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Bulletin of the Seismological Society of America, 91, 6, pp. 1471–1497, December 2001

Source Characterization and Ground-Motion Modeling of the 1892 Vacaville–Winters Earthquake Sequence, California by Daniel R. H. O’Connell, Jeffrey R. Unruh, and Lisa V. Block

Abstract A sequence of several earthquakes in April 1892 produced significant damage in the towns of Winters, Dixon, Allendale, and Vacaville along the boundary between the southwestern Sacramento Valley and northern Coast Ranges of California. The largest event occurred on 19 April 1892 with a maximum Modified Mercalli intensity of IX and was assigned a moment magnitude (M) of 6.5 based on felt area. These earthquakes occurred within a zone of active crustal shortening accommodated by postulated blind thrust faults. Seismotectonic and structural analyses are used to evaluate the depth, geometry, and segmentation of thrust faults that were the probable sources of the 1892 earthquake sequence. Synthetic ground-motion modeling demonstrates that rupture of a 17-km-long segment of the thrust fault system can produce the magnitude and distribution of intensities documented from anecdotal accounts of the 19 April 1892 earthquake, including probable directivity effects east of the range front. Integrated structural and seismotectonic analyses also are used to interpret the role of inferred geometric segment boundaries in arresting the 19 April 1892 earthquake rupture, and the subsequent occurrence of the 21 April 1892 aftershock. Introduction The boundary between the coastal mountains and western Central Valley of California has been recognized as a zone of active crustal shortening and a potential source of earthquakes from blind thrust faults (Wong and Ely, 1983; Wentworth et al., 1984; Wong et al., 1988; Wentworth and Zoback, 1989; Unruh and Moores, 1992; Wakabayashi and Smith, 1994). Wong and Ely (1983) referred to this region as the Coast Ranges–Sierran Block Boundary Zone (CRSBZ) and suggested that the contractional deformation results from westward motion and impingement of the crystalline basement of the Sierra Nevada, which underlies the Central Valley, against the accreted terranes of the California Coast Ranges. Subsequent work established that active crustal shortening along the CRSBZ is primarily accommodated by blind, west-dipping thrust faults (Wentworth et al., 1984; Namson and Davis, 1988; Unruh and Moores, 1992). Workers have attempted to use the morphology and continuity of late Cenozoic contractional structures mapped at the surface to define potential rupture segments for the blind thrust fault system (Wakabayashi and Smith, 1994; WGNCEP, 1996). Wong and Ely (1983) analyzed anecdotal accounts and proposed that seven moderate earthquakes (ML ⱖ4.5) may have occurred within this zone during the 150 yr prior to 1982. The best-documented sequence of thrust fault earthquakes in the CRSBZ occurred during the 1980s along the western margin of the southern portion of the Central Valley

and included the 1982 moment magnitude (M) 5.5 New Idria earthquake (Scofield et al., 1985), the 1985 M 6.1 Kettleman Hills earthquake (Ekstro¨m et al., 1992), and the 1983 M 6.5 Coalinga earthquake (Stein and Ekstro¨m, 1992). The 1892 Vacaville–Winters earthquake sequence occurred north of Coalinga along the southwestern edge of the Sacramento Valley (which comprises the northern portion of the Central Valley) and has been identified as a potential blind thrust fault earthquake of the CRSBZ (Wong, 1984; Eaton, 1986; Unruh and Moores, 1992). Based on anecdotal accounts and the distribution of felt effects, Toppozada et al. (1981) concluded that the mainshock epicenter of the 1892 sequence was located well to the east of active strike-slip faults of the San Andreas system and that one aftershock may have occurred approximately 10 km east of the Coast Ranges mountain front (Fig. 1). Bakun (1999) arrived at a similar conclusion using a seismic intensity misfit procedure and estimated an M 6.4 for the 19 and 21 April 1892 earthquakes and an M 5.5 for the 30 April 1892 aftershock. Although there are gross tectonic and geomorphic similarities between the Coalinga and Vacaville–Winters areas consistent with active thrust faulting (Unruh and Moores, 1992), the precise location and geometry of a blind thrust fault source or sources responsible for the April 1892 earthquake sequence has not been identified. In this article we present the results of crustal velocity

1471

1472

D. R. H. O’Connell, J. R. Unruh, and L. V. Block

-122.300

-122.205

-122.109

Longitude -122.014 -121.918

-121.823

-121.727

40

38.76 800

Esparto

20

600

38.58

Davis

EH MD Winters

10

PV

38.49

Allendale

Latitude

North grid distance (km)

38.67

Sept. 8, 1978 M4.2

400

EH

Dixon

0

38.40

Vacaville -10

38.31

-20

38.22 0

10

20 30 East grid distance (km)

40

50

Figure 1. Color-shaded topography with plan views of the Gordon Valley (blue polygon) and Trout Creek (black dashed polygon) thrust faults, the postulated 19 April 1892 M 6.5 epicenter (black star), 1982-1998 microearthquake epicenters (gray circles), the approximate position of the “fissure’ described in Bennett (1987) (yellow line), and nearby cities (white squares). The black thick-dashed contour encompasses the Imm ⱖ VIII-IX region for the 19 April mainshock derived from Toppozada et al. (1981) and Bennett (1987). The 21 April M ⬃6.2 aftershock Imm VIII-IX intensity contour is contained within the 19 April mainshock Imm VIII-IX contour. The black thin-dashed contour shows the western extent of the Imm ⳱ V-VI intensity contour for the 30 April M ⬃5.5 aftershock from Toppozada et al. (1981). The black and white square, labeled MD, shows the location of the Monticello Dam; the white arrow, labeled PV, points toward Pleasants Valley; and EH shows the English Hills, the easternmost low hills. The 8 September 1978 M 4.2 thrust-faulting earthquake (Eaton, 1986; Wong et al., 1988) located on the Trout Creek thrust fault is shown in magenta; the thick line denotes the strike of both nodal planes; and the thin lines indicate dip, with dips of zero having lengths of half the strike-line length and dips of 90⬚ having zero length. The Putah Canyon is the green shaded gorge east of the Monticello Dam that ends where the Putah Creek turns to the southeast along the west side of the Gordon Valley thrust fault-propagation fold axis. The red dot-dashed curve is the axis of the synclinal trough east of the range front. A thin solid black curve traces the base of the Venado Formation showing its change in strike near latitude 38.45⬚ N.

200

0

Elevation (m)

30

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Source Characterization and Ground-Motion Modeling of the Vacaville–Winters Earthquake Sequence

modeling and analysis of seismic reflection data to locate the probable thrust faults responsible for the 1892 earthquake sequence. These data, combined with analysis of the rangefront structure, are used to develop a segmented source model for the blind thrust system. We show through modeling of strong ground motions that rupture of a 17-km-long segment of the thrust system could have produced the magnitude and distribution of Modified Mercalli intensities (Imm) documented by anecdotal accounts of the 19 April 1892 event (Toppozada et al., 1981; Bennett, 1987), including probable directivity effects east of the modern range front. Finally, we integrate structural and seismotectonic analyses to interpret the role of inferred geometric segment boundaries in arresting the 19 April 1892 rupture.

M

17

.4

Upper sheet fold

5

. 55

km

th

km

or

The physiographic boundary between the Sacramento Valley and Coast Ranges approximately coincides with the base of an uplifted and east-tilted section of Cretaceous and early Tertiary marine sedimentary rocks (Figs. 2, 6). The east-dipping strata form a prominent series of north-northwest-trending ridges and valleys that can be traced for tens of kilometers along the strike. The east-dipping strata flatten abruptly along the western valley margin and extend eastward as an approximately horizontal section beneath the valley floor. Based on the stratigraphic thickness of the eastdipping section the minimum structural relief due to uplift and tilting at the range front is about 7-10 km. Stratigraphic and structural relations show that eastward tilting along the range front initially occurred in late Cretaceous-early Tertiary time, and that the cumulative structural relief represents several episodes of deformation during the Tertiary (Unruh et al., 1995). The most recent episode of uplift, tilting, and folding occurred since deposition of the 3.4- to 1.0-Ma Tehama Formation in the western Sacramento Valley (Unruh and Moores, 1992). This deformation has resulted in development of the English Hills (Fig. 2), a range of low hills between the towns of Winters and Vacaville (Fig. 1) that are underlain by the monoclinally folded Tehama Formation and the Miocene Lovejoy basalt. Anecdotal accounts indicate that the English Hills were the locus of the strongest ground-shaking effects during the 1892 earthquake sequence (Toppozada et al., 1981; Bennett, 1987). Fluvial-geomorphic relations provide evidence for late Quaternary uplift and tectonic activity of the English Hills. Fluvial deposition in this region is dominated by Putah Creek, which has deposited a large Pleistocene alluvial fan complex in the western Sacramento Valley east of the range front. This fan complex overlies the Tehama Formation and thus is less than 1.0 Ma in age. Older remnants of this fan complex in the vicinity of Winters have been uplifted and incised by the Putah Creek (Thomasson et al., 1960). Field reconnaissance of this region reveals that the surfaces on the older fan remnants appear to converge with the modern thal-

a

(38.85°N, Trout Creek blind thrust fault-propagation fold axis 122.1°W) Ru m se y Du Hi nn lls ig an Hi lls an tic lin e

N

Quaternary Deformation and Models for Thrust Fault Geometry in the Southwestern Sacramento Valley

i

ta

n ou

c Va

ns

Lower sheet fold Gordon Valley blind thrust (38.35°N, fault-propagation fold axes 121.9°W)

Figure 2. Shaded-relief topography in a 3D perspective showing the folding in the English Hills identified in the seismic reflection data. The maximum elevation plotted is 500 m, and the lowest elevation is 5 m. The perspective produces extreme vertical exaggeration and horizontal distortion that facilitates visualization of the topography in the English Hills. Two fold axes are associated with the Gordon Valley thrust fault with the lower sheet spanning almost the entire fault and an upper sheet in the northern ⬃10 km of the fault. The northern Gordon Valley thrust fault-propagation fold dies as it intersects the Vaca Mountains and is replaced northward by the Trout Creek fault-propagation fold. The two fold axes exhibit a right-stepping, en-echelon geometry consistent with dextral wrench faulting.

weg of the Putah Creek downstream of the town of Winters, suggesting that eastward tilting of the older fan complex has occurred since 1.0 Ma. A series of fluvial terraces inset below the older alluvial fan surfaces in the vicinity of Winters provides additional evidence of late Quaternary tectonic activity. These terraces have been recognized and mapped in soil surveys of Solano County (Bates, 1977) and Yolo County (Andrews et al., 1972), and the degree of soil profile development in the terrace deposits can be used to place broad constraints on their ages. Soils on the lowest terraces above the modern Putah Creek floodplain are Entisols, and soils on higher terraces include both Alfisols and Palexeralfs, which represent moderate to significant degrees of profile development. Comparison of these soils with an existing soil chronosequence in the Dunnigan Hills approximately 25-km north of Winters (Munk, 1993) suggests that the uplifted terraces are middle-

1474

Backthrust

Subsurface location of blind thrust fault

V1

V1

V2

East A'

V3

V5

V6

V4

Plan view

West A

Range - Front monocline


West A

COAST RANGES

V9

p

s

am tr

ru

d

in

Bl

th

East A'

