and others, 1973; Schulz and others, 1982). Creep is most pronounced along the central California segments of the fault system, where average creep rates match the long-term displacement rates of 32 to 34 mm/yr. Louie and others (1985) documented creep along sections of the seismically active fault segments in the Salton Trough, and Astiz and Allen (1983) documented creep along a section of the Garlock fault that is marked by microearthquake activity. The creep rates in these two areas, however, are more than an order of magnitude less than the long-term deformation rates.
      In the following subsections, we consider the 1980-86 seismicity along and adjacent to the major sections of the San Andreas fault system in more detail. We begin with the Mendocino triple junction in the north and move southward, generalizing slightly Allen's (1968) subdivision of the fault system into four major sections of contrasting seismic behavior: (1) the quiescent 1906 break and subparallel branches, (2) branches forming the central California active (creeping) section, (3) the quiescent 1857 break, and (4) branches forming the southern California active section south of the Transverse Ranges.

      The three lithospheric plates that dominate the modern tectonics of coastal California meet at the Mendocino triple junction, which is marked by a dense cluster of epicenters just off Cape Mendocino (fig. 5.5A; see chap. 3). North of this triple junction, oblique subduction dominates, with the eastern margin of the Gorda plate (the southernmost section of the Juan de Fuca plate) slipping beneath the North American plate north-north-eastward at a rate of 30 to 50 mm/yr (Wilson, 1986). South of the triple junction, the Pacific plate is moving past the North American plate along the San Andreas transform boundary on a heading of 35°-38° W. of N. at a rate of approximately 50 mm/yr (DeMets and others, 1987).

Figure 5.5 - Seismicity in the Mendocino triple junction area, northern California. A, Earthquake locations, showing major branches of the San Andreas fault system in red; faults dotted where concealed. Magnitude symbols shown in explanation are scaled with enlargement of cross section. B, Depth sections perpendicular to N. 75° W. trend of the Mendocino Fracture Zone (A-A' and B-B' outlined in fig. 5.5A). C, Depth sections parallel to N. 75° W. trend of the Mendocino Fracture Zone (C-C' and D-D' outlined in fig. 5.5A).

      The dense lineation of epicenters that trends west-northwest from Cape Mendocino corresponds closely to the Mendocino Fracture Zone (MFZ) and the Pacific-Gorda plate boundary but follows a slightly more northerly trend. Details of how the San Andreas fault ties into the triple junction, however, are unclear. The trace of the San Andreas fault lies off shore north of Point Arena, and the broad area of seismic quiescence south of the triple junction offers few clues to the kinematics of this northernmost segment of the fault.
      The conspicuous linear zone of epicenters marking a northeast-trending slice across the southwest corner of the Gorda plate is the aftershock zone of the M=7.2 Eureka earthquake of November 8, 1980. This was the largest earthquake to occur in California during the interval 1980-86. The N. 53° E. trend of its aftershock zone agrees closely with the strike of the left-lateral slip plane inferred from the focal mechanism of the main shock, which was located about a fourth of the way downstrike from the northeastern end of the aftershock zone (Eaton, 1989). The aftershocks from this earthquake died off within a few months after the main shock-a notably short aftershock sequence for an earthquake of this size.
      A third, more diffuse group of earthquakes in the vicinity of the triple junction shows little tendency to concentrate in linear zones. These epicenters form an irregular zone with the greatest concentration in the vicinity of Cape Mendocino, gradually dying off with distance to the north, east, and west. To the south, the seismicity dies off abruptly across the MFZ and its landward extension.
      The three-dimensional aspects of this triple junction seismicity are illustrated by four cross sections (figs. 5.5B, 5.5C). The two cross sections perpendicular to the N. 75° W. trend of the MFZ (fig. 5.5B) compare the distribution of focal depths within the submarine section of the Gorda plate adjacent to the triple junction (A-A') with that within the adjacent, subducted section of the Gorda plate and the overlying North American plate (B-B'). Earthquake hypocenters along the southern margin of the submarine Gorda plate define a dense, vertically elongate zone beneath the MFZ that dips 70°-75° N. and extends to depths of nearly 35 km. Earthquakes north of this zone cluster in a somewhat less dense, wedge-shaped core outlined by a northward shallowing of maximum focal depths accompanied by a northward deepening of minimum focal depths that converge at a depth of about 20 km (Eaton, 1989). Overlying this relatively dense wedge is a more diffuse distribution of hypocenters, the northernmost of which represent the 1980 M=7.2 event and adjacent aftershocks. The absence of seismicity south of the MFZ indicates that the northeast corner of the Pacific plate behaves as a rigid block in its interaction with the younger, thinner, and internally deforming Gorda plate (Wilson, 1986; Eaton, 1989).
      In the onshore cross section (B-B', fig. 5.5B), the intense seismicity associated with the MFZ loses most of its expression. Here, the southern margin of the subducted Gorda plate is marked only by isolated clusters of hypocenters at depths of 10 and 25 km. To the north, however, a band of hypocenters concentrated at about a depth of 20 km corresponds closely to the base of the wedge-shaped distribution beneath the submarine section of the Gorda plate, including the downward flexing of this band as it approaches the southern margin of the Gorda plate. South of the landward extension of the MFZ, the seismicity shallows rather abruptly, reflecting the edge of the subducted Gorda plate beneath the North American plate and the rather thin seismogenic crust associated with the San Andreas fault system to the south.
      The cross section parallel to and including the MFZ and its landward extension (D-D', fig. 5.5C) reveals that the dense seismicity cluster along the MFZ near the triple junction tapers westward along the MFZ with a wedge-like geometry to a 20-km-deep band of hypocenters, much the same as it does to the north (cross sec. A-A', fig. 5.5B). (The pronounced linear concentration of hypocenters at 15-km-depth beneath the submarine Gorda plate in cross sections A-A' and D-D' represents the default focal depth for the more poorly located earthquakes beyond the perimeter of the onshore seismic network.) The landward extension of the MFZ shows up only weakly as a diffuse scattering of hypocenters extending to a small, isolated cluster of 25-km-deep hypocenters some 50 km east of the triple junction (Δ=190 km, cross sec. D-D') and, possibly, as far as a handful of 30- to 50-km-deep hypocenters 100 km east of the triple junction (Δ=260 km, cross sec. D-D'). Focal depths of the shallow seismicity in the northern Coast Ranges are confined to the upper 10 km of the crust. Farther east, however, focal depths increase again to depths of 35 to 40 km beneath the north end of the Great Valley. An east-dipping quiescent band, about 10 km thick, appears to separate the seismicity associated with the MFZ from that beneath the northern Coast Ranges and the Great Valley to the east. The geometry of this band suggests that it may somehow be related to the geometry of the subducted Gorda plate beneath the western margin of the North American plate.
      The parallel cross section immediately to the north (C-C', fig. 5.5C) reinforces the impression that the distribution of hypocenters carries an image of subducted-Gorda-plate geometry. Maximum focal depths increase systematically from 25 to 30 km beneath the submarine Gorda plate to nearly 80 km beneath the southern Cascade volcanoes. Although the easternmost of these deep earthquakes are small and few, their locations are well constrained (Cockerham, 1984; Walter, 1986). As in the section of the fault to the south, a seismically quiescent, east-dipping band appears to separate earthquakes within the Gorda plate from those in the overlying North American plate.

      Aside from a light scattering of epicenters about the trace of the San Andreas fault through the San Francisco peninsula and the Santa Cruz Mountains to the south, the rupture zone of the 1906 earthquake is nearly aseismic. This pattern has persisted not only through the period 1980-86 shown in figure 5.6 but also at least since the mid-1930's, when instrumental data became available for reliable earthquake locations in the area (see Bolt and Miller, 1975; Hill and others, in press). It was interrupted, however, by the M=7.1 Lorna Prieta earthquake on October 17, 1989, which ruptured the southernmost 45 km of the 1906 break (cross sec. L-L', fig. 5.7; see chap. 6). The cluster of epicenters along the fault just west of San Francisco coincides closely with Boore's (1977) estimate of the epicentral location for the 1906 main shock. We note that the greatest offsets along the 1906 rupture occurred north of the epicenter along the stretch of the fault that now shows the lowest seismicity (see Thatcher and Lisowski, 1987).

Figure 5.6 - Seismicity in the northern (A) and central (B) Coast Ranges. Major branches of the San Andreas fault marked in red; faults dotted where concealed. Earthquakes within rectangles are plotted in corresponding depth sections in figures 5.7 and 5.8. C, Calaveras fault; CN, Concord fault; G, The Geysers; SF Bay, San Francisco Bay; SJB, San Juan Bautista; SSB, Suisun Bay; D, Dixon; V, Vacaville; W, Winters.

Figure 5.7 - Longitudinal depth sections along the San Andreas fault system in the northern and central Coast Ranges. Bar above section L-L' indicates rupture extent of the October 17, 1989, Lorna Prieta M=7.1 earthquake (see chap. 6). SSB, Sui sun Bay. See figure 5.6 for locations of sections and explanation of symbols, which are scaled with scale change of cross sections.

