DEPARTMENT OF THE INTERIOR 

BRUCE BABBITT, Secretary 

U.S. GEOLOGICAL SURVEY 

Gordon P. Eaton, Director 

Any use of trade, product, or firm names in this publication 

is for descriptive purposes only and does not imply endorsement 

by the U.S. Government 

Manuscript approved for publication, January 2, 1997 

Library of Congress catalog-card No. 92-32287 

For sale by the 

U.S. Geological Survey 

Information Services 

Box 25286 

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Denver, CO 80225 
CONTENTS 

Aftershocks of the Lorna Prieta earthquake and their 

tectonic implications ................................................ 

By Lynn D. Dietz and William L. Ellsworth 

Response of regional seismicity to the static stress 

change produced by the Loma Prieta earthquake --------------- 

By Paul A. Reasenberg and Robert W. Simpson 

Spatial variations in stress from the first six weeks of 

aftershocks of the Loma Prieta earthquake ....................... 

By John W. Gephart 

Loma Prieta aftershock relocation with S-P travel 

times from a portable array ......................................... 

By Susan Y. Schwartz and Glenn D. Nelson 

Empirical Green's function study of Loma Prieta 

aftershocks: determination of stress drop ------------------------ 

By H. Guo, A. Lerner-Lam, W. Menke, and S.E. Hough 

U.S. Geological Survey aftershock ground-motion 

Response of U.S. Geological Survey creepmeters to 

the Loma Prieta earthquake ........................................ 

By K. S. Breckenridge and R.W. Simpson 

Increased surface creep rates on the San Andreas fault 

southeast of the Lorna Prieta main shock ........................ 

By Jeff Behr, Roger Bilham, Paul Bodin, Kate Breckenridge, and 

Arthur G. Sylvester 

Effect of the Loma Prieta earthquake on fault creep 

rates in the San Francisco Bay region ............................ 

By Jon S. Galehouse 

Postseismic strain following the Loma Prieta 

earthquake from repeated GPS measurements ------------------ 

By Roland Biirgmann, Paul Segall, Mike Lisowski, and 

Jerry L. Svarc 

Page 

Dl 
Models of postseismic deformation and stress transfer 

associated with the Loma Prieta earthquake ..................... 253 

By M.F. Linker and J. R. Rice 

A magnetotelluric survey of the Loma Prieta 

earthquake area: implications for earthquake 

processes and lower crustal conductivity ......................... 289 

By Randall L. Mackie, Theodore R. Madden, and Edward A. Nichols 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

CONTENTS 

INTRODUCTION 

By Paul A. Reasenberg, 

U.S. Geological Survey 

INTRODUCTION 

While the damaging effects of the earthquake represent 

a significant social setback and economic loss, the geophysical effects have produced a wealth of data that have 

provided important insights into the structure and mechanics of the San Andreas fault system. Generally, the period 

after a large earthquake is vitally important to monitor. 

During this part of the seismic cycle, the primary fault 

and the surrounding faults, rock bodies, and crustal fluids 

rapidlyreadjust in response to the earthquake's sudden 

movement. Geophysical measurements made at this time 

can provide unique information about fundamental properties of the fault zone, including its state of stress and the 

geometry and frictional/rheological properties of the faults 

within it. Because postseismic readjustments are rapid compared with corresponding changes occurring in the 

preseismic period, the amount and rate of information that 

is available during the postseismic period is relatively high. 

From a geophysical viewpoint, the occurrence of the Loma 

Prieta earthquake in a section of the San Andreas fault 

zone that is surrounded by multiple and extensive geophysical monitoring networks has produced nothing less 

than a scientific bonanza. 

The reports assembled in this chapter collectively examine available geophysical observations made before 

and after the earthquake and model the earthquake's principal postseismic effects. The chapter covers four broad 

categories of postseismic effect: (1) aftershocks; (2) 

postseismic fault movements; (3) postseismic surface deformation; and (4) changes in electrical conductivity and 

crustal fluids. 

SEISMOLOGICAL STUDIES 

The earthquake occurred inside the U.S. Geological 

Survey's Northern California Seismographic Network 

(NCSN) and is the largest earthquake within the network 

to be recorded by it. More than 75 high-gain seismographs 

recorded the main shock and nearly 11,000 aftershocks 

during the first 2 years of the earthquake sequence. Many 

of these instruments had been recording for 20 years before the earthquake. The approximately 5,600 earthquakes 

that they recorded in the region before the earthquake 

allow detailed comparisons of the pre- and post-earthquake activity. In addition, some 170 records were obtained from strong-motion instruments within 200 km of 

the epicenter. After the earthquake, 38 temporary seismographs (mostly strong motion instruments) were deployed 

to record the aftershock sequence. The first three papers 

in this chapter analyze the seismological data and focustheir analyses on the structure of the San Andreas fault, 

its interaction with the major nearby faults, and the contemporary tectonics of the greater Pacific-North American plate boundary in the San Francisco Bay area. 