Monticello SACRAMENTO Dam VALLEY V5 6 V V7 V8

late Quaternary in age. Soil maps show that similar terrace sequences are present along smaller drainages south of the Putah Creek that head in the English Hills. Based on these soil-geomorphic relations we interpret that uplift-triggered fluvial incision of the English Hills has occurred in late Quaternary time. Following Wentworth et al. (1984) and Namson and Davis (1988), Unruh and Moores (1992) inferred that Quaternary uplift and eastward tilting in the southwestern Sacramento Valley occurred above an east-tapering, underthrust tectonic wedge. The tectonic wedge model can be described as an east-vergent fault-bend fold (cf. Suppe, 1983) above a ramp-flat transition with one or more east-dipping backthrusts (i.e., roof thrusts) to transfer east-directed displacement back toward the west (Fig. 3). As strata in the hanging wall are progressively pushed up the ramp and onto the flat, they form a range-front monoclinal fold similar to the easttilted panel of marine strata along the Coast Ranges–Sacramento Valley boundary. This model requires that the upper flat or detachment surface beneath the monocline be located at the stratigraphic level of the major east-dipping thrust fault that forms the roof thrust of the tectonic wedge because this is the stratigraphic level at which the crust is being delaminated by propagation of the wedge tip (Fig. 3). The geometry of this model further implies that the vertical surface projection of the top of the thrust ramp must lie slightly west of the range-front monocline and that the surface width of the monocline is a crude measure of the total east-directed slip (i.e., propagation of the wedge tip) on the blind rampflat system beneath the fold (Fig. 3). The primary evidence cited in support of tectonic wedging in the southwestern Sacramento Valley is seismic reflection profiles from the Rumsey Hills–Dunnigan Hills area approximately 50 km north of the epicentral region of the 1892 earthquake. Unruh et al. (1995) recognized a subhorizontal to gently west-dipping reflector imaged by these data at a depth of about 5–7 km. They interpreted this reflector to be the detachment for an underthrust tectonic wedge (Unruh et al., 1995). Previous workers (Kirby, 1943; Redwine, 1972) mapped east-dipping, layer-parallel thrust faults within Cretaceous marine strata in the Rumsey Hills region, and Unruh et al. (1995) interpreted these structures to be backthrusts of the tectonic wedge system. Extending the tectonic wedge model from the Rumsey Hills south to the epicentral region of the 1892 sequence is problematic because well-defined backthrust faults have not been mapped in the English Hills area (e.g., Sims et al., 1973). Such faults may be present but difficult to identify if they are parallel to bedding and do not repeat stratigraphic section. Alternatively, uplift and eastward tilting along the southwestern valley margin could be the result of faultpropagation folding (Fig. 4). In this model the hanging-wall block is folded into a monocline above the propagating tip of a blind thrust fault (Fig. 4). The tip of the fault is tied to a synformal hinge that marks the base of the monocline, and as the tip of the fault propagates updip the synformal hinge


Ba

ck th

ru

Upper flat

st

V1 V2 V3 V4 V5 V6 V7 V8 V9

Figure 3. Schematic plan view (top) and crosssectional view (bottom) of a range-front monocline formed by uplift and propagation of an east-tapering tectonic wedge. The wedge is a fault-bend fold developed above a ramp-flat transition in a blind, westdipping thrust fault. The forelimb of the fault-bend fold delaminates the crust along the stratigraphic contact between units V4 and V5, forming an eastdipping backthrust. Note that the vertical surface projection top of the blind thrust ramp is slightly west of the range-front monocline. The acoustic velocities of the stratigraphic units increase with depth and age with unit V9 exhibiting the highest velocity and unit V1 exhibiting the lowest velocity.

progressively moves toward the foreland region, causing the monocline to widen as the structural relief in the hangingwall block increases (Suppe and Medwedeff, 1990). As strata are flexed and folded by growth of the monocline, layer-parallel shear may accommodate differential flexure and lead to the development of shallowly rooted out-ofsyncline thrust faults (Fig. 4). Both the tectonic wedge and fault-propagation fold models can explain the development of the range-front structure in the epicentral region of the 1892 earthquake sequence. Although both models invoke slip on a west-dipping thrust fault to produce map-scale uplift and eastward tilting, the location of the fault relative to the locus of uplift and folding at the range front differs significantly. Specifically,


V1

V5

V4

V3


East B'

V1

Plan view

V2

West B

Subsurface location of blind thrust fault West B


Monticello Dam SACRAMENTO VALLEY

COAST RANGES

“O

V5 V6 V7 V8

ut

-o

f-s

yn

cli

ne

”

thrust

V9

ult

st fa

hru ind t

East B'

Bl

V1 V2 V3 V4 V5 V6 V7 V8 V9

Figure 4. Schematic plan view (top) and crosssectional view (bottom) of a range-front monocline formed by fault-propagation folding. Layer-parallel shearing during growth of the fold is accommodated by shallowly rooted out-of-syncline thrust faults. Note that the vertical surface projection eastern tip of the blind thrust is approximately near the center of the range-front monocline. The acoustic velocities of the stratigraphic units increase with depth and age, with unit V9 exhibiting the highest velocity and unit V1 exhibiting the lowest velocity.

the fault-propagation fold model (Fig. 4) predicts that the vertical surface projection of the thrust tip is closer to the synformal hinge at the base of the range-front monocline than the tectonic wedge model (Fig. 3). These differences in depth and fault geometry predict strikingly different distributions of strong ground shaking near the range front.

Crustal Structure of the Eastern Coast Ranges Piedmont We performed multidisciplinary analyses of the crustal structure in the 1892 epicentral region to resolve ambiguity about the presence, location, and geometry of blind thrust faults. These analyses, described subsequently, included 3D modeling of the crustal velocity structure and interpretation of depth-migrated seismic reflection profiles.

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Progressive Velocity-Hypocenter Inversion Data and Analytical Approach. A multistep approach was used to locate earthquakes and develop 3D crustal velocity models for the epicentral region of the 1892 earthquake sequence. The U.S. Geological Survey (USGS) Northern California Seismic Network (NCSN) microearthquake waveform data from 1982 to 1998 were combined with digital microearthquake recordings from four temporary stations adjacent to Monticello Dam near Winters to derive a P- and S-wave arrival-time database suitable for velocity-hypocenter inversion. About 10,000 digital seismograms were iteratively processed to obtain 7886 P-wave arrivals and 1099 S-wave arrival times. Seismographic network coverage near the town of Winters was nearly absent until 1987, when the U.S. Bureau of Reclamation contracted to the USGS to add six stations to the NCSN near the Monticello Dam, which is located 12 km west of the town of Winters (Fig. 1). Virtually all the NCSN data were transmitted using analog telemetry, and signal quality varied substantially as a function of time. Aging of the analog telemetry components is clearly evident in the decay of signal quality from 1990 to 1998. Efforts to pick P- and S-wave arrival times were hampered by occasional cross talk between stations, one-sided signal clipping, and low dynamic range coupled with high background noise levels at many of the stations. Best arrival-time picking precision ranges from 0.01 to 0.02 sec (sample rates ranged from 67 to 100 samples/sec). Picking uncertainties for the most impulsive arrivals are about 0.02 sec. Picking uncertainties for emergent arrivals are much larger, ranging from 0.05 to ⬎0.3 sec for most arrival-time picks. A relatively high proportion of emergent arrivals are a result of high background noise and large distances between stations in the relatively sparse seismographic network in the Winters area (Fig. 5). The nonlinear hypocenter-velocity-station correction inverse problem is solved by iteratively inverting the linearized problem until the changes in the root mean square arrival-time residuals are small (ⱕ0.001 sec). To improve the stability of this iterative method, we limited the maximum variation of the model parameters (velocities, hypocenters, and station corrections) during each iteration. The separation-of-parameters technique was used to separate the velocity-station correction inversion from the hypocentral relocation (Pavlis and Booker, 1980; Spencer and Gubbins, 1980). The velocity-hypocenter-station correction inversion is formulated as a constrained least-squares optimization with residual weighting to approximate an exponential (L1) norm inversion (Scales et al., 1988; Vasco et al., 1994), consistent with the distribution of arrival-time residuals. In addition to minimizing the sum of the squared residualweighted arrival-time residuals, the sum of the squared first spatial P-wave velocity derivatives are also minimized. This constraint, or regularization, prevents unreasonable velocity variations at poorly resolved nodes. A conjugate gradient

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Coast Range

Figure 5. Map of the study area showing microearthquake epicenters (gray circles) and seismographic stations (triangles; filled triangles have one or more horizontal components). The bold line outlines the San Francisco (SF) Bay-Sacramento River Delta region and the coastline. Seismographic stations are absent from the region between 38.2 and 38.4⬚ N for longitudes east of 122.2⬚ W. The dashed oval denotes The Geysers geothermal area and the thick-dashed line shows the approximate position of the Coast Range–Great Valley boundary. The dashed rectangle outlines the location of the 3D velocity inversion grid.

Great Valley

39.0 The Geysers area Lake Berryessa

20 km

38.8

Latitude

38.6

38.4

38.2

SF Bay

38.0 Pacific Ocean

37.8 -123 -122.8 -122.6 -122.4 -122.2 Longitude

Sacramento River Delta -122 -121.8 -121.6

algorithm is used to solve the velocity-station correction inverse problem. Hypocentral relocations are performed with progressive undamped residual-weighted least squares. The horizontal (x, y) coordinates are recomputed first, before allowing all four hypocentral parameters to vary. Trial elevations of 5 and 10 km below sea level were used in addition to the most recent hypocenter elevation during each earthquake relocation to reduce the probability of a hypocenter converging to a local elevation minimum. The hypocentral equations are inverted using QR factorization (Householder orthogonalization). Throughout the simultaneous hypocenter-velocity inversion, 3D ray tracing is implemented with the iterative ray bending technique of Um and Thurber (1987), as modified by Block (1991). A 3D velocity grid was established to encompass most of the seismicity and stations in the region (Fig. 5). To accommodate the wide range of station elevations (sea level to 1.5 km), the z grid was established relative to mean sea level, with negative values denoting depth below mean sea level. An x-y grid node spacing of 5 km was used throughout the grid. Elevations nodes were located at 1.5, 0, ⳮ1.5, ⳮ3.0, ⳮ5.5, ⳮ8.0, ⳮ14.0, ⳮ20.0, and ⳮ22.0 km. Linear velocity interpolation was used between nodes for traveltime calculations. Station and earthquake distributions were

sparse along the edges of the grid. Consequently, the two northernmost and southernmost rows of nodes were constrained to vary only in the east-west direction, and the two easternmost and westernmost columns of nodes were constrained to vary only in the north-south direction. Three L1-norm 3D velocity-hypocenter-station correction inversions were performed with arrival-time data from 351 microearthquakes recorded from 1982 to 1998. After each of the first two inversions the arrival-time picks associated with the large (⬎0.2 sec) residuals were checked. Very few of the arrivals times with large residuals were changed. Instead, these arrival times were deleted because analog network telemetry artifacts were apparent, signal-tonoise ratios were too low to justify repicking arrival times, or, in the case of S-wave arrivals, phase identification was too ambiguous. The broad tails of P- and S-wave arrival-time residual distributions are consistent with the L1 assumption for firstarrival-time picking uncertainties. The median absolute value of P residuals (median ||r||; Table 1), is about twice the estimated P picking precision, which suggests the inversion is able to fit P-arrival times roughly consistent with P picking uncertainty. The median absolute value of S residuals is more than double that of the P residuals. Conse-