      The pair of subparallel epicenter lineations through the northern Coast Ranges east of the 1906 break closely follow the Rodgers Creek-Healdsburg-Maacama fault zone (west) and the Green Valley-Bartlett Springs fault zone (east). In both lineations, the epicenters tend to cluster along the eastern margins of these 2- to 3- km-wide fault zones, which are characterized by multiple strands of subparallel, curvilinear fault traces (see maps at front of book; Dehlinger and Bolt, 1984). These fault zones are essentially colinear with the Hayward and Calaveras faults to the south, although an aseismic interval coincident with the eastern arm of the San Francisco Bay (San Pablo and Suisun Bays) obscures the connection between these branches of the fault system. Most of the earthquakes defining the pair of subparallel lineations through the northern Coast Ranges are small (M≤3), and, indeed, these fault zones were not recognized as seismically active branches of the San Andreas system until after the northern Coast Ranges section of the telemetered network (fig. 5.2) became operational in late 1979.
      Dense clusters of epicenters just south of Clear Lake define a northeast trending pattern transverse to and midway along these northern two branches of the San Andreas fault system. The southwesternmost of these clusters represents microearthquake activity associated with the Geysers geothermal field (Eberhart-Phillips and Oppenheimer, 1984; Oppenheimer, 1986). The cluster just to the northeast underlies the Clear Lake volcanic field, which last erupted about 10 ka (Donnelly-Nolan and others, 1981). Scattered clusters farther to the northeast suggest a tenuous link between this Geysers-Clear Lake trend and the north-south-trending lineation of epicenters along the axis of the north end of the Great Valley.
      In longitudinal cross section (H-H', fig. 5.6), earthquakes occurring along the two northern branches of the San Andreas fault system are moderate in number and rather evenly distributed except for the dense, shallow cluster beneath the Geysers-Clear Lake area. The depth to the base of the continuously seismogenic crust shows considerable relief, deepening to between 12 and 13 km south of the Geysers-Clear Lake area and shallowing to only 5 km beneath and just north of the geothermal field. Farther north, the base of the seismogenic crust deepens gradually to a maximum of about 10 km. Note the isolated cluster of small but well-located events at depths of 13 to 18 km beneath Clear Lake, just north of the shallowest depths to the base of the continuously seismogenic crust. Another isolated cluster of deep earthquakes located beneath Suisun Bay have focal depths as great as 15 to 25 km. The more numerous earthquakes along branches of the fault system south of Suisun Bay paint in a dense distribution of hypocenters throughout the upper 10 to 15 km of the crust.
      Although the two transverse cross sections across the northern Coast Ranges (E-E', F-F', fig. 5.8A) show the concentration of hypocenters around the Rodgers Creek and Green Valley branches of the San Andreas fault, they provide no hint of the location of the main branch of the San Andreas fault that ruptured in 1906. The southern of these two cross sections (F-F') includes the dense, shallow cluster of events associated with the Geysers- Clear Lake activity, as well as the cluster of deeper events beneath the eastern margin of the Coast Ranges. Focal depths in the latter cluster, which falls at the northeast end of the Geysers-Clear Lake lineation and at the south end of the north-south-trending lineation beneath the Great Valley, range from 10 to 25 km, a depth range common to earthquakes near the north end of this Great Valley seismicity lineation (cross sec. D-D', fig. 5.5C). Maximum focal depths increase by several kilometers from west to east in both cross sections. Farther east, they increase abruptly to depths of 25 km or so along the Great Valley lineation.

Figure 5.8 - Transverse depth sections across the San Andreas fault system in the northern (A) and central (B) Coast Ranges. See figure 5.6 for locations of sections and explanation of symbols, which are scaled with enlargement of cross sections. Faults: BS, Bartlett Springs; C, Calaveras; GV, Greenville; H, Hayward; M, Maacama; R, Rinconada; SA, San Andreas; SG, San Gregorio; SNA, Sur-Nacimiento. BV, Bear Valley on cross section J-J'. FRT, Franciscan terrane.

      Densely aligned epicenters mark the nearly straight, creeping section of the San Andreas fault that separates the south end of the 1906 break near San Juan Bautista from the north end of the 1857 break near Cholame (fig. 5.6B). Densely aligned epicenters follow the Calaveras fault northward to a point just east of the south tip of the San Francisco Bay, where the Hayward fault branches to the west and the Greenville fault zone branches to the east. Few epicenters fall along the northward extension of the Calaveras fault beyond this branching point, although a diffuse cluster of epicenters coincides with the right-stepping offset between the north end of the Calaveras fault and the Concord fault (fig. 5.6A). This dilatational offset was the site of pronounced earthquake swarms in June 1970 and August 1976 (Lee and others, 1971; Weaver and Hill, 1978/79).
      Although these dense alignments of epicenters coincide closely with the mapped surface trace of the San Andreas fault system as first documented by Eaton and others (1970a), the coincidence is not everywhere perfect (fig. 5.6B). In the region where the Calaveras fault branches from the main trace of the San Andreas fault, for example, the densely aligned epicenters appear to be systematically displaced 3 to 4 km westward of the surface trace of the San Andreas fault and a somewhat smaller distance eastward of the surface trace of the Calaveras fault. Much of this apparent offset results from a strong contrast in rock type and P-wave velocity across the faults that is not adequately accounted for in routine hypocenter locations (Mayer-Rosa, 1973; Pavoni, 1973; Spieth, 1981). When the hypocenter locations are recalculated using a more appropriate, two-dimensional structural model, these offsets are much reduced but not completely eliminated. The remaining offsets reflect deviation of the faults from the vertical, with the San Andreas fault dipping 70° W. (Pavoni,1973; Spieth, 1981) and the Calaveras fault dipping 80° E. (Reasenberg and Ellsworth, 1982).
      Wesson and others (1973) recognized the close correlation between active creep and persistent microseismic activity along the 200-km-Iong section of the San Andreas fault north of Parkfield and proposed that this correlation may hold for other branches of the fault in central California as well. Allen (1981) and Schultz and others (1982) pointed out that this correlation holds for the Calaveras-Hayward fault system, but creep measurements have yet to be made on the subparallel Rodgers Creek-Healdsburg-Maacama and Green Valley-Bartlett Springs faults. Creep is the dominant process for shallow slip along the central section of the San Andreas fault, and geodetic measurements spanning this section of the fault indicate that the long-term slip rate of 32 mm/yr along the fault accommodates nearly all of the local plate motion. Because it appears that little, if any, shear strain is accumulating in the blocks on either side of the fault, most seismologists believe that this section of the fault is unlikely to rupture in a great earthquake (see chap. 7).
      The creeping section of the fault system, however, has produced several moderate earthquakes during historical time (see chap. 6). The most recent of these events, which occurred along the right-branching segments northeast of the San Andreas fault, where creep rates average several millimeters per year, include (1) the M=5.9 Coyote Lake earthquake of August 6, 1979, and the M=6.2 Morgan Hill earthquake of April 24, 1984, which ruptured adjacent 20-km-Iong segments of the Calaveras fault south of its junction with the Hayward fault (Reasenberg and Ellsworth, 1982; Bakun and others, 1984); and (2) the M=5.5 and 5.8 Livermore events of January 24 and 27, 1980, which ruptured a 20-km-Iong stretch of the Greenville fault north of Livermore (Bolt and others, 1981).
      The most noteworthy sequence of moderate earthquakes along the central section of the San Andreas fault involve the five virtually identical M=6 events that have ruptured the same 30-km-Iong stretch near Parkfield at nearly 22-year intervals since 1881 (Bakun and McEvilly, 1984). This stretch of the fault is defined by a 1-km right-stepping offset on the south and a 5° W. bend on the north; it coincides with the transition between the south end of the creeping section of the fault and the north end of the 1857 break (see Bakun and Lindh, 1985, fig. 1). The most recent of these characteristic Parkfield earthquakes occurred in 1966, and an intensive monitoring experiment is underway to capture a detailed instrumental record of the next Parkfield earthquake, which is predicted to occur sometime within a 10-year window centered on 1987-88 (Bakun and Lindh, 1985).
      The scattered seismicity within the Coast Ranges surrounding the San Andreas fault system is distinctly more intense in the Franciscan terrane east of the fault than in the granitic Salinian block to the west. The large, dense cluster of epicenters along the eastern margin of the Coast Ranges adjacent to the south end of the creeping section represents the aftershocks of the M=6. 7 Coalinga earthquake of May 2, 1983, and the M=5.7 Kettleman Hills earthquake of August 4, 1985, both of which involved reverse slip on northwest-striking planes subparallel to the adjacent section of the San Andreas fault (Stein and King, 1984; Eaton, 1989). Scattered clusters of epicenters within the Franciscan terrane show a crude northwest-trending alignment with the southwest edge of the Coalinga- Kettleman Hills aftershock zone and the Ortigalita fault, the north end of which passes beneath the San Luis Reservoir (LaForge and Lee, 1982). This weakly defined lineation is essentially colinear with the Greenville fault, east of the San Francisco Bay (fig. 5.4).
      The Salinian block west of the central section of the San Andreas fault and east of the Sur-Nacimiento fault forms a broad, nearly aseismic swath along the west flank of the Coast Ranges. These two seismically active fault zones separate the granitic Salinian block from the Franciscan terrane on either side (see chap. 3). The Rinconada fault, within the Salinian block, appears to be nearly aseismic except, possibly, toward the south where it approaches the Sur-Nacimiento fault zone. The small cluster of epicenters just east of the midpoint of the Rinconada fault represents a persistent spot of microearthquake activity at depths of 8 to 10 km near the San Ardo oil field (Poley, 1988).
      In the cross section along the central section of the San Andreas fault (L-L', fig. 5.7), the actively creeping section of the fault shows up as a densely mottled distribution of hypocenters within the upper 12 to 15 km of the crust. The density of hypocenters within this creeping section tends to decrease with depth, and the denser clusters generally are concentrated at depths of less than 5 to 8 km. The base of the seismogenic zone undulates about an average depth of some 13 km beneath most of the creeping section, but it deepens to 15 km beneath both the northern and southern transitions to the locked sections of the fault. In contrast to the creeping section of the fault, the sparse seismicity associated with the locked segments that ruptured in 1906 (north) and 1857 (south) tends to be concentrated toward the deeper parts of the seismogenic crust. Note that the 1989 Lorna Prieta M=7.1 earthquake ruptured the 45-km-Iong section of the San Andreas fault with a pronounced U-shaped gap in shallow earthquakes immediately north of the creeping section of the fault (cross sec. L-L', fig. 5.7; see chap. 6).
      The cross sections transverse to the central San Andreas fault system (G-G', fig. 5.8A; I-I', J-J', K-K', fig. 5.8B) reveal the seismically active, creeping branches of the fault as narrow, near-vertical hypocenter distributions coincident with the fault plane. The broadened distribution in cross sections I-I' and J-J' (fig. 5.8B) reflects the oblique projection of earthquakes along the Calaveras fault zone and the clustering northeast of the fault in the Bear Valley region (Ellsworth, 1975), respectively. These transverse cross sections also emphasize the quiescence of the Salinian block relative to the Franciscan terrane on either side (note, however, the isolated cluster of deep events beneath San Ardo in the Salinian block in cross sec. K-K'), and the fairly uniform depth of 12 to 15 km to the base of the seismogenic zone that persists throughout the central Coast Ranges. As in the area farther north, however, maximum focal depths increase rather abruptly to 25 km beneath the eastern margin of the Coast Ranges and the Great Valley. This increase in focal depth is particularly pronounced beneath the dense cluster of hypocenters associated with the 1983 Coalinga earthquake and its aftershocks (cross sec. G-G', fig. 5.8B).