Dietz and Ellsworth provide a comprehensive survey of 

the aftershock locations and magnitudes, the focal mechanisms for the main shock and the larger aftershocks, and 

the time-evolution of the aftershock sequence. From these 

observations they infer possible geometries of the faulting 

structures active in the earthquake and their possible kinematic and geometric relationship to the San Andreas 

fault. The aftershock focal mechanisms were highly diverse in orientation and unlike those of the pre-earthquake 

seismicity, leading Dietz and Ellsworth to conclude that 

the stress drop in the main shock may have been nearly 

complete. 

At greater distances from the earthquake, Reasenberg 

and Simpson examine the post-earthquake seismicity 

changes in central California. They compare the 20 years 

of seismicity recorded in central California before the 

earthquake to the activity during the 20-month postseismic 

period. They find that some regions (for example, the San 

Francisco peninsula) sustained an increase in seismic activity after the earthquake, while at least one other area 

(the Hayward fault) sustained a decrease in activity. These 
D2 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

changes are consistent with the stress changes calculated 

for simple elastic dislocation models of the earthquake 

when the effective coefficient of friction on the faults is 

assumed to be low (less than 0.3). A surprisingly small 

amount of stress change (0.1 bar) was apparently sufficient to affect the seismicity on these nearby faults. 

Another view of the stress changes produced by the 

earthquake is provided by the aftershock focal mechanisms. Gephart uses fault plane solutions for the first 6 

weeks of the aftershock sequence to map the spatial variations in stress in the vicinity of the main shock rupture. 

Near and northwest of the main shock hypocenter, he finds 

low shear stress remaining on the main shock fault plane 

after the earthquake, again consistent with the inference 

of a nearly complete stress drop in the earthquake. Southeast of the hypocenter, near Pajaro Gap and the creeping 

segment of the fault, significant right lateral shear apparently remained on the fault. 

Aftershock observations also were used to hone seismological technique. Schwartz and Nelson use synthetic 

calculations and data from portable (PASSCAL) threecomponent seismographs in order to compare earthquake 

hypocenter location methods based on various combinations of P, S, and S-P arrival times. They find that the 

unsynchronized S-P arrival times observed with a sparse, 

portable array provided remarkably good locations relative to those obtained from the more extensive, permanent 

NCSN array. The PASSCAL data also were used by Guo 

and others to estimate source characteristics of the 

aftershocks. Using an empirical Green's function technique, they estimate Brune-model stress drops from the 

corner frequencies and moments of the aftershocks. They 

report an apparent dependence of stress drop on earthquake moment. 

Other portable seismic instruments, in addition to the 

PASSCAL instruments, were deployed after the earthquake 

to record aftershocks. Altogether, 38 digital seismographs 

were deployed at 195 sites during the 3-month period 

after the earthquake. Wennerberg documents these deployments and the massive and disparate data sets they produced. He provides us with detailed maps, instrument 

descriptions, and data logs for the aftershock deployments. 

He also provides a table of references to studies utilizing 

these data and published elsewhere. In so doing, he assures that future seismological investigations of the earthquake will be able to proceed efficiently with a map of 

these essential, recorded seismological data sets. 

POSTSEISMIC MOVEMENT ON THE 

SAN ANDREAS AND NEARBY FAULTS 

At the time of the earthquake, 27 creepmeters were 

recording surface movements at points on the San Andreas and Calaveras faults within about 200 krn of the 

epicenter. These instruments had been recording fault 

movement for periods up to 25 years before the earthquake. In addition, small-scale (50-200 m) triangulation 

networks at about two dozen sites in the San Francisco 

Bay Area on the San Andreas, Calaveras, Hayward, and 

Concord-Green Valley faults had been observed regularly 

to measure surface fault displacements for up to 13 years 

before the earthquake. These observations form the basis 

for three papers describing and modeling the surface movements of the major faults in the Bay Area after the earthquake. 

In a detailed study of surface fault displacements (creep) 

at the eight nearest creepmeter sites on the San Andreas 

and Calaveras faults, Breckenridge and Simpson find that 

both the sense (right-lateral or left-lateral) and magnitude 

of long-term creep rate changes on nearby faults agree 

with coseismic changes in horizontal shear stress calculated for elastic dislocation models of the earthquake. In 

contrast, the amplitudes of coseismic steps in fault displacement measured at the same sites apparently bear little 

similarity to the calculated static stress changes; they propose that dynamic stresses (shaking) likely account for 

these coseismic steps. The correlation between the regional 

creep rate changes and the earthquake-induced static stress 

changes parallels changes observed in the regional seismicity rates (Reasenberg and Simpson) and suggests that 

over a period of a few years a dominant effect of the 

earthquake on the major Bay area faults was the sympathetic advancement and retardation of ongoing processes 

of seismicity and creep, driven by the static stress changes. 