Table 1 Arrival-Time Residual Statistics Arrival Type

Mean (sec)

Std. Dev. (sec)

Skew

Kurtosis

Median |r| (sec)

L1 r (sec)

P wave S wave

0.012 0.0

0.15 0.80

3.0 30

70 875

0.037 0.092

0.06 0.12

quently, the S data provide little additional constraint on hypocenters relative to using only the P data (Gomberg et al., 1990). The skews and kurtoses are substantial (Table 1). P residuals are skewed toward late picks, consistent with a bias toward missing P first arrivals in low signal-to-noise situations and picking more energetic, later arrivals. Conversely, S residuals are skewed toward early picks, consistent with a weak, pre-S converted phase picking bias. Estimated velocities should have minimal biases since both sets of residuals have nearly zero means and medians. P-velocity resolution is best in the central portion of the grid over the elevation range of ⳮ2 to ⳮ16 km. P-velocity resolution is low and spatially limited below the ⳮ16-km elevation. Seismographic station coverage is sparse in and near the Sacramento Valley, further degrading the earthquake-location capabilities in that region. Consequently, hypocentral elevation uncertainties are about 1–2 km for earthquakes beneath the Coast Ranges and about 3–6 km for earthquakes beneath the Sacramento Valley. Details of the velocity transition along the Coast Range–Sacramento Valley boundary can only be resolved between latitudes of 38.4⬚ N and 38.5⬚ N, and full 3D resolution is confined to longitudes west of 122.1⬚ W due to the generally sparse distributions of stations east of 122.1⬚ W (Fig. 5). The situation has changed little since the assessment by Eaton (1986) that locating earthquakes beneath the Sacramento Valley is a difficult task due to poor station coverage. Analysis and Interpretation of 3D Velocity Structure. We compared east-west cross sections of velocity structure through the 3D model (Fig. 6) with geologic cross sections constructed from the 1:250,000 scale geologic map of this region by the California Division of Mines and Geology (Wagner and Bortugno, 1982), supplemented by strike and dip data from Fox et al. (1973) and Sims et al. (1973). The major lithotectonic units that comprise the crust at this latitude formed at a long-lived convergent margin (Dickinson, 1981; Ernst, 1981). From east to west, these units are: (1) Mesozoic strata of the Great Valley group and overlying Tertiary strata, which were deposited in a marine forearc basin; (2) ultramafic rocks of the Coast Range ophiolite, which formed the forearc basement; and (3) the Franciscan complex, which represents the remnants of an accretionary prism (Dickinson and Rich, 1972; Dickinson, 1981; Ernst, 1981). The east-west cross sections through the velocity model in Figure 6 depict gross crustal structure consistent with monoclinal folding of these lithotectonic units along the

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range front. The deformed structural boundary that separates the Franciscan complex from the Coast Range ophiolite and Great Valley group strata is associated with a west-to-east lateral velocity gradient. In detail, isovelocity lines indicate that this gradient is coincident with the east-facing homoclinal flexure of rocks with Franciscan velocities (i.e., about 5.1 km/sec and greater) in the upper 10 km of the crust. West of longitude 122.20⬚ W, rocks of Franciscan velocities are at or near the surface, and isovelocity lines in the upper 4–5 km of the crust are subhorizontal with minor fold relief. In the vicinity of longitude 122.20⬚ W, isovelocity lines dip east and rocks with Franciscan velocities abruptly increase in depth to about 10-km depth beneath the range front. Total structural relief on rocks with velocities of about 5.5 km/sec across the monocline is about 7 km, comparable to relief on the basal Great Valley group strata estimated from exposed stratigraphic thickness and projection of surface exposures to depth. Given sufficient resolution the velocity model could be used to determine whether tectonic wedging or faultpropagation folding is responsible for development of the range-front monocline. In the tectonic wedge model, the wedge is composed of higher-velocity rocks that are thrust up and over a subhorizontal section of lower-velocity rocks (Fig. 3). The depth of detachment for the wedge is marked by the point at which the isovelocity lines diverge from parallelism (i.e., the contact between units V4 and V5 in Fig. 3), which also is the wedge tip. A simple kinematic tectonic wedge model fit to the east-west velocity structure is shown in Figure 6b. For the fault-propagation fold model the tip of the thrust fault is located at the base of the monocline where isovelocity lines diverge from parallelism (i.e., the contact between units V5 and V6 in Fig. 4). The fault dips beneath the monocline and predicted displacement of high-velocity rocks over lower-velocity rocks is less pronounced than in the tectonic wedge model. Also, the region of uplifted higher-velocity rocks should extend to the edge of the range-front monocline. Based on these assumptions a faultpropagation fold model is fit to the velocity structure in Figure 6a. Of the two models we favor the fault-propagation fold model to explain the monoclinal structure depicted in the velocity model. The tectonic wedge interpretation requires a subhorizontal detachment to be located within the Franciscan complex. Although this is not impossible, detachments typically develop parallel to planar anisotropies in rock bodies or along through-going zones of weakness. We infer that such anisotropies in the Franciscan complex are likely to reflect deformation within the ancestral western California subduction zone and thus dip moderately to steeply east. If present, we do not expect such zones to be optimally oriented to serve as subhorizontal detachments. Also, there is no clear expression of a wedge of high-velocity rocks being thrust over lower-velocity rocks in the velocity model, as predicted by the wedge model (Fig. 3). Although the resolution of the velocity model may not be sufficient to preclude

1478


West

(a)

East Latitude = 38.499°N COAST RANGES

Maacama Fault

-122.70

-122.60

KJf

-122.40

Longitude -122.30

-122.20

-122.10

TQ

0

Elevation (km)

JK gvg -122.50

Monticello Dam (projected from the north)

Wragg Canyon Fault

KJf

-122.00

Ba

KJf

Ku gvg

se

Venado Fm

KJf

-5

-121.90

JK gvg

-10

KJf

-15

Base

-20 -30

-20

-10

0

10

20

30

Distance from 122.3°W (km) Projection of the Wragg Canyon Fault

West

(b)

Latitude = 38.463°N

Northern projection of the Green Valley Fault

COAST RANGES -122.70

-122.60

-122.50

East

Monticello Dam (projected from the north)

SACRAMENTO VALLEY

Longitude -122.40

-122.30

-122.20

-122.10

JKgvg

0

-121.90

TQ Ku gvg

Ba

se

?

JKf Elevation (km)

-122.00

Venado Fm

-5

JK gvg

"Tectonic Wedge"

JKf

-10

-15

-20 -30

-20

-10 0 Distance from 122.3°W (km)

10

20

30

Explanation TQ - Tertiary-Quaternary rocks Ku gvg - Upper Cretaceus Great Valley Group JK gvg - Jurassic-Cretaceus Great Valley Group KJf - F ranciscan Assemblage

2

3

4 5 6 P-wave Velocity (km/s)

Figure 6. Simplified geology superimposed on east-west cross sections at (a) 38.499⬚ N and (b) 38.463⬚ N through the P-wave 3D velocity model. The solid circles are microearthquake hypocenters located within 2 km of the cross sections.

7

8

1479


the presence of a tectonic wedge, we conclude that the velocity model more closely resembles the predicted velocity structure for fault-propagation folding shown in Figure 4. Seismic Reflection Data Eight proprietary 2D seismic lines (totaling about 135 linear kilometers) were acquired and processed to produce depth-migrated seismic images of the range-front structure. Because these data are proprietary, it is not possible to show the precise positions of the seismic lines or to completely reproduce the seismic reflection data in this article. Consequently, the eight dip lines were combined into four general dip profiles (from the north to the south, labeled as seismic profiles A-D; Figs. 7–10, respectively). A single strike line is labeled profile S (Fig. 11). Processing Approach. All the original reflection seismic data were collected using Vibroseis sources. Recording durations allowed 5 sec of two-way travel time to be recovered for line A (Fig. 7), 6 sec for profiles B, D, and S (Figs. 8, 10, and 11, respectively), and 7 sec for profile C (Fig. 9). The surface-consistent processing scheme was designed to avoid artifacts associated with automatic statics and coherency statics processing. The general processing scheme consisted of Vibroseis correlation, common midpoint binning (including crooked-line geometry corrections, where nec-

essary), amplitude recovery, refraction statics, surgical and trace noise mutes, interactive velocity analysis, and dipmoveout stacking. The velocity models derived from the velocity analyses were spatially smoothed using 0.5-1.0 km smoothing dimensions to produce interval velocity models for depth migration. Although finite-difference prestack migration was used for several lines it was found that Kirchoff poststack depth migration produced comparable results. Consequently, Kirchoff poststack depth migration was used to produce the final depth migrations. All processing was performed using ProMax 2D version 7.0. Near-surface refraction P velocities obtained from the seismic reflection data showed shallow refractor velocities of approximately 1.7 km/sec at the eastern limits of the lines near Davis and Woodland. Refractor P velocities increase almost linearly with distance westward to velocities of approximately 2.7 km/sec about 1 km east of the eastern slope of the English Hills. Refractor velocities in the English Hills increase from ⬃3 km/sec along the eastern margins to approximately 3.7 km/sec along the eastern edge of the Coast Ranges. This westward-increasing lateral velocity gradient is reproduced in the interactive velocity analysis of the dip lines (profiles A-D) for two-way travel times extending to ⬃3 sec (corresponding to depths of 5–6 km). A sonic P velocity log to about a 6-km depth was available within several kilometers of two dip lines and the strike

West

East

0

Cenozoic, undivided Upper Cretaceous

Elevation (km)

-2

-4

Onlap

-6

unconformity

-8

F -10

TC

“Faulted unconformity” -12 -10

-5 0 5 Distance from ~122°W along N70W (km)

Figure 7. Thin curves show tracings of prominent reflectors from a depth migration of seismic line A, a dip line located in the northern portion of the Trout Creek blind thrust segment. The tracings from the original lines were projected along a line extending N70⬚W to obtain a uniform dip profile. Faults and unconformities are shown by bold lines, dashed were inferred. Gross geologic ages on the units are as labeled. The thin-dashed line shows a hinge line associated with fault-propagation folding. The approximate position of the top of the Trout Creek thrust fault is labeled as TCF.

10

1480


Fault-propagation fold West

East

0

Cenozoic

Elevation (km)

-2 Upper Cretaceous

-4

Onlap

-6 Unconformity -8 “Faulted unconformity”

TCF

-10 -12 -15

-10

-5

0

5

10

Distance from ~121.9°W along N70W (km)

Figure 8. Thin curves show tracings of prominent reflectors obtained from a depth migration of seismic line B, a dip line located in the central portion of the Trout Creek blind thrust segment. The line tracings from the original lines were projected along a line extending N70⬚W to obtain a uniform dip profile. Faults and unconformities are indicated by bold lines. Gross geologic ages on the units are as labeled. The thindashed line shows the eastward hinge line of the fault-propagation fold. The approximate position of the top of the Trout Creek thrust fault is labeled as TCF.

profile (Fig. 12). Velocities obtained via interactive velocity analysis were virtually identical to those obtained from the sonic log. Sonic log P velocities are also similar to nearby P velocities estimated in the 3D velocity-hypocenter inversion. The 3D model correctly resolves the long-wavelength retical velocity gradient near the well (Fig. 12), and absolute velocity differences between the sonic and 3D model velocities are less than 10% below 1-km depth. The 3D model isovelocity contours follow along-strike depth variations of the Upper Cretaceous strata (Fig. 11) throughout the western margin of the Sacramento Valley and show isovelocity contours climbing upsection to the west through the monocline (Fig. 6) on the eastern margin of the Coast Ranges. Although the short-wavelength velocity details are not resolved by the 3D model (Fig. 12), the 3D velocity model provides large-scale constraints on velocity variations with depth and horizontal position suitable for constructing velocity relations for depth migration in the 6–12-km depth interval. Consequently, interactive velocity analysis was used to construct P velocities from the surface to a 6-km depth, and the 3D velocity model results were used to constrain P velocities below a 6-km depth. Based on the agreement between the independent estimates of P velocities discussed previously, estimated long-wavelength depth uncertainties in the depth migrations are less than 5% for depths less than 8 km and less than 10% for 8- to 12-km depths.