      The 1857 rupture of the San Andreas fault began near Parkfield at the south end of the creeping section of the fault and propagated southeastward along the straight segment, through the Carrizo Plain and around the Big Bend near Tejon Pass, and thence along the relatively straight, east-southeastward trend of the Mojave segment to Cajon Pass, where the San Jacinto fault branches to the south (figs. 5.3, 5.9A). Right-lateral offsets associated with this great earthquake generally decreased from 9 m along the Carrizo Plain segment, through 6 m around Fort Tejon, to 3-4 m along the Mojave segment (Sieh, 1978).

Figure 5.9 - Seismicity in the southern Coast Ranges and Transverse Ranges. A, Earthquake locations, showing major branches of the San Andreas fault system in red; faults dotted where concealed. Magnitude symbols shown in explanation are scaled with enlargement of cross sections. C, Cholame; CP, Cajon Pass; P, Parkfield; PI, Palmdale; SB, Santa Barbara; 5MB, Santa Monica Bay; TP, Tejon Pass. B, Depth sections outlined in figure 5.9A. Faults: B, Banning; G, Garlock; MC, Mission Creek; N.Br.SA, northern branch of the San Andreas; PM, Pinto Mountain; SA, San Andreas, SJ, San Jacinto; WW, White Wolf.

      The pronounced bends in the San Andreas fault at either end of the east-southeast-trending Mojave segment involve strong structural complexities and clusters of persistent seismic activity. Both bends, for example, spawn major left-lateral faults that form conjugate sets to the San Andreas system. Sykes and Seeber (1985) proposed that these two major bends in the San Andreas fault system represent large-scale asperities that exert a strong influence on the rupture patterns of great earthquakes along the San Andreas fault in southern California. The San Andreas fault appears to maintain its integrity as a single, more or less continuous zone through the 30° bend at Fort Tejon. Convergence resulting from the pronounced counterclockwise cant of the Mojave segment of the San Andreas fault (N. 66° W.) with respect to the average N 37° W. orientation of the transform boundary largely accounts for the crustal uplift and shortening expressed in the Transverse Ranges.
      The straight Carrizo Plain segment is, except for a small cluster of events near Simmler (lat 35°25' N., long 120°00' W.), almost completely aseismic, much like the Point Arena segment of the 1906 break. The straight part of the Mojave segment also is seismically very quiet. The southernmost 80 km of the 1857 rupture zone produces few earthquakes, forming a sinuous lineation around the fault. Again, an analogy can be drawn with the 1906 rupture zone: The southernmost 80 to 100 km of the 1906 rupture, which exhibited less slip than the rupture zone farther north, also has a slightly higher background seismicity rate than that to the north. The Mojave earthquakes are temporally clustered (Sauber and others, 1983) and, because of their reverse focal mechanisms and displacement from the surface trace of the San Andreas, are thought to be on secondary fault structures rather than on the San Andreas fault itself.
      Within the area of the Big Bend of the San Andreas fault near Tejon Pass, the level of seismicity is much higher than in the adjoining regions. Much of this activity is associated with the Pleito- White Wolf fault system, which abuts the San Andreas fault in the northern bend, some 40 km west of the junction with the Garlock-Big Pine faults (fig. 5.9A; cross sec. N-N', fig. 5.9B; see fig. 5.12). The White Wolf fault ruptured in 1952 with the M=7.7 Kern County earthquake, accompanied by left-oblique reverse slip on a southeast-dipping fault plane (Richter, 1958; Stein and Thatcher, 1981). This is the largest earthquake to occur in California since the M≈8 San Francisco earthquake. In contrast to the quiescent 1906 rupture, however, the White Wolf fault continues to be marked by persistent aftershocks of the 1952 Kern County earthquake. The Garlock and Big Pine faults are essentially quiescent within 40 to 50 km of the San Andreas, although two small clusters of epicenters form a nearly symmetrical pattern on either side of the junction of these sinistral faults with the San Andreas.
      Seismicity increases markedly near the south end of the 1857 rupture zone, where the San Jacinto fault and the east-west-striking Cucamonga fault, which forms the south front of the central Transverse Ranges, intersect the San Andreas fault. A bulbous lobe of epicenters extends westward along the Cucamonga fault from this junction. Scattered epicenters fill in the wedge of the Transverse Ranges between the Cucamonga and San Jacinto faults. In figure 5.9A, note the tight lineation of epicenters that appears to follow the trace of the Cucamonga fault westward to its intersection with the Elsinore fault (Chino branch). The significance of this lineation is unclear because the Cucamonga fault presumably dips north beneath the central Transverse Ranges (Morton, 1987).
      The cross section along the Mojave segment of the San Andreas fault (M-M', fig. 5.9B) shows a wide variation in focal depths. The few events along the straight section of the Mojave segment are strongly clustered around depths of 10 km, with the maximum focal depth always above 15 km and almost no shallow earthquakes. This pattern is similar to the concentration of hypocenters in the lower half of the seismogenic crust beneath the locked 1906 segment (compare with cross sec. L-L', fig. 5.7). In contrast, the earthquakes at Tejon Pass cover the full depth range from 0 to 25 km. At the southeast end of the 1857 rupture zone at Cajon Pass, maximum focal depths increase with the seismicity level to a maximum of 20 km. In this section, there are no shallow (less than 5 km deep) events.
      Note, in particular, that focal depths of more than 20 km are broadly associated with the convergent tectonics of the Transverse Ranges; they are not limited to the seismicity clusters in the vicinity of the bends in the San Andreas fault system as might be inferred from cross section M-M' (fig. 5.9B). Cross sections N-N' and 0-0' (fig. 5.9B), for example, illustrate that the seismogenic crust reaches thicknesses of 20 to 25 km beneath the Santa Barbara Channel and the western Transverse Ranges, as well as beneath the Tehachapi Mountains to the east.