Behr and others also investigate the increases in creep 

rate at the five USGS creepmeter sites on the San Andreas fault studied by Breckenridge and Simpson, and at 

three additional sites between San Juan Bautista and Pajaro 

Gap. In contrast to the ephemerally enhanced fault creep 

observed after some earthquakes, they find a sustained 

level of elevated creep rate during the first 3 years after 

the Loma Prieta earthquake. Their two-dimensional elastic boundary models of the San Andreas and Calaveras 

faults can qualitatively explain the observed decrease in 

postseismic slip with distance from the earthquake. 

Triangulation measurements over small (50-200 m) aperture arrays spanning Bay area faults tell a similar story. 

Galehouse reports that the creep rate on the San Francisco 

peninsula section of the San Andreas fault was nil both 

before and after the earthquake, while creep rates approximately doubled on the San Andreas fault southeast of the 

rupture. At the same time, creep on the Hayward and 

southern Calaveras faults decreased and, at some stations 

along the southern Hayward fault, creep may have stopped 

or become left lateral. Galehouse's findings are consistent 

with the seismicity and other creep observations presented 

in this chapter and with static stress change calculations 

(Breckenridge and Simpson), suggesting that the East Bay 

faults were at least temporarily relaxed by the earthquake. 
INTRODUCTION D3 

How long this relaxation will last before being overtaken 

by continuing tectonic and additional seismogenic loading will surely be the subject of continuing measurements 

and modeling efforts. 

POSTSEISMIC DEFORMATION OF THE 

CRUST 

Prior to the earthquake, horizontal deformation in the 

epicentral region had been monitored for about 20 years 

by land-based trilateration measurements, and for several 

years by Global Positioning System (GPS) measurements. 

Thus, from a geodetic monitoring point of view, the 

earthquake's location was indeed advantageous. Resurveying of these sites and the rapid and repeated survey of 

new sites by several campaigns launched after the earthquake has provided an extensive and varied data set for 

constraining the surface deformation and inferring the deformation at depth associated with the earthquake. The 

data and conclusions of these studies compliment the seismological and fault slip studies; together they provide one 

of the most extensive sets of geophysical observations of 

the crustal response to a large earthquake. 

Burgmann and others investigate the deformation of the 

crust after the earthquake using data from repeated GPS 

measurements in two fault-crossing networks and compare it to the deformation inferred from preseismic observations. They find that horizontal displacement rates 

northwest of the epicenter did not change significantly at 

the time of the earthquake, while the fault-parallel station 

velocities in the epicentral region increased significantly 

over pre-earthquake rates. They also find evidence for 

San Andreas fault-normal shortening near the epicenter 

and within the Foothills thrust. Using linear, elastic dislocation models, they find that a two-plane slip model with 

maximum depth of about 16 km best explains these data. 

They suggest that the inferred reverse slip on the Foothills thrust fault appears to have been triggered by stress 

changes produced by the earthquake. 

At the northwest end of the main shock rupture, just 

northwest of the study region of Biirgmann and others, a 

10-km aperture trilateration network was repeatedly measured by Langbein in the 2-year period after the earthquake. Although no preseismic observations of these lines 

had been made, comparisons are made to preseismic measurements of the Black Mountain geodetic net 15 krn to 

the northwest. Langbein's analysis of these data finds that 

the rapid temporal decay in the postseismic deformation 

rate in the network is typical for a postseismic geodetic 

response. It fails to find evidence for the fault-normal 

horizontal contraction inferred by Biirgmann and others 

from the regional GPS data. 

Linker and Rice use the geodetic observations to 

explore viscoelastic, time-dependent models of the 

postseismic deformation. Their models contrast with those 

of Burgmann and others in that they allow for transient, 

postseismic slip in the deep, aseismic region of the fault 

zone, well below the maximum depth of slip in the elastic 

dislocation models. While the deep, viscoelastic models 

are compatible with most of the strike-slip deformation 

observed in the epicentral region, they do not explain the 

fault-normal contraction observed there and may 

overpredict strike-slip deformation near Black Mountain 

to the northwest of the earthquake. The paper explores a 

wide range of alternative assumptions about fault zone 

rheology and slip distribution in the crust. 