Interpretation of the Depth-Migrated Seismic Reflection Sections. The four dip lines (Figs. 7–10) all display similar characteristics of the subsurface structure of the southwestern Sacramento Valley. East-dipping layered reflectors in the upper 8–10 km at the west ends of the lines project westward to outcrops of the east-dipping Great Valley group strata exposed in the range-front monocline. Reflectors associated with east-dipping strata are juxtaposed against subhorizontal to west-dipping reflectors at a depth of about 9–10 km. The discordance in reflector dip occurs across a gently westdipping reflector that can be traced eastward beneath the Sacramento Valley. Near the centers of the dip lines the eastdipping strata above this reflector flatten across a synformal hinge that marks the base of the range-front monocline. East of this hinge subhorizontal reflectors associated with strata in the upper 5–6 km appear to form an on-lapping relationship with the west-dipping reflector. Localized short wavelength, low-amplitude folding is evident in the angular unconformity between the Upper Cretaceous and Eocene strata east of the synformal hinge at the base of the range-front monocline. We interpret the subhorizontal to west-dipping reflector beneath east-dipping Upper Cretaceous strata to be a detachment surface that has accommodated east-vergent blind thrust faulting beneath the range front; this detachment has transmitted minor displacements east of the range-front monocline into the southwestern Sacramento Valley. Be-


West

Cenozoic, undivided East

0

Upper C

retaceo

us

Elevation (km)

-5

April 21 GVF

-10 GVF 19 April TCF

-15 -4

-2

0

2

4

Distance from ~122°W along N70W (km) Figure 9. Thin curves show tracings of prominent reflectors from a depth migration of seismic line C, a dip line located in the segmentation overlap of the Gordon Valley (GVF) and Trout Creek (TCF) blind thrust segments. See Figure 7 for labeling and projection information. The northern GVF bifurcates into two segments associated with the 19 April 1892 mainshock and 21 April 1892 aftershock. An out-ofsyncline backthrust is shown as the dashed line. An inactive older thrust is shown as a dash-dot line.

cause most of the seismic lines we examined do not extend west of the valley margin into the outcrops of the Upper Cretaceous rocks, it is difficult to determine the geometry of the detachment surface beneath and west of the range front. In our interpretation of profile A (Fig. 7) the westward dip of the detachment surface steepens beneath the range-front monocline because we assume that the westward increase in structural relief is associated with an increase in fault dip. Additional evidence that the detachment dips moderately west beneath the monocline and flattens eastward is the development of a small fault-propagation fold directly east of the range-front monocline in profile B (Fig. 8). As strata are thrust eastward from the steeper part of the detachment onto the subhorizontal part, bedding in the hanging wall

1481

form a small fault-bend fold as indicated by east-dipping reflectors truncated by the detachment east of the change in dip. A short, east-dipping backthrust transmits displacement on the detachment back to the west, where it is consumed in the formation of a west-vergent fault-propagation fold directly east of the range-front monocline (Fig. 8). From the relationships in Figure 8 we infer that the lowangle detachment at about 8 km beneath the valley margin extends westward beneath the range front and steepens below the range-front monocline. We thus interpret the rangefront monocline to be an east-vergent fault-propagation fold, consistent with our inference from analysis of the 3D velocity model (i.e., Fig. 6b). Most of the reverse displacement on the detachment apparently has been consumed in growth of the monocline, but relatively small displacements have been transmitted eastward, resulting in minor folding of Cretaceous strata above the detachment, as well as in folding of the basal Eocene angular unconformity. Through field reconnaissance, we confirmed that many of the low-amplitude folds east of the range front imaged on the seismic lines have minor geomorphic expression as very low hills in the Sacramento Valley. These hills are underlain by deformed late Quaternary valley-fill deposits. From this we infer that active thrust faults or detachments are present east of the range front and may have been the source for the 30 April aftershock of the 1892 earthquake sequence (discussed in greater detail subsequently). The north–south strike profile (Fig. 11) provides additional information about the continuity and geometry of thrust faults beneath the western valley margin. We used our interpretations of the detachment on the dip profiles to locate the detachment on strike profile S and trace it laterally to the north and south. North of the town of Winters the detachment is expressed on the strike profile as a discontinuity in the layered Great Valley group rocks in the 8- to 10-km depth range that exhibits relatively higher amplitude (Fig. 11). The detachment appears to rise gradually through the section from about 9-km depth at Winters to about 7-km depth at the northern end of the strike profile. The detachment was also located on the strike line from an intersection with dip profile D south of Winters (Fig. 10). Here, the detachment is located at a depth of about 8.5 km and it rises northward toward Winters where the expression of the fault becomes lost or obscured in a low-amplitude fold. We interpret this geometry to be the 2D expression of a right en-echelon step in the detachment system and axis of the range-front monocline at Winters. The structures with surface expression of uplift above the blind thrust faults (i.e., the English Hills) trend northwest, whereas the strike line trends north-south. Moving from south to north, the strike line crosses the northwest-trending English Hills obliquely such that the underlying thrust fault appears to rise northward. On approaching Winters, the displacement on the fault is primarily consumed in monoclinal folding of the English Hills. The strike line crosses the fold and begins to approach the axis of the next en-echelon segment of the range-front

1482


Pliocene-Pleistocene East

West 0

Eocene Elevation (km)

-2 Upper Cretaceous

-4

Onlap (on unconformity ?)

-6 -8 -10

Approximate eastern limit of slip on detachment

GVF

-12 -15

-10

-5 0 5 Distance from ~121.8°W along N70W (km)

10

15 “Basement”: deformed lower Great Valley group strata?

Figure 10. Thin curves show tracings of prominent reflectors from a depth migration of seismic line D, a dip line located in the southern portion of the Gordon Valley blind thrust segment. The line tracings from the original lines were projected along a line extending N70⬚W to obtain a uniform dip profile. Faults and unconformities are the bold curves, dashed where inferred. Gross geologic ages on the units are as labeled. The approximate position of the Gordon Valley thrust fault is labeled as GVF.

monocline north of Winters, where the fault also appears to rise northward in the plane of the strike section because the fault strikes oblique to the seismic line. The interpreted en-echelon step at Winters is well expressed in the map-scale topography of the outcrop belt. The crest of the Vaca Mountains, which forms the prominent north-northwest-trending strike ridge approximately 10 km west of Vacaville, Allendale, and Winters (Fig. 1), is underlain by the northeast-dipping Venado sandstone of the Great Valley group. The Venado sandstone forms the foundation of the Monticello Dam (Fig. 1), and is fully involved in the monoclinal folding along the range front. Moving southeast to the northwest along the Vaca Mountains, the Venado sandstone strike ridge bends to the northeast at the latitude of Allendale and resumes a northwest trend at about the latitude of Winters. The large bend in the Venado sandstone is coincident with an eastward step in the locus of the Quaternary uplift along the range front at Winters, which is expressed by the low hills (i.e., about 200-m elevation or less) bordering the southwestern Sacramento Valley (Fig. 1). In particular, a right en echelon step in the crest line of the low hills at the latitude of Winters is well expressed in the perspective view of the topography in Figure 2. The strike of the Venado sandstone and the trend of the crest line of low hills bordering the valley both are more westerly than the general north-northwest trend of the Coast Range Mountains. If it is assumed that the topography reflects the firstorder geometry of the active thrust faults at depth, then these

relations are characteristic of a right-stepping en-echelon pattern of blind thrust faults beneath the range front.

Geometric Model for Blind Thrust Faults Beneath the Southwestern Sacramento Valley Based on analysis of the 3D velocity structure and seismic reflection images, we interpret the northeast-facing monocline along the boundary between the Coast Ranges and southwestern Sacramento Valley as the surface expression of fault-propagation folding rather than tectonic wedging. The monocline has developed above blind thrust faults that dip southwest beneath the Coast Ranges. The thrust faults flatten eastward to form a subhorizontal to gently west-dipping detachment beneath the southwestern Sacramento Valley at a depth of 8–9 km. This geometry is characteristic of a ramp-flat transition. The presence of lowamplitude folds in Cretaceous and Tertiary strata above the flat in the western Sacramento Valley indicate that some displacement has been transmitted along this surface east of the range front. However, the majority of structural relief on Great Valley group and Franciscan rocks, as well as the locus of late Cenozoic uplift and eastward tilting, occurs at the range front. From this, we interpret that the majority of coseismic slip primarily occurs on the more steeply westdipping thrust ramps beneath the eastern Coast Ranges. We further interpret that two potentially seismogenic thrust ramps are separated by a right en-echelon step at the

1483


latitude of the town of Winters. The southern ramp extends between the towns of Winters and Vacaville and terminates in a fold visible on the north-south strike profile S (Fig. 11). For convenience we refer to this southern ramp as the Gordon Valley segment of the range-front thrust system. Based on features visible in strike profile S and the topographic limit of the fault-propagation fold in the southern English Hills, the total length of the Gordon Valley segment is 18 Ⳳ 2 km. The northern ramp extends from Winters northward, and we refer to it as the Trout Creek segment. The Gordon Valley segment appears to bifurcate into two thrust splays about 10 km from its northern terminus (Fig. 11). This interpretation provides an explanation of the reported Modified Mercalli intensities associated with the 19 and 21 April 1892 earthquakes as discussed in a following section. Bifurcation of the thrust fault provides an explanation for the complex strike-perpendicular folding pattern in the ⳮ3 to 3-km distance range on the south flank of the large-scale segmentation anticline located between distances of 0–5 km in Figure 11. The bifurcation is consistent with the strike change in the fault-propagation fold shown in Figure 2. As discussed in greater detail subsequently, we interpret

that structural segmentation of the thrust system due to the en echelon geometry has resulted in the formation of a rupture barrier at the boundary between the Gordon Valley and Trout Creek ramp segments. More specifically, the structure at the segment boundary is characterized by en-echelon offsets of late Quaternary monoclinal folds and distinct depth offsets between adjacent thrust sheets (Fig. 11) and the rightstepping en-echelon offsets of the fault-propagation folds (Fig. 2).