      Southeast of the 1857 rupture, the San Andreas fault splays into several branches associated with intense seismicity in the Banning-San Gorgonio area. Like the Tejon Pass bend in the 1857 rupture zone, the San Gorgonio bend spawns a major left-lateral fault (the Pinto Mountain fault) conjugate to the San Andreas system. Unlike the situation at Tejon Pass, however, the San Andreas fault at San Gorgonio splays into a complex pattern of branching and intersecting fault segments. South of San Gorgonio, the San Andreas fault reconverges into a single strand and bends again to the more southeasterly trend that characterizes the southern section of the fault system.
      This section of the fault system south of the Transverse Ranges is transitional from oblique spreading along the axis of the Gulf of California to the obliquely convergent strike-slip displacement that dominates deformation along the continental section of the San Andreas transform boundary to the north. Several major strike-slip faults run west of and subparallel to the main strand of the San Andreas fault south of the Transverse Ranges. These faults, which are considered part of the San Andreas system and include the Imperial, San Jacinto, and Elsinore faults, accommodate a significant proportion of the plate-boundary motion. The Imperial and San Jacinto faults, in particular, have produced more moderate earthquakes than any other segment within the fault system (see chap. 6; Hanks and others, 1975).

      The most intense seismicity along the main trace of the southern section of the San Andreas fault is associated with the San Gorgonio bend and is concentrated between the two principal branches of the San Andreas fault: (1) the Mission Creek fault, or northern branch of the San Andreas; and (2) the Banning fault, which runs nearly due west from the south end of the Mission Creek fault toward an ambiguous junction with the San Jacinto fault just south of San Bernardino (see fig. 5.3A and maps at front of book). Neither strand forms a continuous structure through the bend. This San Gorgonio seismicity cluster produced numerous M=5.0-6.5 earthquakes in the 1930's and 1940's, and in 1986 it produced the M_L=5.6 North Palm Springs earthquake, which involved dextral strike-slip displacement on the north-dipping Banning fault (Jones and others, 1986). The background seismicity in this area is the highest in southern California, but it is distributed throughout a volume and cannot be clearly associated with any fault. To the west, seismicity associated with the Banning cluster abuts the dense lineation of epicenters coincident with the northernmost segment of the San Jacinto fault. Nicholson and others (1986) suggested that much of the seismicity within the upper 10 km of the crust in this cluster involves left-lateral slip on a series of northeast-striking faults; however, Jones (1988) pointed out that the evidence for northeast-trending lineations of epicenters within the Banning cluster is less than clear.
      Diffuse seismicity extends northward from the Banning cluster into the San Bernardino Mountains and eastward into the Pinto Mountains, with no clear lineations along the sinistral Pinto Mountain fault. Indeed, a diffuse, north-south-trending lineation of epicenters seems to cut directly across the Pinto Mountain fault from the west-central Pinto Mountains. Two M=5.2 earthquakes (see events 75, 76, fig. 5.11A) with right-lateral strike-slip planes parallel to this trend occurred at the north end of this zone in 1975 and 1979. Somewhat farther south, however, a broad, east-west-trending lineation appears to coincide with the Blue Cut fault. Even farther south, a second broad lineation extends eastward from near the junction of the Banning and Mission Creek branches, although this lineation does not coincide with a mapped fault.
      The southernmost section of the San Andreas fault, the Indio segment, which extends from the junction of the Banning and Mission Creek branches southward to the end of the San Andreas at the Salton Sea, has been almost completely aseismic in historical time. At the north end of this segment, periodic swarms of small (M≤4) earth- quakes a few kilometers east of the San Andreas appear to occur on small northeast-trending structures (for example, Norris and others, 1986; Jones, 1988). The sparse background seismicity is also offset a few kilometers to the east from the surface trace of the San Andreas. Although the possibility of systematic offsets related to P-wave-velocity contrasts across the fault has not been investigated in detail, the observed offset seems too large to be explained entirely by lateral velocity contrasts.
      Although it has not ruptured with a major earthquake in historical time, the aseismic Indio segment of the San Andreas fault seems to have much in common with the 1857 and 1906 rupture zones. Sieh (1986) presented geologic evidence for at least four major ruptures along the Indio segment since A.D. 1000; the last occurred approximately 300 yr ago. Unlike the two major locked sections, however, the south end of the Indio segment adjacent to the Salton Sea shows minor aseismic creep (Louie and others, 1985) and has shown episodes of sympathetic slip accompanying M≈6 earthquakes on the Imperial fault and the southern section of the San Jacinto fault (Sieh, 1982). Not only is the Indio segment aseismic, but also the entire Coachella block extending from the San Andreas fault on the northeast to the crest of the San Jacinto Mountains on the southwest.
      A cross section of hypocenters along the southern branch of the San Andreas fault (M-M', fig. 5.9B) shows that the earthquakes associated with the bend at San Gorgonio are among the deepest in southern California, with maximum focal depths approaching 25 km. The maximum focal depths deepen southward along the San Andreas fault from about 12 km beneath the Mojave segment to 25 km beneath the San Gorgonio fault. At the south end of the San Gorgonio area, however, maximum focal depths abruptly decrease to 10 km. This shallowing of seismicity is associated with a shift in the most concentrated seismicity from between the two segments (Mission Creek and Banning) of the San Andreas to east of the Mission Creek fault. The sparse seismicity of the Indio segment is limited to depths of 5 km or less.

      Although the southernmost section of the San Andreas fault is almost completely aseismic, associated subparallel faults are extremely active. These faults are marked by the three bold north-south- to northwest-trending alignments of epicenters that dominate the seismicity pattern within the San Andreas fault system south of the Transverse Ranges (fig. 5.10A), from east to west: (1) the Brawley seismic zone (Johnson, 1979), defined by a dense, spindle-shaped cluster of epicenters connecting the north end of the Imperial fault and the south end of the Indio segment of the San Andreas fault; (2) the northwestward alignment of densely clustered epicenters along the San Jacinto fault zone, which appears to branch from the northern section of the Imperial fault; and (3) the northwestward alignment of more diffusely clustered epicenters along the Elsinore fault, which appears to branch from somewhere near the south end of the Imperial fault.

Figure 5.10 - Seismicity along the southern section of the San Andreas fault system. A, Earthquake locations, showing major branches of the San Andreas fault system in red; faults dotted where concealed. Magnitude symbols shown in explanation are scaled with enlargement of cross sections. BZ, Brawley seismic zone; LB, Long Beach; MSJ, Mount San Jacinto; SB, Santa Barbara, SBd, San Bernardino. B, Depth sections outlined in 5.10A. Faults: CU, Cucamonga; NI, Newport-Inglewood; W, Whittier.

Figure 5.11 - Focal mechanisms for larger earthquakes. A, California. 5MB, Santa Monica Bay; SS, San Simeon. Numbers refer to table 5.2. B, Coalinga-Kettleman Hills region (events 47-50, fig. 5.11A). Letters and numbers refer to table 5.3. Circle size increases with magnitude from 3.5 to 6.7.