Near the southeast end of the main shock rupture, a 

borehole strainmeter recorded the tensor strain during and 

after the earthquake. This instrument had been recording 

changes in tensor strain for 10 years before the earthquake. Gladwin and others compare the pre- and postearthquake strain rates at this site and find an apparent 

increase in fault-parallel shear strain rate. The timing of 

this strain anomaly relative to the Loma Prieta earthquake 

and its regional extent are addressed in detail through 

comparisons with surface creep and geodetic measurements. The occurrence of the Chittenden earthquake sequence in April 1990, located near the southeast end of 

the main shock rupture and only 10 krn from the borehole, apparently increased the shear strain rate at this site. 

To explain these strain observations, Gladwin and others 

propose that a shallow asperity southeast of the main shock 

rupture initially resisted the increased shear load on the 

San Andreas fault, and then yielded in the April 1990 

sequence, thereby unpinning the San Andreas fault and 

resulting in even higher shear strain rates. 

CRUSTAL FLUIDS AND ELECTRICAL 

CONDUCTIVITY 

Finally, Mackie and others present the results of a 

magnetotelluric survey of the epicentral area. Working 

with observations made one year after the earthquake, 

they find evidence for a vertical tabular zone of anomalously high electrical conductance in the lower crust below the earthquake. Unfortunately, no corresponding 

measurements were made before the earthquake, so their 

interpretation of these data cannot conclusively associate 

the anomaly with the earthquake; it might instead be a 

permanent feature of the San Andreas fault. However, 

they suggest an intriguing scenario in which fluids (water) that were already present in the crust experienced 

increased flow due to an increase in porosity and fluid 

connectivity in the fault zone caused by fracturing during 

the earthquake. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND 

THEIR TECTONIC IMPLICATIONS 

CONTENTS 

By Lynn D. Dietz and William L. Ellsworth, 

U.S. Geological Survey 

Formation ........................................................... 

Decay ............................................................... 

Structure of the aftershock zone ...................................... 

Main zone ........................................................... 

San Andreas fault ................................................ 

Castro and southern Sargent faults ----------------------------: 

Faults north of the main zone and east of the San Andreas 

Hanging wall activity - ............................................ 

Relationship of the preshocks to the main shock and 

aftershocks - ...................................................... 

Discussion ............................................................... 

The earthquake and the San Andreas fault ------------------- Acknowledgements - ..................................................... 

References cited ......................................................... 

Page 

D5 

6 

6 

6 

7 

10 

12 

12 

14 

14 

14 

15 

15 

15 

19 

2 1 

23 

37 

37 

3 9 

39 

40 

41 

41 

42 

43 

ABSTRACT 

The Loma Prieta earthquake and its aftershocks ruptured faults within a broad zone along the San Andreas 

fault in the southern Santa Cruz Mountains. The main 

body of aftershocks forms a southwest-dipping zone, rising from the main shock hypocenter at 15.9 km depth to 

the San Andreas fault, and extending bilaterally along 

strike for a distance of 42 km. This zone, however, does 

not mark a simple fault plane, as it has a width of 1 to 2 

km. Aftershocks also occur on faults well removed from 

the main zone on both sides of the San Andreas fault. The 

aftershock zone assumed essentially its final form within 

less than 1 day and only showed apparent growth to the 

southeast, where aftershock activity occurred on the normally active part of the San Andreas and Castro faults. 

Aside from these two faults, the aftershocks occurred on 

structures that were essentially aseismic in the 20 years 

preceding the earthquake. Focal mechanisms of the aftershocks bear little resemblance to the main shock or preceding seismicity. The mechanisms tend to have 

fault-normal P-axes, although a very wide range of mechanism types is observed. The diversity and orientation of 

the aftershock mechanisms and the lack of spatial correlation with the preshocks argue for a major reorientation of 

the stress field by the main shock and are consistent with 

a complete stress drop on the main shock fault. 

The main shock nucleated with a foreshock 1.6 s before 

the start of large-amplitude high-frequency seismic 

radiation. Although the foreshock was too feeble to trigger many of the strong-motion stations in the epicentral 

region (vertical acceleration~.Ol g), it had an amplitude 

magnitude of 5. Nearfield displacement seismograms 

indicate significantly larger moment release at long 

periods, equivalent to MW=5'/2. The foreshock thus appears to mark the beginning of a continuous and relatively smooth process of moment release that only began 

to propagate away from the initial hypocenter after 1.6 s. 

The aftershock sequence has no distinct beginning, with 

the earliest locatable events emerging from the main 

shock coda. 