Elastic Deformation Modeling We employ elastic deformation modeling to assess how well the interpreted positions and geometries of blind thrust faults reproduce present-day topography along the boundary between the western Sacramento Valley and eastern Coast Ranges mountain front, including the epicentral region of the 1892 earthquake sequence in the English Hills. We assume that the current long-wavelength topography in the English Hills is produced by static deformation associated with repeated episodes of reverse faulting. Elastic deformation modeling is performed using the methods of Manshinha and Symlie (1971) and Okada (1985) to calculate the static elas-

Cenozoic, undivided

South

North

0 Elevation (km)

-2

Upper Cretaceous

-4 -6 -8

92

ril 18

Ap 19 April 1892 21

19 April 1892

-10 G ordo n Va lle y -12 -10

s e gme nt

C re e k

Tro ut

0 10 Distance from ~38.5°N along N20°W (km) 2

P-wave velocity (km/s) 3 4 5 6

s e gm

7

20

8

Figure 11. Reflector tracings (thin lines) from the depth-migrated line S, obtained near the eastern margin of the English Hills. The corresponding P-velocity profile from the 3D velocity-hypocenter inversion is superimposed in color. Thrust fault segments are labeled bold lines. The triangles show nearby seismic network stations within ⬃10 km of the profile, with stations east of the profile gray shaded. Most of the interpretation of the Gordon Valley segment is derived from lines C and D (Figs. 9 and 10) and inspection of the proprietary wiggle trace data (not shown). The two Gordon Valley blind thrust segments involved in the April 1892 Vacaville-Winters earthquake sequence are labeled as discussed in the text. Note the north-vergent monoclinal folding in the Cretaceous and Cenozoic strata above the Gordon Valley-Trout Creek segmentation boundary.

e nt


0

(a)

1

0

2

Elevation (km)

Depth (km)

1484

Sonic log

3

38.432°N Elevation plotted at 7.5X exaggeration

-5

Gordon Valley thrust segment

-10 -15 0

4

121.80°W

122.25°W

10

20 30 Grid distance east (km)

(b) 3D Model

2.0 2.5 3.0 3.5 4.0 4.5 Interval P-wave velocity (km/s)

121.80°W

0

5.0

Figure 12. Smoothed sonic P velocities (solid line) from the Great Basins well (38.64⬚ N, 121.89⬚ W) and averaged P velocities from nearby nodes in the 3D velocity-hypocenter inversion model (dashed line).

tic vertical displacements at the surface due to subsurface faulting. The shear modulus is set to zero to obtain the longterm deformation field. Two east-west topographic profiles were obtained from USGS 3-arcsec topography to evaluate the surface expression of folding associated with slip on the Gordon Valley and Trout Creek thrust fault segments (Fig. 13). Limited data are available to interpret the western limb of the fault-propagation folds within the English Hills (e.g., Fig. 8), so the analysis was directed toward finding the fault dips that produced the best match to the long-wavelength topography on the eastern limb of the fault-propagation folds. The depths of the thrusts were inferred from the seismic reflection data, and it was assumed that the thrusts extend to the 14-km base of the seismogenic zone in the Coast Ranges (Wong et al., 1988). The best-fitting fault dips obtained from the models are 30⬚ for the Gordon Valley segment and 20⬚ for the Trout Creek segment. Elevated topography extends further east adjacent to the Trout Creek segment relative to the Gordon Valley segment (Figs. 1 and 13), consistent with a lower dip on the Trout Creek segment relative to the Gordon Valley segment. Although these models are nonunique, the ability to reproduce first-order observed topography along the western margin of the Sacramento Valley provides additional confidence in our interpretation of the thrust fault geometries based on analysis of the 3D crustal velocity model and seismic reflection data.

Ground-Motion Modeling Approach

Elevation (km)

6 1.5

50

38.572°N

122.25°W

5

40

Elevation plotted at 7.5X exaggeration -5

Trout Creek thrust segment -10 -15 0

10

40

50

Figure 13. Observed topography (thick lines) and predicted elevation change (dashed gray lines) resulting from repeated uniform slip blind thrust earthquakes for preferred fault positions (thick lines with slip direction arrows) as a function of distance east from a position west of the Monticello Dam for (a) the Gordon Valley thrust segment and (b) the Trout Creek thrust segment. The starting and ending longitudes of the topography profiles are as labeled and the latitude of each section is shown at the top of the section.

compute near-field seismograms for finite-fault rupture models. Isochrones are all the positions on a fault that contribute seismic energy that arrives at a specific receiver at the same time. The simplest way to employ the isochrone method in the near field is to assume that all significant seismic radiation from the fault consists of direct S-wave arrivals. This assumption is reasonable in the near field, particularly for a deeply buried, blind thrust fault that produces dominantly near-vertical source-receiver paths. A further simplification is to use a simple trapezoidal slip-velocity pulse. Let f (t) be the slip function, For simplicity we assumed ¨f (t) ⳱ d(tⳮ tr) ⳮ d(tⳮth) , where tr is rupture time and th is healing time. Then, all seismic radiation from a fault can be described with rupture and healing isochrones. Surface ground-motion velocities, v, and accelerations, a, from rupture or healing can be calculated from (Spudich and Frazer, 1984) v(x, t ) = f (t ) ⊗

∫ (siG)c dl

(1)

y ( t ,x )

a(x, t ) = f (t ) ⊗



∫ c

y ( t ,x )

Bernard and Madariaga (1984) and Spudich and Frazer (1984, 1987) developed the isochrone integration method to

20 30 Grid distance east (km)

 ds  2  dG  dc is  + i(siG) dl (2)  iG  + c   dq  dt  dq  

2

where c is isochron velocity, s is slip velocity (either rupture

1485


or healing), dq is the spatial derivative, y(t,x) defines the isochrone, and G is a hybrid ray theory Green’s function combined with synthetic 3D scattering functions described subsequently. Spudich and Frazer (1984) showed that c can be eliminated from equation (1) by integrating along an isochrone over a finite time window defined by the isochrones t ⳮ dt and t Ⳮ dt. By limiting the integration to frequencies lower than 10 Hz (dt ⳱ 0.05 sec), band-limited ground motions are obtained, making the isochrone method useful for qualitatively evaluating accelerations. This approach was used here with a first-order approximation where equation (1) was reduced to a point-source summation of over the isochrone strip corresponding to the finite time window:

(3)

where B(t) ⳱ [H(t Ⳮ 1) ⳮ H(t ⳮ 1)]/2 is the boxcar function and ts is the isochrone. We used this approach to efficiently calculate seismograms for a large number of receiver positions relative to a fault. For the isochrone integrations the ray theory portion of G was approximated as G s ( x, ξ ) =

F ( x, ξ )W ( x, ξ ) 4

ρ ( x ) ρ (ξ ) β ( x ) β (ξ )5 / 2 ( x − ξ )2

Range of receiver positions

where x is the receiver position, n is position on the fault, ˜ F(x,n) is the source radiation term, W(x,n) is the free surface amplification factor, q is the density, and b is the S-wave velocity. In this approximation only first S-wave arrivals are included in the calculation. Although ray spreading factors are simply approximated by the inverse source-receiver distance, a 1D approximation to the 3D S-wave velocity model beneath the English Hills from the velocity-hypocenter inversion was used to calculate, b(n), takeoff angles for F(x,n), ˜ and incidence angles for W(x,n) to incorporate first-order geometric effects of vertically heterogeneous velocity structure which varied from b ⳱ 1.7 at the surface to b ⳱ 3.5 km/sec at 9-km depth. This produced realistic partitioning of SV velocities between vertical and horizontal components of ground motion. The amplitude effects of SV transmission were approximated by computing the median vertical-incidence amplification of a band-limited (0.5–5 Hz) SV plane wave propagated through a 3D heterogeneous media. A self-similar, fractal correlation model of random spatial variations of crustal seismic velocities with an autocorrelation function, P, of the form P ( kr ) ≈

an 1 + ( kr a ) n

(5)

where a is the correlation distance, kr is the radial wavenumber, and n ⳱ 3 was used in a 3D elastic finite-difference

East

West

(4)

0

Elevation (km)

1  (t − t s )  v(x, t ) = f (t )∫∫sr iG s B   dA, dt  dt  A

calculation to compute SV transmission amplification through the top 9 km of the crust. The free surface was omitted since equation (4) was used to calculate free-surface amplification. The 3D randomization of the 1D velocity model was normalized to produce a standard deviation of 5% of b for akr k 1 with a ⳱ 2.5 km. O’Connell (1999a) showed that 3D scattering in the upper crust can have a significant influence on the scaling of near-field peak ground motions. Scattering of direct SV waves by correlated-random velocity variations tends to reduce peak horizontal velocities (PHVs) associated with nearfield rupture directivity by increasing the time and phase dispersion of direct SV waves. Scattering is included here using the 3D elastic finite-difference program of Graves (1996) to produce synthetic three-component SV scattering functions. SV plane waves were propagated at incidence angles of 80⬚ and azimuths of 45 and 135⬚ relative to the strike of the steeply dipping Upper Cretaceous rocks in the Coast Range through the upper 2 km of the crust using the 3D model of velocity structure near the Monticello Dam derived by O’Connell (2001) (Fig. 14). Three-component SV scattering functions were obtained for receiver positions within

-1

-2

-3

SV pla ne wav e

-4 -2

-1 0 1 2 Downstream distance (km) 1.5 2.0 2.5 S-wave velocity (km/s)

Figure 14.

East-west-depth S-wave velocity profile from a 3D randomization of a 2D velocity model for the top 2 km of the crust below the Monticello Dam with n ⳱ 2 and a ⳱ 1 km in equation (5). A 2km-thick homogeneous region is inserted at the bottom to introduce a uniform amplitude SV plane wave. The scattering functions for the ground-motion simulations were derived using n ⳱ 3 and a ⳱ 2.5 km to produce the 3D randomization of the 2D velocity model. The standard deviation of the velocity randomization was set to 5% for all cases. The horizontal extent of the east-west region used to extract scattering functions is shown at the top.

1486

A kinematic rupture model is used that mimics the spontaneous dynamic rupture behavior of a self-similar stress distribution model of Andrews and Boatwright (1998). The kinematic rupture model is also similar to the rupture model of Herrero and Bernard (1994). Self-similar effective stresses (and slip velocities) are generated over the fault with rise times that are inversely proportional to effective stress. Peak rupture slip velocities evolve from ratios of one relative to the sliding (or healing peak) slip velocity at the hypocenter to a maximum ratio of 4:1 consistent with the dynamic rupture results of Andrews and Boatwright (1998) that shows a subdued Kostrov-like growth of peak slip velocities as the rupture grows over a fault. The kinematic model used here produces slip models with 1/k2 distributions consistent with

Fourier velocity Fourier velocity Fourier velocity (cm/s/Hz) (cm/s/Hz) (cm/s/Hz)

velocity (cm/s) velocity (cm/s)

10-1 10-2

East

10-3 10-4 10-1

1

10-2

10 North

10-3 10-4 10-1

1

10

10-2 Vertical

10-3 10-4

1 10 Frequency (Hz)

Example of a synthetic three-component Monticello Dam SV scattering function with normalized velocity waveforms on the left and Fourier velocity spectra on the right for the components as labeled.