      The Brawley seismic zone and the cluster of epicenters at the south end of the Imperial fault (coincident with the Cerro Prieto volcanic-geothermal field in Mexico) represent the two northernmost in the series of small spreading centers offset by right-lateral transform faults that characterize oblique spreading in the Gulf of California (Lomnitz and others, 1970; Johnson and Hill, 1982). The Imperial fault itself, which is marked by a scattered alignment of epicenters, serves as the transform fault between these two small spreading centers. The M=7.1 El Centro earthquake ruptured the entire length of the Imperial fault in 1940, and the M=6.6 Imperial Valley earthquake of 1979 ruptured the north two-thirds of the fault; intensity data suggest that moderate earthquakes (5.5<M<6.3) in 1906, 1915, 1917, and 1927 may also have been located on the Imperial fault (Johnson and Hill, 1982). Most of the aftershocks associated with the 1979 Imperial Valley earthquake were concentrated in the south half of the Brawley seismic zone, which was first recognized because of the many earthquake swarms it produced from 1973 through mid-1979 (Hill and others, 1975; Johnson 1979; Johnson and Hutton, 1982). Many of the individual swarm sequences, as well as individual clusters of events in the aftershock sequence, defined lineations transverse to the strike of the Imperial fault and the long axis of the Brawley seismic zone. Most earthquakes within the Brawley seismic zone have strike-slip focal mechanisms; thus, kinematically, these transverse lineations represent conjugate structures to the dominant north-northwestward trend of the Imperial-Brawley fault system.
      Irregular clusters of epicenters mark the San Jacinto fault zone, which runs along the southwest base of the Santa Rosa and San Jacinto Mountains. These clusters tend to be concentrated near bends and junctions within the complex set of multiple fault strands that form the surface expression of this fault zone. In several places, particularly within the southern and northern sections of the fault zone, epicenters define linear concentrations that tend to be closely aligned with mapped fault traces. The San Jacinto fault zone has produced at least 10 earthquakes of M=6.0-6.6 since 1890, the most recent of which were the M=6.2 earthquake of 1954, the M=6.6 Borrego Mountain earthquake of 1968, and the M=6.6 Superstition Hills earthquake of 1987. Thatcher and others (1975) pointed out that this series of historical M>6 earthquakes along the San Jacinto fault zone has left two seismic gaps: one along the northern 40 km of the fault, and the other along a 20-km-Iong stretch of the central section of the fault zone (the Anza gap). The Anza gap shows up in figure 5.10A as a relatively quiescent stretch of the fault zone between two dense clusters, with a third cluster located off the fault zone some 20 km southwest of the gap (see Fletcher and others, 1987; Sanders and Kanamori, 1984).
      The Elsinore fault zone is defined not so much by a coincident alignment of epicenters as by the loci of western end points for clusters of epicenters elongate northeastward between the Elsinore and San Jacinto fault zones. This pattern is most pronounced along the southeast half of the fault; the northwest half, which defines the northeast scarp of the Elsinore Mountains, is marked by scattered clusters of epicenters. As the Elsinore fault enters the Los Angeles Basin to the north, it splays into the Whittier and Chino faults. Historical seismicity levels are considerably lower along the Elsinore fault than either the San Jacinto fault zone or the Imperial fault/Brawley seismic zone. The largest historical earthquake on the Elsinore fault was an M=6 event in 1910 in the central section. The Whittier Narrows earthquake (M_L =5.9) of 1987, which caused over $300 million in damage, was located at the north end of the Elsinore-Whittier fault. Because its mechanism was thrust faulting on an east-west-striking plane with a shallow dip, however, it does not appear to be simply related to the Elsinore system.
      Seismicity in the relatively quiescent southwestern corner of California between the Elsinore fault and the coast shows up in figure 5.10A as small, sparsely scattered clusters of epicenters. Activity picks up again, however, in the vicinity of the major northwest-striking faults along the coast (the Rose Canyon fault through San Diego and the Newport- Inglewood and Palos Verdes faults along the western margin of the Los Angeles Basin). Except for weak alignments along the Newport-Inglewood fault, which ruptured with an M=6.3 earthquake in 1933 (Richter, 1958), the seismicity patterns associated with these faults show little tendency to align along mapped fault traces.
      Moving northwestward along the San Jacinto fault zone", the base of the seismogenic crust deepens systematically to a maximum of 20 km beneath the stretch adjacent to San Jacinto Mountain (which at 3,293 m, is the second highest point in southern California) midway along the fault zone (cross sec. P-P', fig. 5.10B). The base of the seismogenic crust maintains this 20-km depth farther northwestward along the fault zone to its junction with the Banning fault just south of San Bernardino (fig. 5.10A), beyond which it begins to shallow again. Note, in particular, that earthquakes tend to be concentrated between 10- and 20-km depth beneath the San Jacinto fault zone, leaving the upper 10 km of the crust relatively quiescent along the middle stretch of the fault zone. The dense knot of hypocenters in the upper 5 km of the crust midway along cross section P-P' corresponds to the cluster of epicenters 15 km southwest of the fault zone near the Anza gap (fig. 5.10A). The Anza gap itself shows up between Δ=120 and 140 km in cross section P-P' as a quiescent zone below and southeast of the shallow cluster of hypocenters (Fletcher and others, 1987; Sanders, 1987). The distribution of hypocenters beneath the Elsinore fault zone (cross sec. R-R', fig. 5.10B) is in many ways similar to that beneath the San Jacinto fault zone. Maximum focal depths increase northwestward from 12-15 km at the southeast end of the fault near the United States-Mexican border to about 20 km midway along the fault zone (generally coincident with the highest topography in this section of the Peninsular Ranges) and then gradually decrease farther northwestward toward the Los Angeles Basin. Maximum focal depths show evidence of increasing again at the northwest end of the fault as it approaches the Transverse Ranges and branches into the Whittier and Chino faults. The hypocenters along the south half of the Elsinore fault also tend to concentrate in the lower 10 km of the seismogenic crust, although this pattern is not as well defined in the diffuse seismicity of the Elsinore fault zone as in the dense clustering along the San Jacinto fault zone.

      Focal mechanisms of selected earthquakes recorded in California from 1933 through 1988 are shown in figure 5.11A, and the corresponding source parameters are listed in tables 5.2 and 5.3. Primary considerations in the selection of these events were (1) size -larger events were chosen where available, because they represent large-scale processes along major boundaries; (2) date of occurrence-the quality of data for focal-mechanism determinations improved significantly during the mid-1970's; and (3) location-some larger events were omitted because they were redundant in terms of mechanism and location, and some smaller events were included because they occurred in regions of significant seismicity where no larger events were available. Most focal mechanisms were determined from first arrivals at stations in the northern and southern California seismic networks. The evolving capability of these networks for such studies is reflected in the number of stations in the networks, summarized in table 5.1. Fault-plane solutions for the few large earthquakes on the list before the mid-1970's were supported by observations from stations outside the California networks.

Table 5.2 - Locations and focal-mechanism parameters for selected earthquakes in California

Table 5.3 - Locations and focal-mechanism parameters for earthquakes in the Coalinga-Kettleman Hills region

      Since the mid-1970's, focal mechanisms have been determined for only a fraction of the events for which adequate local first-motion data were available. Therefore, in addition to the three considerations listed above, there was a fourth, the interests of the investigators who analyzed the data. These interests included topical studies of large earthquakes and aftershock sequences, analyses of regional traveltimes on the basis of M≥4 earthquakes, and a special study of the focal mechanisms of earthquakes on or near the San Andreas fault in southern California (Jones, 1988).
      Focal mechanisms discussed in the first two subsections below are for earthquakes in the contiguous Coast Ranges-Transverse/Peninsular Ranges-Mojave Desert region associated with the principal seismic expression of the San Andreas fault system, where the seismic networks are best developed. Outside that region, except for the Cape Mendocino area and the vicinity of Long Valley caldera, the few well-determined focal mechanisms that are available provide only limited information on tectonic processes.

      Most moderate and large (M≥3) earthquakes along the San Andreas fault and its major branches produce nearly pure right-lateral displacements along near-vertical planes that closely follow the surface traces of the respective fault segments. This relatively simple kinematic pattern holds for the great earthquakes that rupture "locked" sections of the fault every few hundred years (Sieh, 1981), as well as for nearly all the moderate earthquakes that rupture limited patches along persistently active segments of the fault system (Ellsworth and others, 1982; Jones, 1988). Displacements associated with these earthquakes dominate the kinematic pattern along the transform boundary in California. DeMets and others (1987) and Minster and Jordan (1987), for example, argued that the cumulative displacement from earthquakes along the faults in the San Andreas system, together with the contribution from aseismic slip along its creeping segments, accounts for 60 to 70 percent of the total displacement between the Pacific and North American plates.
      The fault-parallel strike-slip displacements typical of San Andreas earthquakes are illustrated in figure 5.11A by focal mechanisms along the San Andreas fault and its major branches from the United States-Mexican border to north of Clear Lake. In central California, such mechanisms mark the San Andreas fault itself from San Francisco to Cholame (events 26, 36, 38, 45, 46), the Calaveras-Greenville fault (events 23, 28-34) and the Hayward fault (event 27). Farther north, such mechanisms occur along the Green Valley-Bartlett Springs fault (event 15) and the Rodgers Creek-Healdsburg-Maacama faults (events 16, 17, 19, 20). In southern California, such mechanisms mark the San Jacinto fault (events 78, 82-85) and the Imperial fault (event 89). Along the coast west of the San Andreas fault, similar focal mechanisms occur along the San Gregorio-Palo Colorado fault (events 39, 40) in northern California and along the Newport-Inglewood fault zone (events 62, 71), the Rose Canyon fault (event 73), and the San Clemente fault (event 70) in southern California.
      Exceptions to this simple pattern for moderate (M≥4) events along the San Andreas fault and its major branches appear to be limited to regions of unusual complexity, such as the major bends in the San Andreas near Cajon Pass (event 69) and San Gorgonio Pass (event 80). Jones and others (1986) attributed the July 8, 1986, earthquake (event 80) to right-lateral slip on the Banning segment of the San Andreas fault where it dips 45° N. beneath the San Bernardino Mountains. The October 17, 1989, M=7.1 Lorna Prieta earthquake involved nearly equal amounts of right and reverse slip along a section of the San Andreas fault that takes a slight westerly bend through the Santa Cruz Mountains and dips 70° SW. (see chap. 6). Smaller (M<4) events near, but probably not on, the fault show a great variety of focal mechanisms that reflect varying conditions along the fault; these mechanisms range from reverse or reverse-oblique slip on easterly-striking planes (events 37, 67), through right-lateral strike slip on planes parallel to the San Andreas fault (events 65, 66, 68, 87), to normal or normal-oblique slip on northerly-striking planes (events 77, 81).
      Moderate earthquakes with strike-slip focal-mechanisms that are not located on major faults of the San Andreas system but yet are broadly associated with it commonly have right-slip planes, with strikes ranging from northwestward (event 42) to north-southward (events 35, 44, 74, 75, 76, 86). In most cases, these right-slip planes agree in strike with local mapped faults or with alignments of epicenters that strongly suggest active faults (events 75, 76).