The Loma Prieta fault, as mapped by the main zone of 

aftershocks, exists as a structure distinct from the San 

Andreas fault along most of its length. Along its southern 

end, where the sense of displacement changes from reverse to strike slip, it merges seamlessly with the San 

Andreas fault near Pajaro Gap. We speculate that the fault 

represents a former part of the San Andreas that has been 

abandoned as a strike slip fault and has evolved into a 

right-reverse fault to accommodate motion around the compressional bend in the San Andreas fault through the southern Santa Cruz Mountains. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

INTRODUCTION 

The Loma Prieta earthquake presents an unparalleled 

opportunity to study the source process of a major earthquake using multiple and complementary data provided 

by geology, geodesy, and seismology. Our objectives in 

this paper are to describe the spatial distribution and temporal characteristics of the sequence, the connection between the earthquake and the San Andreas fault, and the 

relationship between the preshocks, the main shock, and 

the aftershocks from the vantage point of short-period network seismology. 

The Loma Prieta earthquake occurred within a portion 

of the San Andreas fault system in one of the most densely 

instrumented regions for the observation of microearthquake activity. Systematic cataloging of accurately 

located hypocenters, magnitude 21.5, as recorded by the 

U.S. Geological Survey's telemetered seismic network 

(CALNET), began in this area over two decades before 

the 1989 earthquake. The CALNET thus provides a 

wealth of information on not only the complete aftershock sequence of a major event, but also on its relationship to background seismicity over an extended period of 

time. 

This study builds on the work of many of our colleagues who have analyzed various aspects of the CALNET 

data to determine the crustal structure, hypocentral locations, focal mechanisms, and state of stress as reported in 

this chapter and other chapters in this volume [also 

Eberhart-Phillips and others (1990), Michael and others 

(1990), Foxall and others (1993), Olson (1990), 

Oppenheimer (1990), Schwartz and others (1990), and 

Zoback and Beroza (1993)l. In this paper we extend the 

results of our earlier paper (Dietz and Ellsworth, 1990) to 

present a comprehensive picture of the earthquake and 

attendant seismicity through the first two years of the 

aftershock period, and we compare the patterns of hypocenters and focal mechanisms with those observed during 

the preceding 20 years using uniform analysis methods. 

DATA AND METHODS 

The primary data sets examined in this study are the 

locations, magnitudes, and focal mechanisms of earthquakes that occurred in the vicinity of the Loma Prieta 

earthquake. These data are derived from the analysis of 

seismograms recorded by CALNET, and are based on the 

measured arrival times and polarities of P-wave onsets 

and the duration of coda waves, recorded as the time from 

the P-wave onset to the time when the average absolute 

amplitude (2-second window) falls below a threshold value 

of 60 millivolts (Baton, 1992). 

Earthquakes within the dashed box in figure 1 from 

January 1, 1969, through October 17, 1991, are considered in this study. About 5,600 events were recorded in 

this region before the Loma Prieta earthquake and nearly 

11,000 were processed in the first 2 years of the sequence. 

The region completely encompasses the primary aftershock 

sequence and is approximately twice as long and twice as 

wide as the extent of main shock rupture determined by 

geodetic and seismic means. 

HYPOCENTER ANALYSIS PROCEDURES 

Earthquake magnitude, denoted Md, is determined using the formulation of Eaton (1992). Supplementary magnitudes, principally Mi, are used for the largest events. 

During the 1970's, the CALNET's detection capability 

increased to lower magnitudes as the network was enlarged and densified. Also, a change in network amplification in 1977 is thought to have introduced an artificial 

shift in catalog magnitudes (Marks and Lester, 1980; 

Habermann and Craig, 1988). Since 1980 the detection 

threshold in the study area has remained relatively constant and no network change affecting magnitudes is known 

to have occurred. Figure 2 shows the frequency-magnitude 

density distribution, binned by 0.1 magnitude units, of the 

Figure 1.-Map showing study region (dashed line) and locations of 

CALNET seismic stations (triangles) within -60 km of the Loma Prieta 

main shock (star). Solid triangles are assigned to velocity model 1; open 

triangles to model 2 (see table 1). Shaded area represents the extent of 

the aftershock zone. Labeled faults are the San Andreas, SAF; Sargent, 

S; Calaveras, CAL; Hayward, HAY; and San Gregorio, SG. 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D7 

pre-Loma Prieta earthquake catalog. Due to the network 

changes mentioned above, we make no estimate of catalog completeness or b-value for the early time range (fig. 

2A). From 1980 through the end of the study, the catalog 

appears to be complete to magnitude 1.0 (fig. 25). It should 

be understood, however, that a small percentage of the 

earthquakes with magnitudes above these thresholds are 

missing because their seismograms are obscured by larger 

events, such as in the coda of the main shock. The b-value 

of the preshocks is 0.64k0.02 calculated from M2l .O events 

(0.74k0.04 from M22.0 events). The aftershock b-value, 

omitting the incomplete record of the first day's activity, 

is 0.83k0.01 calculated from M21.0 events (0.74k0.03 from 

M22.0 events). 