1.0 0.5

East

0.0 -0.5 0.0 0.5 1.0 1.5 2.0 0.5 0.0 North -0.5 -1.0 -1.5 0.0 0.5 1.0 1.5 2.0 1.0 Vertical 0.5 0.0 -0.5 -1.0 0.0 0.5 1.0 1.5 2.0 Time (s)

Fourier velocity Fourier velocity Fourier velocity (cm/s/Hz) (cm/s/Hz) (cm/s/Hz)

Rupture Modeling Approach

1.0 East 0.5 0.0 -0.5 -1.0 0.0 0.5 1.0 1.5 2.0 0.4 North 0.2 0.0 -0.2 -0.4 -0.6 -0.8 0.0 0.5 1.0 1.5 2.0 0.4 Vertical 0.2 0.0 -0.2 -0.4 -0.6 0.0 0.5 1.0 1.5 2.0 Time (s)

Figure 15.

velocity (cm/s) velocity (cm/s) velocity (cm/s)

the high-velocity portion of the lower Venado Formation as shown in Figure 14. The depth limit of 4 km in the 3D velocity model was dictated by the need to produce scattering functions to a maximum frequency of 10 Hz, the horizontal dimensions necessary to sample a 2.5-km correlation distance adequately (12 km), and the limitations of fitting a 3D elastic finite-difference calculation into computer memory. The 3D scattering functions were normalized so that the median peak velocity for each component matched the median SV transmission amplifications derived earlier. Geometric spreading and free-surface amplification were applied using equation (4). A 3D scattering function was chosen at random at each integration position in equation (5) from a total of 5200 scattering functions used in the ground-motion simulations. The scattering functions possess fairly simple waveforms (examples are shown in Figs. 15 and 16). The scattering function responses are consistent with the rather simple site responses observed at the Monticello Dam from an earthquake located at 11-km depth, about 5 km from the dam (Fig. 17); coda durations are relatively short, and direct S waveforms are fairly simple. Shear-wave splitting was commonly observed for sites located on Upper Cretaceous Great Valley sedimentary rocks near the Monticello Dam (Fig. 17). Consequently, kinematic time shifting of horizontal components was used to delay east-west-polarized, direct S waves relative to north-south components using the method described in O’Connell (1999c). Sokolov and Chernov (1998) showed that seismic intensities correlate well for rather narrow ranges of Fourier amplitude spectra of ground acceleration, with 0.7–1.0 Hz being most representative of Imm ⱖ VIII, while the 3–6-Hz range best represents Imm in the V-VII range. The frequency bandwidth was limited to 5 Hz to accommodate the wide range of site conditions and frequency band of interest for sites near the range front, and to facilitate comparison of the PHVs and peak horizontal accelerations (PHA) with intensities in the Imm ⳱ V-IX range.

velocity (cm/s)


10-1 10-2

East

10-3 10-4 10-1 10-2 10-3 10-4 10-5 10-1

1

10 North

1

10

10-2 10-3 10-4

Vertical 1 10 Frequency (Hz)

Figure 16. Another example of a synthetic threecomponent Monticello Dam SV scattering function with normalized velocity waveforms on the left and Fourier velocity spectra on the right for the components as labeled.

estimates of earthquake slip distributions (Somerville et al., 1999) and x2 displacement spectra in the far field. To calculate peak slip velocities, healing slip velocities, and displacements for moment calculations, we used an approximate Kostrov slip-velocity function from O’Connell and Ake (1995)


velocity (µ/s)

4

Figure 17.

EQ1 P-wave

2

EQ2 P-wave

0 -2

Vertical

-4 0

2

4 Time (s)

6

8

velocity (µ/s)

8 EQ1 fast S-wave EQ2 fast S-wave

4

Monticello Dam left abutment three-component broadband velocity waveforms from two 3 March 1998 M ⬃ 1.8 blind thrust earthquakes located 5 km northwest of the Monticello Dam at an elevation of -11 km. The P- and S-wave direct arrivals from each earthquake are labeled as EQ1 and EQ2. The vertical line and small arrow show the traveltime delay of the slow S wave on the east component (bottom) relative to the fast S wave on the north horizontal component (middle).

0 -4 -8

North (Cross Canyon) 0

2

8 velocity (µ/s)

1487

4 Time (s)

6

8

EQ1 Slow S-wave EQ2 slow S-wave

4 0 -4

East (Downstream)

-8 0

2

s( r , t ) C (υ r ,α , β )

4 Time (s)

6

σE β A( H (t − Tinitial ) − H (t − Theal )) µ

was used. It has a maximum error of less than 2% relative to the results of Dahlen (1974) and Richards (1976).

 t  t  t      0.7 exp  −  + 0.15 exp  − 0.4b  + 0.15 exp  − 0.8b   0.9a       − [ H (t − Tstop ) − H (t − Theal ) + H (t − Tinitial ) − H (t − Theal )]

(t − Tstop )   h 

(6)

where r is the distance on the fault from the hypocenter, ␣ is the compressional velocity, b is the shear velocity, l is the rigidity, rE is the effective stress, C(tr,␣, b) is a number determined from tr, the rupture velocity, and ␣/b (Dahlen, 1974; Richards; 1976), H is the Heaviside step function, t is time of rupture, Tinitial is the time of rupture initiation, Theal is the time the rupture begins to stop, and Tstop is the time rupture ceases. A is an amplitude factor calculated from the Kostrov slip-velocity function of Archuleta and Hartzell (1981) at dt after Tinitial that allows equations (4) to (6) to replicate the Kostrov slip-velocity function, and a, b, and c, are times from the onset of rupture until the slip velocity reaches, 0.5, 0.1, and 0.05, of the difference between the peak slip-velocity amplitudes and the steady slip-velocity amplitudes, respectively. Equation (6) was normalized over the entire fault to produce a maximum ratio of peak rupture slip velocity to peak healing slip velocity of 4:1. The approximation to C(tr,␣, b) of O’Connell and Ake (1995) C(υr ,α , β ) = 0.446437 + 0.707423υr − 0.151251

8

α β

(7)

Blind Thrust Rupture Models for the 19 April 1892 Mainshock Based on the seismic reflection data, geologic information, and seismicity depth extent, the Gordon Valley blind thrust was assigned a length of 17 km, a dip of 30⬚, and extends from a depth of 8 to 14 km. Its northern tip extends just north of 38.5⬚ N as indicated by the clear segmentation boundary in line S (Fig. 11) and the topography (Fig. 2). Its southern limit was inferred from the southern terminus of the English Hills (Fig. 1). The moment-fault-area relationship of Somerville et al. (1999) yields M 6.3 for this fault segment, but this relationship does not account for the style of faulting. The Wells and Coppersmith (1994) momentfault-area relationship for reverse faults was used to assign M 6.5 for the ground-motion simulations (LaForge, 1999). The fault tip was placed about 0.5 km west of the crest line of the English Hills (Fig. 23), consistent with interpretations of the seismic reflection data and deformation modeling of the topography. The hypocenter was placed near the southwest corner of the fault at 13 km to produce east-northeast forward directivity consistent with the northeast-oriented bias of the observed seismic intensities (Toppozada et al., 1981; Bennett, 1987) that suggest a nearly unilateral rupture of the fault. Two rupture simulations were performed, and the ground motions were averaged to smooth the short-wavelength amplitude fluctuations associated with the complexity

1488


South 16 Downdip distance (km)

of an individual earthquake rupture. Effective stresses ranged from 3 to 12 MPa (Fig. 18), producing peak slip velocities in the 40–370 cm/sec range (Fig. 19). Total slips were less than 2 m (Fig. 20) and have a ⬃1/k2 wavenumber spectrum that results from the ⬃1/k effective stress wavenumber spectrum and the variable rise time distribution (Fig. 21). The average rupture velocity is about 0.87*b but is allowed to vary from approximately ⬃0.7*b to ⬃1.05*b (Fig. 22), consistent with modeling results that show that supershear rupture velocities may occur along portions of fault rupture (Archuleta, 1984; Olsen et al., 1997; Hernandez et al., 1999; Belardinelli et al., 1999; O’Connell, 1999b). Three-component ground motions were produced for freesurface locations using a 1-km spacing on a 50 by 50-km2 grid using the simulation approach outlined subsequently. The PHV and PHA were calculated by rotating the horizontal components at 1⬚ increments through 180⬚ to find the maximum resolved peak horizontal ground motions.

20 22 24 26 28 0

5 10 Strike distance (km) 4

Effective stress (MPa) 6 8

15

10

Figure 18. Gray-shaded ⬃1/k effective stress distribution on the fault for the first of two rupture models used to simulate the 19 April 1892 M 6.5 earthquake on the Gordon Valley blind thrust fault. The perspective is normal to the fault plane and downdip distance is measured from the free surface. The hypocenter is the white star.

22 24 26

50

5 10 Strike distance (km)

15

100

150

200

250

300

350

Figure 19. Gray-shaded ⬃1/k peak slip-velocity distribution on the fault for the first of two rupture models used to simulate the 19 April 1892 M 6.5 earthquake on the Gordon Valley blind thrust fault. The perspective is normal to the fault plane and downdip distance is measured from the free surface. The hypocenter is the white star.

South 16 Downdip distance (km)

Downdip distance (km)

18

20

Peak slip velocity (cm/s)

The PHVs and PHAs produced by two rupture models were averaged and are contoured in Figures 23 and 24. The shallow refraction velocities derived from the 2D seismic reflection data that extend from the central Sacramento Valley to the English Hills show that shallow P-wave velocities are relatively low (1.6–1.8 km/sec) in the Sacramento Valley

North

18

28 0

Comparison of Synthetic PHV and PHA to Observed Modified Mercalli Intensities

South 16

North

North

18 20 22 24 26 28 0

5 10 Strike distance (km) 50

Total slip (cm) 100

15

150

Figure 20. Gray-shaded total slip distribution on the fault for the first of two rupture models used to simulate the 19 April 1892 M 6.5 earthquake on the Gordon Valley blind thrust fault. The perspective is normal to the fault plane and downdip distance is measured from the free surface. The hypocenter is the white star.



South 16

North

18 20 22 24 26 28 0

5 10 Strike distance (km) 0.5

Rise time (s) 1.0 1.5

15

2.0

Figure 21. Gray-shaded rise time on the fault for the first of two rupture models used to simulate the 19 April 1892 M 6.5 earthquake on the Gordon Valley blind thrust fault. The perspective is normal to the fault plane and downdip distance is measured from the free surface. The hypocenter is the white star.


South 16

North

18 20 22 24 26 28 0

5 10 Strike distance (km)

15

υ r /V s 0.70

0.75

Figure 22.

0.80

0.85

0.90

0.95

1.00

Gray-shaded rupture velocity-toshear-velocity ratio (tr /Vs) on the fault for the first of two rupture models used to simulate the 19 April 1892 M 6.5 earthquake on the Gordon Valley blind thrust fault. The perspective is normal to the fault plane, and downdip distance is measured from the free surface. The hypocenter is the white star.