      One of the more important results to emerge from high-resolution focal-mechanism studies in recent years is that earthquakes occurring even a short distance off faults of the San Andreas system can involve displacements that diverge sharply from local San Andreas strike-slip displacements. This pattern is particularly pronounced in the strong component of reverse slip at large angles (more than 60° ) to the local strike of the San Andreas fault on both sides of the San Andreas fault system in both the Transverse and Coast Ranges.
      North-south convergence within the Transverse Ranges is dominated by reverse slip on easterly-striking planes. The M = 7.7 Kern County earthquake of 1952 (event 64), which occurred on the south-dipping White Wolf fault along the north flank of the Transverse Ranges about 25 km north of the junction of the San Andreas and Garlock faults, and the M=6.6 San Fernando earthquake of 1971 (event 71), which ruptured a 20-km-Iong stretch of the northeast-dipping San Gabriel-San Fernando thrust faults (Whitcomb, 1971; Heaton, 1982), are two striking examples of this deformation. So, also, is the alignment of M=5-6 reverse-slip earthquakes (events 59, 60, 63) along the southern margin of the Transverse Ranges. The reverse slip on east-west-striking planes associated with these earthquakes suggests that the north-dipping Santa Monica-Cucamonga fault serves as an important convergent boundary between the Peninsular and Transverse Ranges.
      Figure 5.11A also shows that the east-west-trending zone of convergence associated with these earthquakes curves northward near Santa Monica Bay and continues northwestward along the coast at least as far as Point Sal and probably as far as San Simeon. Focal mechanisms of earthquakes along this zone from Point Sal to Whittier (events 55-60, 63) are predominantly reverse slip, with slip directions nearly normal to the local trend of the zone. The focal mechanism of event 43 near San Simeon, which indicates right-oblique reverse slip on a northeast-dipping plane parallel to the coast, is intermediate between those of event 40 at Point Sur and event 55 at Point Sal.
      Reverse-slip focal mechanisms for offshore events 18 and 41 in central California and for event 72 in southern California suggest that the offshore crust is undergoing compression normal to the coastline throughout the length of the San Andreas fault system. The Coalinga-Kettleman Hills earthquake sequence of 1982-85 (events 47-50, fig. 5.11A) emphasizes the important role of crustal convergence along the southern Coast Ranges-Great Valley boundary in central California. The principal events in this sequence (event 48, Coalinga, and event 49, Kettleman Hills, fig. 5.11A) involved reverse slip on subparallel planes at depths of 10 to 12 km that dip gently (approx 20°) southwest. Much of the aftershock activity, however, occurred at shallower depths and involved high-angle reverse slip on planes dipping steeply (45°-70°) northeast (events f, g, i, o, q, fig. 5.11B). Displacements associated with these earthquakes, which are nearly perpendicular to the San Andreas fault, represent a convergent process in which Franciscan melange on the west is being wedged between crystalline basement and the overlying Great Valley sedimentary sequence on the east (Wentworth and others, 1983; Eaton, 1990).
      The boundary between the Coast Ranges and Great Valley is marked by reverse-slip earthquakes throughout much of its length: event 54 southeast of the Kettleman Hills, the Coalinga-Kettleman Hills sequence, event 21 near Winters east of Lake Berryessa, and event 14 west of Oroville. The similarity in focal mechanism of event 21 near Winters to the Coalinga and Kettleman Hills main shocks suggests that the convergent process acting in the southern Coast Ranges is common to the entire eastern margin of the Coast Ranges. Indeed, the strong earthquakes that shook the Winters-Vacaville-Dixon area in 1892, just south of event 21, resemble the Coalinga-Kettleman Hills sequence in both setting and intensity distribution. Focal mechanisms of smaller earthquakes along the Coast Ranges-Great Valley boundary in central California studied by Wong and others (1988) also suggest convergence across that boundary.
      Convergence normal to the strike of the San Andreas fault is not limited to the coast and the Coast Ranges-Great Valley boundary described above. In a detailed examination of the focal mechanisms of aftershocks of the 1984 Morgan Hill earthquake, Oppenheimer and others (1988) concluded that the direction of maximum compression immediately adjacent to the Calaveras branch of the San Andreas fault is at an angle of about 80° to the N. 10° W. strike of the fault. Along the entire stretch of the San Andreas fault from Parldield to the Salton Sea, Jones (1988) found a constant angle of 65° between the strike of the fault and the maximum-principal-stress direction for earthquakes occurring off the fault.
      This evidence from earthquake focal mechanisms and other stress indicators (such as borehole breakouts and fold axes) that the maximum principal compressive stress may be oriented at a high angle to the local strike of the San Andreas fault seems to contradict long-accepted ideas for brittle failure in the crust based on laboratory experiments in rock mechanics. Zoback and others (1987) and Oppenheimer and others (1988) suggested that these relations can be explained by an exceptionally low average shear strength for the San Andreas fault system. As pointed out by Lachenbruch and McGarr in chapter 10, however, the strength and state of stress along the San Andreas fault are still matters for discussion.

      The three moderate-earthquake focal mechanisms for the Sierra Nevada and its western foothills shown in figure 5.11A all indicate normal faulting on northerly-striking planes and suggest pervasive east-west extension throughout the Sierra Nevada. Event 52 is in a dense north-south-trending band of epicenters about 15 km east of the Kern River canyon, and event 53 is in a north-south-trending band of earthquakes about 10 km west of the Sierra frontal fault near Walker Pass. These relations suggest that the east-west spreading and associated normal faulting on northerly-striking faults of the Great Basin are encroaching into the southeast corner of the Sierra Nevada block (Jones and Dollar, 1986).
      Event 13 is the main shock (M=5.7) of an earthquake sequence on a north-south-striking, west-dipping normal fault near the Oroville Dam that occurred in 1975. The uplift of the Sierra Nevada relative to the Great Valley to the west indicated by the focal mechanism of this event is also visible in the Chico monocline, which marks the Sierra Nevada-Great Valley boundary northwest of Oroville.

      The Sierra Nevada-Great Basin boundary zone is represented in figure 5.11A by three focal mechanisms. Events 11 and 12 lie northwest of Lake Tahoe along the edge of a minor gap in the band of seismicity along the east edge of the Sierra Nevada. Both events appear to have resulted from left-lateral slip along steeply dipping, northeast-striking faults; both events had aftershock regions that were elongate northeast-southwest. About 250 km southeast, the M=6.4 July 21, 1986, Chalfant Valley earthquake (event 51) resulted from right-lateral strike-slip displacement on a north-northwest-striking surface dipping 60° SW. An M=5.7 foreshock on July 20 resulted from left-lateral strike slip on a northeast-striking, northwest-dipping surface. These two conjugate slip surfaces merge at their north ends (Cockerham and Corbett, 1987; Smith and Priestly, 1988).
      The zone of intense seismicity in the vicinity of Long Valley caldera and the Sierra Nevada block to the south produced 11 M≥ 5.5 earthquakes from 1978 through 1984 (Savage and Cockerham, 1987), as well as many thousands of smaller events and numerous earthquake swarms. Most of the larger events occurred in the Sierra Nevada block south of Long Valley caldera, involving left-lateral slip along near-vertical, north-south- to north-northeast-striking faults. One of four M≈6 events that occurred on May 25-27, 1980, however, was located within the south moat of the caldera along the west-northwest-striking fault zone that produced most of the earthquake swarms (see Hill and others, 1985a, b).

      The 1980 Eureka M=7.2 earthquake occurred along a fault break that extended from the continental slope 40 km west of the coastline at lat 41° N. for a distance of 140 km southwestward to the MFZ, virtually cutting off the southeast corner of the Gorda plate. Focal mechanisms of the main shock and largest aftershock (events 1, 3, fig. 5.11A) both indicate left-lateral strike-slip displacement along a vertical fault that coincides with the line of aftershocks. Some early aftershocks, including event 4 and other moderate events farther east along the MFZ, have focal mechanisms that indicate right-lateral slip along the MFZ. Although the main shock occurred beneath the Continental Shelf, it seems clear that the 1980 earthquake primarily involved the Gorda plate because the fault broke well beyond the base of the continental slope and the edge of the North American plate. Moreover, left-lateral slip along the 1980 break stimulated right-lateral slip along the adjacent part of the MFZ. Ongoing right-lateral displacement along the MFZ is also indicated by event 5 (Dec. 1983).
      Two moderate earthquakes near Cape Mendocino in 1981 and 1987 (events 6 and 7, respectively) had focal mechanisms similar to that of the 1980 Eureka earthquake, indicating left-lateral strike-slip displacement on steeply dipping, northeast-striking planes. Aftershocks of the 1987 M=5.8 event outlined a narrow, steeply dipping, northeast-trending, 20-km-Iong zone between about 15- and 25-km depth that extended southwestward from the shoreline just north of Cape Mendocino to the MFZ. This aftershock zone appears to cut off the southeasternmost corner of the Gorda plate just north of the abrupt eastward termination of intense seismicity along the MFZ, at a point that might be taken as the Mendocino triple junction from the viewpoint of seismicity.
      Relative horizontal extension at seismogenic depths is suggested by events 2 and 8. Event 2 (Nov. 10, 1980; 7 km deep) was the largest in a detached cluster of shallow aftershocks 20 km east of the 1980 main shock, and event 8 (Apr. 9, 1987; 26 km deep) occurred about 100 km east of the 1980 main shock in the zone of seismicity associated with the subducting Gorda plate.