Event locations were derived from a merged set of handand machine-picked P-wave arrivals at CALNET stations 

(fig. 1) using the location program HYPOINVERSE 

(Klein, 1989), with the station corrections and P-wave 

velocity model described below. In locating events, the 

data are weighted as the product of a function of their 

quality (inversely proportional to the variance of the pick 

uncertainty), and a function of the epicentral distance of 

the station. The distance weight is unity out to an epicentral distance of 52.5 km, and decreases to zero at a distance of 105 km. 

When examining the spatial distribution of hypocenters, 

we restrict our attention to the most accurate locations 

(fig. 3). The selected hypocenters have root mean square 

residual (rms) <0.15 s, horizontal standard error <1.0 krn, 

vertical standard error <2.0 km, number of stations 28, 

and magnitude 21.0. The 3,085 selected preshocks include 

78 percent of all Md21.0 and 82 percent of all Md22.0 

events. The 4,998 selected aftershocks include 79 percent 

of the total number of events with Md21.0 and 86 percent 

of those Md22.0. The mean rms is 0.07 s. 

CRUSTAL MODEL AND LOCATION ACCURACY 

In developing a traveltime model for locating the events, 

our objective was to find a simple model that accounted 

for the pronounced near-surface variations in velocity 

known to exist in the region (Wesson and others, 1973a) 

without introducing biases in the overall distribution due 

to incorrect modeling of the true three-dimensional earth 

structure. In the main shock, for example, traveltimes depart from a laterally homogeneous model by up to 2 s 

(fig. 4). Fortunately, the fact that traveltime residuals at 

individual stations change slowly as a function of event 

location within the sequence implies that near-station 

structure and path-averaged delays account for most of 

the observed variance in traveltime. Accordingly, we 

model the traveltimes using a one-dimensional velocity 

model with station corrections. Obviously, locations made 

with these assumptions will contain biases due to 

unmodeled three-dimensional effects, but these appear to 

be both predictable and of minor importance in light of 

similar locations from the three-dimensional studies of 

Michael and Eberhart-Phillips (1991), and Foxall and 

others (1993). 

Using P-wave traveltimes from the main shock and 107 

aftershocks distributed evenly throughout the aftershock 

zone, we calculated station traveltime corrections and 

one-dimensional velocity models with the joint 

hypocenter-velocity inversion program VELEST (Roecker, 

1981; Kradolfer, 1989). To account for the obvious differences in crustal geology across the San Andreas fault, 

we partitioned the stations into two sets corresponding to 

the northeast (Franciscan basement) and southwest 

(Salinian basement) sides of the San Andreas fault (fig. 

1). Then we simultaneously derived station corrections 

and a separate velocity model (seven layers over halfspace) 

0123456701234567 

MAGNITUDE MAGNITUDE 

Figure 2.-Pre-Loma Prieta seismicity catalog. Number of earthquakes in each 0.1 magnitude-unit bin (A) from 

1969 through 1979 and (B) from 1980 through October 17, 1989. 
D8 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

for each side of the fault. Following the recipe to calcu- are poorly resolved, principally due to inadequate samlate a minimum one-dimensional velocity model (Kissling pling. As a consequence, a tradeoff exists between focal 

and others, 1994), we initially considered several candi- depth and velocity for the deepest events in the sequence. 

date models; our preferred velocity model (table 1) is used In particular, the depth of the main shock could be in 

to locate all events in this study. Velocities below 18 km error by about 1 km. 

Figure 3.-Best located events within the study region (dashed line). (A) January 1, 1969, through October 17, 1989, and (B) October 18, 1989, 

through October 17, 1991. Symbols vary by magnitude: star = M27.0, circle = kP5.0, diamond = M24.0, triangle = M23.0, and + = M21.0. Dots 

labeled Z and Z' are depth section endpoints. 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D9 

When considering the location results, two types of er- fidence ellipse, of the hypocenter location and averages 

rors need to be distinguished: relative location error, which k0.3 km in epicenter and k0.6 km in depth for our catameasures the precision of the location in the reference log. This means that internal structures with dimensions 

frame of other locations; and absolute location error. The of about 1 km may be reliably interpreted in most cases. 

relative error is estimated from the standard error, or con- Figure 5 illustrates the distribution of standard error for 

Figure 3 .-Continued. 
Dl0 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

two portions of the aftershock zone, shown in transverse 

cross section. In figure 5A, the central portion of the aftershock zone has an apparent width of about 2 km, which 

is supported by the confidence ellipsoids. In contrast, the 

San Andreas fault at the southern end of the zone may 

have zero width (fig. 5B). 