1489

and increase substantially in the English Hills (3–4 km/sec). The Green’s function approach used here does not account for site-response differences across the region or for systematic 3D crustal velocity structure. The Green’s functions incorporated a rock site response. Field et al. (1998) showed that amplification differences between stiff soil and rock sites during the 1994 M 6.7 Northridge earthquake were about a factor of 1.5. Invoking different physical mechanisms, Field et al. (1998) (soil nonlinearity) and O’Connell (1999b) (direct S scattering) provided physical explanations of the diminished stiff soil amplifications during the mainshock relative to the aftershocks. Soil site-response characteristics in the Sacramento Valley are not well known. If soil responses are similar in the Sacramento Valley to those in the near-field region of Northridge, then the ground-motion responses calculated east of the English Hills could be somewhat higher than shown in Figures 23 and 24. Soil nonlinearity and shallow crustal low Q could substantially reduce the PHA in the Sacramento Valley, without substantially affecting the PHV. Consequently, the synthetic PHAs shown in Figure 23 may be overestimated for the Sacramento Valley portion of the map. Additionally, shallow P velocities decrease rapidly between the English Hills (⬃3 km/sec) and the Sacramento Valley (⬃2 km/sec). This strong lateral velocity gradient may produce a zone of basin edge amplification within several kilometers of the eastern margin of the English Hills and basin-edge-generated surface waves that could increase damage sustained in the Sacramento Valley. As Toppozada et al. (1981) caution, some intensity variations may be related to differences in construction and site responses. Also, the actual rupture velocity and stress drop characteristics of the 19, 21, and 30 April 1892 earthquakes are not known. Combining sources of uncertainty in source characterization and wave propagation inherent in this type of analysis, it is clear that the comparison and evaluation of the synthetic peak ground motions and intensity data in Figures 23 and 24 can only be semiquantitative. Wald et al. (1999) developed PHA- and PHV-Imm relations using data from six recent California earthquakes. Their results provide a basis for a semiquantitative comparison between synthetic peak ground motions and observed Imm, with the caveat that the absolute scaling of the synthetic peak ground motions is not well constrained. The contour outlining the Imm VIII-IX region correlates with the ⬎30 cm/ sec (Fig. 23) and ⬎300 cm/sec2 (Fig. 24) regions of the synthetic ground motions. The correlations between synthetic PHV and PHA and the Imm ⱖ VIII region in Figures 23 and 24 are consistent with the PHA and PHV regressions of Wald et al. (1999) against Imm for Imm ⱖ V using the recent California earthquakes shown in Table 2. The Imm ⱖ VIII region for the 19 April 1892 mainshock shown in Figures 1, 23, and 24 was constructed using the results of Toppozada et al. (1981) and Bennett (1987). Bennett (1987) provides detailed accounts of the intensity of ground shaking for the region between Vacaville and Winters. A consistent observation was that the intensities dimin-

1490


Longitude -122.300 40

-122.205

-122.109

20

-122.014

-121.918

-121.823

-121.727 38.76 800

Esparto

30

38.67

20

Winters

PV Allendale

38.49

400

Elevation (m)

10

MD

Latitude

BV

600

38.58

Davis

40


40

Dixon

VM 0

38.40

Vacaville

20

-10

38.31

-20

38.22 0

10


40

50

200

0

Figure 23. Color-shaded topography with the plan view of the Gordon Valley thrust fault (blue polygon), the position of the top of the fault for the 19 April 1892 M 6.5 rupture simulations (black dashed line), the postulated 19 April 1892 epicenter (black star), microseismicity epicenters (gray circles), 20-cm/sec interval contours of PHV (black curves, with the 100-cm/sec contour shown in red), the approximate position of the fissure described in Bennett (1987) (yellow line), and nearby cities as labeled (white squares). The position of the Monticello Dam is the black and white square labeled MD. PV, Pleasants Valley; BV, Berryessa Valley (now Lake Berreyessa); and VM, the crest of the Vaca Mountains. The pink thick-dashed contour encompasses the Imm ⱖ VIII region for the 19 April M 6.5 mainshock. ished abruptly west of the Vaca Mountains. A steep gradient of diminishing rock fall intensity was reported by S. B. Dunton as he cleared rock fall from east to west along the road located in Putah Canyon between the town of Winters (Fig. 1) and the present position of the Monticello Dam after the 19 and 21 April 1892 earthquakes (Bennett, 1987). Dunton was in the Putah Canyon when the 21 April earthquake occurred (Bennett, 1987). In particular, Dunton states (Bennett, 1987):

All of the great bowlders, which came crashing down the mountain sides, were in the lower canyon. There were not any rocks in the road at the “Devil’s Gate” [present site of Monticello Dam], or “middle canyon”, and none in the upper part of Putah Canyon. The earthquake was scarcely felt in Berryessa valley [currently Lake Berryessa impounded by Monticello Dam], hence no damage was done there. I have conversed with many citizens living there, and all tell the same story. Just

1491


-122.300 40

-122.205

Longitude -122.109 -122.014 -121.918

-121.823

-121.727 38.76 800

Esparto 30

38.67

00

2

Winters

PV

10

38.49

Allendale

Dixon

Latitude

0

MD

400

Elevation (m)

300 BV

600

38.58

Davis

30


200

20

VM 38.40

20

0

0

Vacaville

200

-10

38.31

-20

38.22 0

10


40

50

0

Figure 24. Equivalent of Figure 23 for PHA using 100-cm/sec2 contour intervals and with the 1000-cm/sec2 contour shown in red. See Figure 23 for details.

Table 2 Ranges of Peak Ground Motions for Modified Mercalli Intensities Imm 2

PHA (cm/sec ) PHV (cm/sec)

V

VI

VII

VII

IX

XⳭ

38–90 3.4–8.1

90–180 8.1–16

180–330 16–31

330–640 31–60

640–1200 60–116

1200 116

above the “Seeley Black Rocks” [these are located near the eastern entrance to Putah canyon about 4 km downstream of Monticello Dam] immense rocks, many in number, some of which weight from ten to fifty tons, came down from five hundred feet above into the bed of the creek. (p. 81) Geotechnical investigations in the vicinity of the Mon-

ticello Dam (Powell, 1999) reveal an abundant supply of perched rocks on both abutments above the dam; the lack of rock fall observed by Dunton in the vicinity of the presentday position of the Monticello Dam was not a consequence of insufficient source rock. Thus, Dunton’s observations suggest that peak ground motions diminished rapidly between the entrance to the Putah Canyon, 4.5-km east of the Monticello Dam, and the present position of the dam, during

1492 the 19 and 21 April earthquakes. Conversely, intensities increased very rapidly east of the entrance of the Putah Canyon in Pleasants Valley (Bennett, 1987). The maximum reported Imm values of IX (Bennett, 1987) extend from Pleasants Valley east to Winters and Allendale, and terminate just west of Dixon (Figs. 23 and 24). Synthetic PHVs and PHAs are consistent with these detailed intensity gradient observations; PHAs greater than 300 cm/sec2 and PHVs greater than 40 cm/sec that would consistently produce strong rock fall are confined to the eastern 1-2 km of the Putah Canyon, and both PHV and PHA rapidly increase eastward and terminate west of Dixon (Figs. 23 and 24). The 19 April earthquake Imm ⱖ V-VI contour has a large asymmetric lobe extending NNE across the northern Sacramento Valley (Bennett, 1987). Reports of Imm ⳱ VI extend 100–150 km north of the Monticello Dam to Willows and Chico (and 50 km north of the limits of the Imm ⱖ VVI contour shown in (Toppozada et al. [1981] and Bennett [1987]). Sokolov and Chernov (1998) showed that accelerations in the 3- to 6-Hz frequency range have the strongest correlation with intensities in the Imm V-VII range. The thickness of low-velocity, Upper Cretaceous sedimentary rocks thins from 10 to 12 km in the western Sacramento Valley (this article; also, Godfrey et al., 1997) to 4–5 km along the central and eastern portions of the Sacramento Valley (Godfrey et al., 1997). Although the Upper Cretaceous units in the top 8–10 km of the crust have low seismic velocities of ⬃4 km/sec, seismic attenuation in these units is relatively low. An M 3.5 earthquake in 1997 located 130 km north of the Monticello Dam in the Sacramento Valley produced S waves at the Monticello Dam that show a 6-Hz high-frequency S-wave acceleration response limit (Fig. 25). Invoking reciprocity, Figure 25 demonstrates that 3- to 6Hz accelerations associated with S waves can propagate to distances greater than 100 km from the English Hills to sites in the northern Sacramento Valley. Consequently, extension of the Imm ⱖ V-VI contour of the 19 April 1892 earthquake into the northern Sacramento Valley is consistent with NNE unilateral rupture of the Gordon Valley blind thrust fault combined with a crustal structure conducive to propagation of high-frequency S waves to distances greater than 100 km.

Aftershock Locations Relative to the Gordon Valley Blind Thrust Fault While the area enclosed within the Imm ⱖ VIII contour for the 21 April earthquake (890 km2) is only slightly smaller than for the 19 April earthquake (1100 km2), peak intensities were lower, with fewer Imm IX intensities reported for the 21 April earthquake than for the 19 April earthquake (Toppozada et al., 1981; Bennett, 1987). Toppozada et al. (1981) note that the estimated shaken area for the 21 April earthquake may be somewhat overstated due to the weakening of structures during the 19 April earthquake. The position of


the Imm ⱖ VIII contour is very similar for both events (Toppozada et al., 1981; Bennett, 1987). The centroid of the best-fitting magnitude contours for the 21 April 1892 earthquake from Bakun (1999) are about 15 km north of the corresponding contours for the 19 April 1892 mainshock. Report intensities are also consistent with a more northerly position for the 21 April earthquake as reported in the 22 April 1892 edition of the Morning Call published in San Francisco (from Bennett, 1987): It is quite clear that the focus of today’s [21 April] disturbance has been shifted to the northwest and has been located near Winters. Elmira and Vacaville got off lightly. The direction of the shocks has also perceptibly changed. (p. 77) One model for the 21 April earthquake is that it ruptured the northern 10 km of the Gordon Valley blind thrust segment (Fig. 26). As shown in Figure 11 it is possible that the Gordon Valley blind thrust fault bifurcates into two thrust sheets about 10 km south of its northern terminus. In this scenario shown in Figure 26 the 19 April earthquake ruptured most of the Gordon Valley segment length on the lower thrust sheet, and the 21 April earthquake ruptured the northern 10 km on the upper thrust sheet. Thus, the 21 April earthquake produced more localized strong ground shaking in the vicinity of Winters. Strike profile S (Fig. 26) shows a locus of intense strike-parallel folding and deformation south of the Gordon Valley–Trout Creek segmentation zone near the latitude (38.46⬚ N) where the strike of the fault-propagation fold changes to a more northerly strike. The existence of two thrust sheets along the northern 10 km of the Gordon Valley blind thrust segment can explain the abrupt change in strike of the fault-propagation fold that forms the crest of the English Hills, and the localization of the fold axis adjacent to the eastern slope of the Coast Range north of 38.46⬚ N (Fig. 26). The Wells and Coppersmith (1994) magnitudefault-area relation and the fault geometry shown in Figure 26 produces M 6.2, consistent with the estimated magnitude of the 21 April earthquake (Bakun, 1999). Several eyewitness accounts mention that ground shaking appeared to start in the hills to the west and propagate eastward toward the Sacramento Valley (Bennett, 1987). In particular, an account of the 21 April aftershock (Bennett, 1987) is consistent with updip rupture of the west-dipping Gordon Valley blind thrust fault as shown in Figure 26. . . . An eyewitness furnished a brief, but vivid description of the passage of this wave on Thursday (21 April) morning. He said that he was cultivating a field a little to the west of the Devilbiss residence (about one mile west of Winters), when he noticed that the cultivator, on which he was riding, plunged violently. At the same moment, there was a loud, roaring noise, and a cloud of dust sweeping rapidly along toward the town of Winters. (pp. 81–82)

1493

3000 2000 1000 0 -1000 -2000 -3000 -4000

10000

Vertical

Acc. spectra (counts)

Velocity (counts)


0 -2000 -4000 4000

Velocity (counts)


2000

East

2000 0 -2000 -4000 30

1000

100 10 10000

North


Velocity (counts)

4000

Vertical

35

40 Time (s)

45

50

S “ Fmax”

North 1000

100 10 10000

East

S “ Fmax”

1000

100 10 0.5

1

5 Frequency (Hz)

10

20

Figure 25. S-wave time windows from the Monticello Dam north abutment, threecomponent broadband velocity waveforms are shown on the left from the 11 November 1997, 13:00 UTC, ML 3.5 earthquake located at 39.674⬚ N, 122.146⬚ W at an elevation of ⳮ17 km. Corresponding acceleration spectra are plotted on the right. The earthquake was located ⬃130-km north of the dam between Willows and Orland. The S “Fmax” labels and arrows on the horizontal components show that S-wave attenuation is relatively low for frequencies up to about 6 Hz. Units are in counts since only the spectral shapes are of interest.