      The Pacific plate moved northwestward with respect to the North American plate by 300 to 400 mm during the 7-yr interval 1980-86. Earthquakes occurring along the San Andreas fault system during the same interval, however, accommodated only a small fraction of this relative plate motion. Only four earthquakes of M>5 occurred along branches of the San Andreas fault system during 1980-86: the pair of M=5.9-5.3 Livermore earthquakes (events 29, 30, fig. 5. 10A) on the Greenville fault (Jan. 24-27, 1980), the M=6.2 Morgan Hill earthquake (event 33) on the Calaveras fault (Apr. 24, 1984), and the M=5.6 North Palm Springs earthquake (event 80) on the Banning segment of the San Andreas fault (July 8, 1986). Each of these moderate San Andreas earthquakes ruptured fault segments limited to 20 to 30 km in length, with average displacements over the respective rupture surfaces of 100 to 200 mm (see Hartzell and Heaton, 1986). As is typical of earthquakes along the San Andreas fault system, each of these events involved nearly pure right-lateral strike-slip displacement coincident with the local strike of the fault. As is also typical of San Andreas earthquakes, slip on the first three events occurred on near-vertical fault planes with a northwestward to north-northwestward strike. The North Palm Springs earthquake, which ruptured a section of the east-west-striking Banning fault in the structurally complex San Gorgonio bend in the fault system at the southern margin of the Transverse Ranges, represents an important deviation from typical San Andreas earthquakes. Although its displacement was dominantly right-lateral strike slip, it occurred along a plane that dips 45° N. (Jones and others, 1986) and included a small but significant component of reverse slip (Mendoza and Hartzell, 1988). With the arguable exception of the North Palm Springs earthquake (arguable because of the complex section of the fault system in which it occurred), however, none of these M>5 earthquakes ruptured the main trace of the San Andreas fault. Indeed, the two most recent M>5 earthquakes to clearly do so were the M=6 Parkfield earthquake of 1966 (Bakun and McEvilly, 1984) and the M=7.1 Lorna Prieta earthquake of 1989 (see chap. 6).
      Thus, aside from the displacement accommodated by steady aseismic slip at a rate of 32 to 37 mm/yr along the creeping section of the fault in central California, most relative plate motion across the San Andreas transform boundary during this 7-yr interval accumulated as elastic shear strain. Accordingly, the earthquakes plotted in figures 5.3 through 5.9 are symptomatic of accumulating strain along the San Andreas fault system rather than of effective strain release. The latter requires rupture with a major earthquake along one of the locked stretches of the San Andreas fault.

      What changes in spatial-temporal patterns of earthquake occurrence might we expect to see as the next great earthquake on the San Andreas fault approaches? Both historical and instrumental seismicity records indicate that the spatial distribution of earthquakes in California changes only slowly over periods of decades to centuries, although the intensity of activity within this distribution fluctuates year to year (Ellsworth and others, 1981; Hill and others, in press; Hutton and others, in press). Temporal fluctuations in activity during the interval 1980-86, for example, were dominated by a short-lived aftershock sequence following the 1980 Eureka M=7.2 earthquake and by the long-lived aftershock sequence following the 1983 Coalinga M=6.7 earthquake. The overall spatial distribution of earthquakes in California, however, remained nearly stationary throughout this 7-yr interval. Furthermore, the spatial pattern defined by 1980-86 seismicity is much the same as that outlined by the record of M≥5 earthquakes that extends back nearly 200 yr (see chap. 6).
      Variations in the historical rate of moderate to large (M>5) earthquakes in central California before and after the 1906 San Francisco earthquake appear to mimic those described by Fedotov (1965) and Mogi (1968) for the earthquake cycle associated with great, subduction-zone earthquakes in Japan, Kamchatka, and the Kurile Islands (see chap. 6; Ellsworth and others, 1981). The history of instrumentally recorded M <5 earthquakes in California is too short, however, to indicate whether we might expect to see distinctive changes in the seismicity pattern a short time (months to years) before the next great earthquake on the San Andreas fault. We have yet to see, for example, whether the quiescent (locked) segments of the San Andreas fault remain aseismic except for the rupture of a great earthquake, or whether these segments become active with small to moderate earthquakes as foreshock activity to great earthquakes.

      The two largest earthquakes in California during the interval 1980-86 occurred off the faults of the San Andreas system, and their occurrence emphasizes the importance of deformation within the plate margins along the San Andreas transform boundary. The M=7.2 Eureka event (Nov. 8, 1980), for example, involved deformation internal to the Gorda plate; and the M=6. 7 Coalinga event (May 2, 1983) involved crustal shortening with reverse slip perpendicular to the San Andreas fault. These two earthquakes and the many smaller, "off fault" events (fig. 5.4A) reflect local deviations from the simple rigid-plate approximation of plate tectonics.

      As the small, youthful Gorda plate is subducted obliquely northeastward beneath the North American plate, it is being subjected to north-south compression in response to a component of convergence between the larger, older Juan de Fuca plate to the north and the Pacific plate to the south (Jachens and Griscom, 1983; Wilson, 1986). Distorted marine magnetic anomalies within the Gorda plate indicate that it has undergone progressive internal deformation over the past 5 Ma in response to this compression (Silver, 1971), and current seismicity within the plate (fig. 5.4) indicates that this deformation continues to the present. The 1980 Eureka M=7.2 earthquake emphasizes that part of this deformation occurs with left-lateral slip on northeast-striking faults within the plate. The seismicity map and cross sections (fig. 5.4) demonstrate that deformation associated with the Gorda plate terminates abruptly against the Pacific plate in a steeply north-dipping zone of interaction along the MFZ, which can be followed on shore beneath the North American plate as a gently east-dipping, subhorizontal zone of widely scattered small to moderate earthquakes. Thus, convergence between the Gorda and Pacific plates across the MFZ apparently occurs by crushing and thickening of the southern margin of the Gorda plate as it is jammed against the anvil-like mass formed by the thicker and colder Pacific plate. Diminished east-west stress in the Gorda plate resulting from the subducting limb of the plate farther east serves to increase the difference between the maximum (north-south) and minimum (east-west) compressive stresses within the plate, leading to left-lateral strike-slip displacements along northeast-striking faults, as in the M=7.2 Eureka earthquake. This process accommodates the convergent component of Gorda-Pacific plate motion along the east end of the MFZ at the expense of fragmentation and eastward expansion of the Gorda plate north of the MFZ.

      Much of the seismicity adjacent to the San Andreas fault system is attributable to differences between the long-term slip rate and direction (slip vector) along the San Andreas fault system and that predicted for relative motion between the Pacific and North American plates along the San Andreas transform boundary on the basis of global models of plate motion. Minster and Jordan (1978, 1987) predicted that the direction of dextral slip between the Pacific and North American plates along the San Andreas transform boundary in central California is N. 35° W. The main trace of the San Andreas system, however, strikes N. 41° W. through central and northern California and N. 65°-70° W. through the Transverse Ranges in southern California. DeMets and others (1987) concluded that the marine magnetic anomalies at the mouth of the Gulf of California constrain the slip rate to an average of 49 mm/yr over the past 3 to 4 Ma. Both long-term geologic offset data and geodetic data measured over the past several decades, however, indicate that the average slip rate along the San Andreas fault system is only about 35 mm/yr. The contribution to deformation of the western margin of the North American plate from spreading across the Basin and Range province is about 10 mm/yr in a N. 56° W. direction (Minster and Jordan, 1987). Ellsworth (see chap. 6) suggests that most of the San Andreas discrepancy can be explained if the component of dextral slip associated with historical Basin and Range earthquakes reflects a long-term trend superimposed on the N. 56° W. spreading direction. If so, then the residual component of Basin and Range extension perpendicular to the San Andreas fault system is approximately balanced by convergence across the Coast Ranges and continental margin.