Absolute location errors are more difficult to assess. It 

is well known (for example, Reasenberg and Ellsworth, 

1982) that many velocity models may be constructed which 

produce equivalent patterns of relative event locations that 

differ in absolute location by up to several kilometers. We 

found this to be true of the Loma Prieta data in deriving 

our final velocity model. Using the method to calculate a 

minimum one-dimensional velocity model (Kissling and 

others, 1994), we constructed several models that fit the 

data equally well and have the same pattern of relative 

locations but have significantly different absolute locations. The absolute locations vary in a northeast-southwest 

direction, normal to the fault and the geologic fabric of 

the region. To choose among candidate models, we selected the one that best reproduces the locations of surface explosions in the center of the aftershock zone (fig. 

6). This model relocates the explosion nearest the main 

shock epicenter to within 0.3 km, and the nearby shots to 

within 1 km or better (table 2). 

A dense network of portable instruments deployed 

shortly after the main shock in the center of the aftershock zone recorded several events with sufficient spatial 

resolution to permit a direct contouring of their traveltimes. 

The event shown in figure 7 exhibits clear evidence for 

slower velocities southwest of the San Andreas fault as 

demonstrated by the more closely spaced isochrons to the 

southwest of the fault and by the suggestion of a refraction angle at the fault. Locations produced by the 

three-dimensional model of Michael and Eberhart-Phillips 

(1991) and by our model agree well with each other, but 

both are systematically displaced -1 km southwest from 

the apparent center of the isochrons. We expect the isochron center to be displaced to the northeast of the true 

epicenter when higher velocities exist to the northwest of 

the fault, as appears to be the case here. Thus, this qualitative evidence supports the absolute accuracy (-0.5 km) 

of both our location and the three-dimensional location. 

The performance of the model at the southeastern end 

of the area is substantially poorer. The southernmost shot, 

in the San Andreas fault zone near San Juan Bautista, 

locates almost 3 krn west of its actual position (fig. 6, 

table 2). Earthquake hypocenters in this vicinity also locate about 3 krn west of the San Andreas surface trace. 

The shot's mislocation suggests that our model does not 

account for velocity variations in the southernmost portion of the aftershock zone. This segment of the fault is 

known to have a very strong velocity contrast across it 

(Wesson and others, 1973a; Michael and Eberhart-Phillips, 

1991), with very low velocity material located to the northeast of the fault. Traveltime contours for the main shock 

(fig. 8) show a refraction angle in the isochrons along the 

fault at San Juan Bautista, strong evidence for low velocities to the northeast of the San Andreas in this region. 

This velocity contrast is of the opposite sense than the 

smaller contrast observed along the main body of the aftershock zone (Lees, 1990; Mooney and Colburn, 1985; 

fig. 7). We infer that the seismicity near San Juan Bautista 

occurs on the San Andreas fault and that the westward 

mislocation is an artifact of improperly modeled velocity 

perturbations in this region. 

FOCAL MECHANISM DETERMINATION 

Focal mechanisms were calculated for the best-recorded 

aftershocks from P-wave first motions using FPFIT 

(Reasenberg and Oppenheimer, l985), a grid-search algorithm that finds acceptable double-couple mechanisms. A 

minimum of 15 hand-picked first motions and a misfit 

value S0.15 were required. For the mechanisms studied 

Epicentral Distance (km) 

Figure 4.-Main shock traveltimes to CALNET seismic stations versus epicentral distance. Times are reduced by 6 kmls. 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS Dl 1 

Table 1 .-Seismic velocity models (P-wave) 

Depth to Top Model 1 Model 2 

of Layer SW of San Andreas NE of San Andreas 

krn km/s km/s 

here, the average uncertainties in strike, dip, and rake are 

16O, 21Â° and 30Â° from an average of 57 first motions per 

event. In addition to determining the most probable mechanism for a particular earthquake, we also use FPFIT to 

test the hypothesis that a particular mechanism or class of 

mechanisms will fit the first motion data of an event within 

the confidence limits of its unconstrained (most probable) 

solution. For example, we can search for all earthquakes 

that have mechanisms that resemble the main shock. 

In almost all parts of the study area, we find that the 

FPFIT focal mechanisms for our HYPOINVERSElocated events explain the first-motion data as accurately 

as mechanisms of the same events located by 

three-dimensional methods. However, minor first-motion 

discrepancies exist for most mechanisms, whether computed using one-dimensional or three-dimensional models. The main shock, for example, has discrepant motions 

at a number of stations (fig. 9). The three-dimensional 

models of Michael and Eberhart-Phillips (199 1) and Foxall 

and others (1993) do not account for these discrepancies. 

Discrepant first-motions at stations to the northwest of 

the epicenter cannot be explained by a lateral refraction, 

which would be apparent in the main shock isochrons of 

figure 8. These discrepancies suggest that significant 

lateral variations in structure may be present near the 

hypocenter. 