This observation suggests that rupture occurred approximately from west to east consistent with postulated updip rupture of the Gordon Valley thrust fault for both the 19 and 21 April earthquakes. The position of the 30 April aftershock beneath the Sacramento Valley is delimited by the clear translation of the western margin of the maximum Imm contours about 20 km east of the maximum Imm contours of the 19 and 21 April 1892 earthquakes (Fig. 26). The 30 April M ⬃5.5 aftershock is postulated to have occurred on the subhorizontal detachment located east of the primary Gordon Valley thrust shown in Figure 9 beneath the western Sacramento Valley (Fig. 26). A square fault with 5-km-long sides would be consistent with an M 5.5 earthquake (Wells and Coppersmith, 1994). The strike position of the 30 April event is quite uncertain, and the rupture plane could lie almost anywhere between

Dixon and Winters and reproduce the intensity distribution reported in Toppozada et al. (1981). A subhorizontal rupture plane with rupture initiating near the western edge of the rupture would reproduce the skewing of the Imm V–VI contour to the east, toward Sacramento, shown in Toppozada et al. (1981).

Discussion There are no known alternative fault sources that can be associated with the April 1892 Vacaville–Winters earthquake sequence. The Pittsburg–Kirby Hills strike-slip fault system (McCarthy et al., 1984; Unruh et al., 1997; WeberBand et al., 1998) is located 20 km to the south, too distant to produce the localization of maximum intensities within the English Hills and relatively lower intensities south of

1494


-122.300

Longitude -122.109 -122.014 -121.918

-122.205

-121.823

-121.727 38.76

40

800

Esparto 30

38.67

Davis

21 April

Winters

MD

il

10

30 Apr

PV

PR/DR

38.49

400

EH

Elevation (m)

WCF

600

38.58

Latitude


20

Dixon

Allendale 0

38.40

ril 19 Ap Vacaville -10

38.31

-20 0

10

GVF

20 CF

200

38.22

30 East grid distance (km)

40

50 Northern extend of the Pittsburg-Kirby Hills fault

0

Figure 26.

Color-shaded topography with the plan view of the two segments of the Gordon Valley thrust fault postulated to have ruptured on 19 April (black) with M 6.5 and on 21 April (blue) with M ⬃6.2. A plan view of the postulated location of the 30 April M ⬃5.5 aftershock is shown in magenta. The yellow line labeled PR/DR shows the approximate position of the axis of Plainfield Ridge and Dixon Rise (from Unruh and Moores, 1992). The rest of the annotation is as described in Figure 1. Dashed gray lines show other faults, including the northern terminus of the Pittsburg-Kirby Hills fault (based on seismicity), and the Green Valley (GVF), Cordelia (CF), and Wragg Canyon (WCF) strike-slip faults.

Vacaville (Fig. 25). The Green Valley, Cordelia, and Wragg Canyon strike-slip faults are located west of the crest of the Coast Ranges (Fig. 26). Consequently, these sources are not capable of producing both low intensities west of the crest of the Coast Ranges and high intensities in the English Hills and western Sacramento Valley. The correlation between synthetic PHV and PHA and observed Imm ⱖ VIII strongly supports the conclusion that the 19 and 21 April 1892 earth-

quakes were associated with rupture of the Gordon Valley blind thrust fault segment. The magnitudes postulated for these earthquakes are consistent with the intensity-moment misfit magnitude estimates of Bakun (1999) for these events. The isochrone approach used in this article to model ground motions is fairly simplistic. However, O’Connell (1999c) reproduced observed Northridge PHV distributions very well using a simpler isochrone approach. This suggests


that rupture directivity generally dominates local site-response variations for sites located within about one source dimension of the fault when evaluating PHV in the near field. Consequently, the association of the Gordon Valley blind thrust fault with the 19 and 21 April 1892 earthquakes is appropriate, particularly because none of the other adjacent fault sources are capable of reproducing the observed intensity distributions. Wong (1984) and Toppozada (1987) compared the relative areas of Imm ⱖ VI between the 19 April 1892 earthquake and the M 6.5 Coalinga earthquake. They concluded that these earthquakes probably had nearly comparable magnitudes, with the 19 April 1892 earthquake possibly having a 0.1-0.3 larger magnitude. In contrast to the across-fault unilateral rupture inferred for the 19 April 1892 earthquake, the hypocenter position and inferred fault plane position for the 1983 Coalinga earthquake (Eberhart-Phillips, 1989; Stein and Ekstro¨ m, 1992) are consistent with bilateral rupture. Bilateral rupture diminishes rupture directivity and produces a more symmetric pattern of intensity contours as observed for the Coalinga earthquake (Stover, 1983; Toppozada, 1987). As discussed previously the 19 April 1892 probably produced an expanded area of Imm ⱖ VI in the northeastern Sacramento Valley from a combination of rupture directivity with an optimally oriented crustal wave guide and relatively low crustal S-wave attenuation. Taking all these factors into account it is likely that the 19 April 1892 and 1983 Coalinga earthquakes both had magnitudes of about M 6.5. The agreement between the synthesized PHV and PHA and observed Imm ⱖ VIII contour for the 19 April 1892 earthquake strongly supports the thrust-faulting geometry inferred for the Gordon Valley blind thrust fault segment based on the 3D velocity structure, interpretation of the 2D seismic reflection data, and elastic deformation modeling. The results of Magistrale and Day (1999) suggest that multisegment rupture between the Gordon Valley and Trout Creek thrust segments is unlikely because the segments are vertically offset by several kilometers in a restraining step-over. In contrast, the lack of clear separation between the two thrust sheets that form the northern 10 km of the Gordon Valley thrust-faulting segment suggests that rupture of the Gordon Valley thrust segment could occur on the shallow northern branch combined with the deeper southern branch as shown in Figure 11.

Conclusions The analyses of 3D P-wave velocity structure, 135 km of 2D seismic reflection data, elastic deformation modeling, and synthetic ground-motion modeling of the April 1892 Vacaville–Winters earthquake sequence provide strong constraints on the locations and downdip geometries of the Gordon Valley and Trout Creek blind thrust fault segments. These data support a model of coseismic fault-propagation folding of the English Hills bordering the western Sacramento Valley in response to east-vergent blind thrust fault-

1495

ing beneath the eastern margin of the Coast Ranges. The positions of the easternmost anticlines in the English Hills, and the seismic reflection data were used to map out the position of the Gordon Valley and Trout Creek blind thrust faults and their east-vergent subhorizontal detachments beneath the Sacramento Valley. Elastic deformation modeling of the topography of the western Sacramento Valley–Coast Ranges boundary suggests that the Gordon Valley and Trout Creek blind thrust faults dip 30 and 20⬚, respectively, to the west, that the horizontal position of the tops of these faults are located with 1 km of the anticlinal fold crests, at depths of 8 and 9 km, respectively, and that the tops of the faults are horizontally offset along strike about 4–5 km. The two fold axes exhibit a right-stepping, en-echelon geometry. The downdip extension of these thrust faults intersect portions of the strike-slip Wragg Canyon fault west of the range front at the maximum depth of observed Coast Ranges seismicity (14 km). The Gordon Valley and Trout Creek blind thrust faults thus appear to be classic “tulip” thrust structures adjacent to transpressional strike-slip faults (Emmons, 1969; Naylor et al., 1986). Fault segmentation between the Gordon Valley and Trout Creek thrust segments is clearly evident in the seismic reflection data in the form of strike perpendicular folds formed by several kilometer offsets of the adjacent thrust sheets in depth and along strike. The lack of obvious tear faulting in the segmentation zones and large vertical offsets between thrust sheets, combined with the dynamic rupturing results of Magistrale and Day (1999), suggests that the Gordon Valley and Trout Creek thrust segments probably are not capable of multisegment rupture. Synthetic ground-motion modeling of intensity data from the M 6.5, 19 April 1892 Vacaville–Winters earthquake shows that this earthquake probably occurred on the Gordon Valley blind thrust fault. Synthetic PHVs and PHAs were produced using the inferred position of the Gordon Valley blind thrust fault, a rake of 105⬚, and a hypocenter near the southwest corner of the fault. The synthetic peak ground motions reproduce observed high-intensity and intensity gradient patterns over a by 50- by 50-km region containing the fault. An M 6.2 aftershock on 21 April 1892 probably occurred along the northern 10 km of the Gordon Valley blind thrust fault in the Gordon Valley–Trout Creek fault segmentation region near Winters. An M 5.5 aftershock that produced maximum Imm near Davis, probably occurred along the subhorizontal portion of the Gordon Valley blind thrust fault beneath the Sacramento Valley. The higher peak intensities and larger felt area of the 19 April 1892 earthquake relative to the M 6.5 1983 Coalinga earthquake are probably a result of greater rupture directivity associated with a nearly unilateral, updip rupture on a more steeply dipping fault in the 19 April 1892 earthquake, relative to the bilateral rupture of the M 6.5 1983 Coalinga earthquake.

Acknowledgments Support for this research was provided by the Dam Safety Program of the U.S. Bureau of Reclamation under project SPVGM. We thank Eld-

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ridge Moores, Ralph Archuleta, Roger Denlinger, Paul Spudich, Jon Ake, Joe Andrews, and Rob Graves for helpful discussions and suggestions; Arben Pitarka and John Wakabayashi for reviews and helpful comments, Paul Spudich for isochrone software; Rob Graves for 3D elastic finite difference software; Dave Copeland, Jon Ake, and Mike Ferrari for spud bar and shovel excavation, mixing and pouring concrete, and other feats of hard labor installing and removing temporary broadband seismic stations near Monticello Dam; Dave Copeland and Chris Wood for providing strong-motion instrumentation information; Chris Wood for software and hardware development. Reclamation’s Strong Motion Program shared equipment and phone lines. Several members of the Solano Irrigation District provided extensive assistance to the project; Don Burbey, Jay Shepard, and Mike Ferrari provided valuable information, site access, and instrument installation assistance. Valuable short-period earthquake data were obtained from the NCSN via the Northern California earthquake data center operated jointly by the University of California, Berkeley, California, Seismographic Station and the U.S. Geological Survey, Menlo Park, California. We thank David Oppenheimer for his extensive efforts in providing instrument response databases and other valuable information concerning the NCSN and Douglas Neuhauser for help in retrieving waveform data.

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