      Focal mechanisms of earthquakes occurring off the San Andreas fault system suggest that the component of the San Andreas discrepancy normal to the fault system may, indeed, be accommodated by distributed brittle deformation on either side of the fault system. These mechanisms range from dextral strike slip on planes subparallel to the San Andreas fault, through oblique-reverse slip, to nearly pure reverse slip with a slip direction perpendicular to the San Andreas fault.
      The Coalinga-North Kettleman Hills earthquake sequence provides clear evidence for crustal convergence perpendicular to the San Andreas fault system in the Coast Ranges. The several smaller events with similar mechanisms to the north along both the eastern and western (coastal) margins of the Coast Ranges (fig. 5.11) suggest that the convergence responsible for the Coalinga earthquake may be active the length of the Coast Ranges (Wong and others, 1988; Eaton and Rymer, 1990). The subparallelism of fold axes within the Coast Ranges with the San Andreas fault indicates that fault-normal convergence has been important for the past 3 Ma in central California (fig. 5.12; Page and Engebretson, 1984). N amson and Davis (1988) proposed that the entire system of Coast Range folds may be genetically related to Coalinga-like earthquake sequences and low-angle (blind) thrust faults that are rooted in a decollement near the base of the seismogenic crust. The reverse focal mechanisms for earthquakes associated with offshore faults along the western margin of the Coast Ranges suggest that, here, convergence involves westward thrusting of the Coast Ranges over oceanic crust of the Pacific plate.

Figure 5.12 - Seismicity from 1980 to 1986 superimposed on digital shaded-relief image of central California, showing faults (blue) and fold axes (red). Size of symbol for epicenters (yellow) increases with magnitude from 1 to 6. Shaded relief by Raymond Batson, U.S. Geological Survey (illumination from north at 30°); overlays from Ross Stein (unpub. data, 1989). SAF, San Andreas fault.

      The pronounced discrepancy in the strike of the San Andreas fault through the Transverse Ranges with respect to the Pacific-North American plate slip direction provides an obvious source of local crustal convergence (Hill and Dibblee, 1953; Atwater, 1970), and the associated structural complexities serve to distribute brittle deformation (seismicity) much more widely about the San Andreas fault system in southern California than about the relatively straight sections of the fault system in central and northern California. The largest earthquake in California since the great 1906 San Francisco earthquake occurred near the northern margin of this convergent regime; this M=7.7 Kern County earthquake ruptured some 35 km of the southeast-dipping White Wolf fault with left-oblique reverse slip on July 21, 1952.
      The focal mechanisms of larger Transverse Range earthquakes, together with the mapped attitudes of major faults with Holocene offsets, show that much of this convergence occurs with slip on north-dipping thrust faults within and along the southern margin of the central Transverse Ranges (fig. 5.11A). For earthquakes in the western Transverse Ranges, the direction of reverse slip is more southwestward, consistent with thrusting of the western Transverse Ranges over the Pacific plate similar to that in the Coast Ranges to the north.

      The fault-normal convergence that dominates deformation adjacent to the San Andreas fault system through both the Coast Ranges and Transverse Ranges gives way rather abruptly to the extensional regime of the Salton Trough near the southern margin of the intensely active San Gorgonio bend in the fault. Focal mechanisms of earthquakes occurring on secondary structures adjacent to the seismically quiescent Indio segment of the San Andreas fault, for example, show a mix of strike- and dip-slip mechanisms. As is the case farther north, however, P-axes for these earthquakes tend to be oriented at a high angle (60°-65°) to the fault, suggesting that the Indio segment of the fault may also be relatively weak (Jones, 1988).
      One particularly noteworthy aspect of seismicity south of the Transverse Ranges is the, tendency for earthquakes to occur along conjugate strike-slip structures. Recall that the Sierra Nevada-Great Basin boundary zone also shows this tendency and that both regions are subject to extensional deformation, earthquake swarms, and late Quaternary volcanism. Earthquake sequences within the southern section of the San Andreas fault system commonly produce epicenter lineations that intersect at nearly a 90° angle with the northwest-striking right-slip plane and the northeast-striking left-slip plane. Earthquake-swarm sequences in the Brawley seismic zone, for example, typically occur along northeast-striking lineations normal to the trace of the adjacent Imperial fault (Johnson, 1979), and the M=5.7 Westmorland earthquake of 1981 involved left-lateral slip along several subparallel, northeast-striking planes (Johnson and Hutton, 1982). The diffuse lineations of epicenters spanning the area of the Peninsular Ranges between the San Jacinto and Elsinore faults also tend to be orthogonal to these two branches of the San Andreas fault system (fig. 5.10A). An impressive recent example of this orthogonal conjugate pattern is the M=6.2 and 6.6 Superstition Hills earthquakes of November 24, 1987 (Magistrale and others, 1988).
      The kinematics of these conjugate structures remains a matter of conjecture. Dextral slip along through going faults of the San Andreas system must certainly dominate deformation, and the shorter, northeast-striking structures must play only a secondary role. Nicholson and others (1986) proposed that the northeast-striking lineations represent the boundaries between blocks rotating clockwise much like roller bearings, between subparallel pairs of dextral strike-slip faults. Hill (1977) and Weaver and Hill (1978/79) suggested that within local spreading centers, such as the Brawley seismic zone, conjugate strike-slip structures form miniature triple junctions with a dike or normal fault that subtends the acute angle between the conjugate strike-slip faults.

      Maximum focal depths of earthquakes beneath the San Andreas transform boundary range from less than 5 km beneath the Geysers geothermal field in the northern Coast Ranges to more than 20 km beneath the Transverse Ranges, the eastern margin of the Coast Ranges, and the San Jacinto and Elsinore faults in southernmost California. Beneath relatively straight segments of the San Andreas fault system through central California, maximum focal depths range from 12 to 15 km (figs. 5.7, 5.8). Sibson (1983) pointed out that these variations in maximum focal depth along the San Andreas fault system are inversely correlated with surficial heat flow, and he argued that the maximum depth of earthquakes coincides with the temperature-dependent transition from brittle failure in the upper crust to aseismic, quasi-plastic flow in the lower crust and upper mantle. For quartz-bearing rocks typical of the upper crust and deformation rates typical of the San Andreas fault system (1x10^-14 to 1x10^-13 s^-l), this brittle/ductile transition occurs at about 300 °C (Sibson, 1983). By this interpretation, the thin seismogenic crust beneath both the Geysers and Brawley geothermal fields in northern and southern California, respectively, reflects elevated temperatures in the shallow crust, whereas the relatively thick seismogenic crust beneath the Transverse Ranges and the eastern margin of the Coast Ranges reflects depressed temperatures in the midcrust associated with crustal convergence. Although temperature may dominantly influence the thickness of the seismogenic crust, local variations in rock composition (particularly the presence or absence of modal quartz and structural water) and in strain rate can also be important. These variations, for example, may help explain isolated clusters of deep earthquakes, such as the 20- to 24-km-deep events north of San Pablo Bay in central California (see cross secs. F-F', G-G', fig. 5.8B).
      In any case, the thickness of the seismogenic crust beneath the San Andreas transform boundary seems to be much more strongly related to temperatures in the crust than to the structural thickness of crust defined by the depth to the Moho (see chap. 8). This relation is strikingly illustrated by the twofold increase in thickness of the seismogenic crust beneath the rootless Transverse Ranges.

      A theme common to models of crustal convergence along the San Andreas fault system involves low-angle reverse slip on decollement surfaces near the base of the seismogenic crust (Wentworth and others, 1983; Webb and Kanamori, 1985; Namson and Davis, 1988; Eaton and Rymer, 1990). A natural extension of this theme leads to a view of the seismogenic crust as a conglomeration of relatively rigid blocks interacting by frictional slip along weak preexisting faults (block boundaries) in response to regional stresses transmitted through both the brittle crust and quasi-plastic deformation in the underlying lithosphere (Hill, 1982). However, the nature of a decollement surface at the base of the brittle crust and the relation of the seismogenic San Andreas fault system to the aseismic transform boundary in the underlying lithosphere remain speculative. It is not yet clear, for example, whether the San Andreas fault continues below the seismogenic crust as a narrow, near-vertical boundary (possibly offset a substantial distance from the seismogenic fault by slip on the horizontal decollement surface) that slips by quasi-plastic, mylonitic deformation or whether it broadens rapidly with depth into a wide shear zone spanning, say, the entire width of the Coast Ranges (see chap. 7; Sibson, 1983).

      The spatial-temporal pattern of earthquake occurrence within the seismogenic crust along the San Andreas fault system is the brittle manifestation of distributed deformation of the lithosphere between the Pacific and North American plates along the San Andreas transform boundary. As we develop a more complete model of the long-term behavior of the seismogenic crust, including relations between great, plate-boundary earthquakes that periodically rupture the principal strand of the San Andreas fault system and the persistent background of small to moderate earthquakes on adjacent structures, our image of the deeper deformation will improve. Together will come a more complete understanding of the processes controlling deformation along the transform boundary and of the earthquake cycle.

REFERENCES CITED