The one area where unmodeled effects of lateral structure appear to significantly bias the focal mechanisms is 

the southeastern part of the study region. For events on 

the San Andreas fault near San Juan Bautista, the velocity 

contrast across the San Andreas fault is so strong that first 

P-arrivals to all stations on the northeast side undergo 

significant lateral refraction (Spieth, 198 1). Because this 

refraction is not accounted for in the one-dimensional 

velocity model, the calculated take-off angles are incorrect, and the resulting focal mechanisms for these San 

Andreas events have an apparent shallow dip on their 

northwest-trending fault plane, whereas the northeasttrending auxiliary plane is vertical and strikes normal to 

the fault. We believe that these events actually have nearly 

vertical nodal planes and represent horizontal strike slip 

movement along the San Andreas fault proper. The apparent dip of the San Andreas fault as defined by the hypo10 0 

DISTANCE (KM) 

Figure 5.-HYPOINVERSE error ellipse axes of well-located MdX?.O aftershocks projected in cross section for (A) the 

region just north of the main shock (section B-B'of figure 21A) and (B) the creeping section of the San Andreas fault 

(section G-G' of figure 21A). Solid inverted triangle marks the surface trace of the San Andreas fault; open triangle 

marks the Sargent fault. 
Dl2 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

centers (fig. 5B) may also be an artifact of the unmodeled Note that r scales as ACT"'". Because a change in Ao of 

strong velocity contrast across the fault. a factor of 3 translates into a change in r of only a factor 

of 0.7, the plotted symbols only weakly depend on the 

assumed value of Ao. The analysis of small subsets of 

DISPLAY OF EARTHQUAKES aftershocks by Hough and others (1991) and Fletcher and 

Boatwright (1991) suggest that the mean Ao value for 

Many cross sections of hypocenters in this paper por- the sequence may be greater than our assumed value by 

tray each earthquake as a circular area with its dimension as much as a factor of 3. Thus our rupture areas may 

scaled to correspond to the rupture area of a constant overestimate the area involved in faulting up to a factor 

stress drop circular crack. By displaying earthquakes as of 2. 

rupture areas, cross sections in the plane of the earth- Earthquake focal mechanisms appear in backquake display the approximate faulting area of the events. hemisphere projection on cross sections and as 

For our figures, we assume a 30 bar (3 MPa) stress drop lower-hemisphere projections in map view, with compres- 

(Ao) and use Md to determine seismic moment (My), and sional quadrants shaded. Strain axes for the equivalent 

thereby radius (r), using the formulas double couple appear as correctly proportioned horizontal 

projections of the three-dimensional axes. 

log#o = 1.5Md + 16.0 

and OVERVIEW OF SEQUENCE 

Figure 6.-Map of study region shows the main shock (star), actual shot 

locations (numbered squares, listed by shot number in table 2) and shot 

locations using our velocity model (x). Dots show landmarks: SC, Stevens 

Creek reservoir dam; LR, Lake Ranch reservoir; LX, Lexington dam; 

EL, Lake Elsman; LP, Loma Prieta peak; GF, Grizzly Flat; HZ, Hazel1 

Dell and Mt Madonna Rd. intersection; PG, Pajaro Gap; SJ, San Juan 

Bautista. The study area is divided into 7 subregions: SCV, Santa Clara 

Valley; LPN, Loma Prieta North; LPC, Loma Prieta Central; LPS, Loma 

Prieta South; CSA, Creeping San Andreas; SAR, Sargent and Castro 

faults; MON, Monterey Bay area. 

Hypocenters of the main shock and its aftershocks activated a 45- to 60-km-long and 10- to 20-km-wide volume 

of the crust along the San Andreas fault in the southern 

Santa Cruz Mountains. The events occur over the entire 

depth range of the seismogenic crust. The main body of 

aftershocks defines a tabular, southwest-dipping zone, with 

the main shock near the base and center of the distribution (fig. 10). Aftershocks deeper than 10 km define a 

plane dipping 6S0k50SW and striking N51Â°Wk20 consistent with the main shock focal mechanism. Rupture in the 

main shock evidently spread upward and bilaterally along 

the strike of this tabular zone. Most aftershocks cluster 

around the perimeter of the distribution and surround a 

relatively aseismic interior (fig. 10) which generally corresponds to the extent of main shock rupture, as described 

in detail in Spudich (1996). The less-seismic interior is 

somewhat obscured in figures 10C and 10D by off-fault 

activity projected onto the main aftershock surface; later 

figures offer a clearer view. At its southeastern end, the 

aftershock zone warps into a near-vertical surface and 

merges with the San Andreas fault. 

The diversity of focal mechanisms in the aftershock 

series is particularly great, with few events possessing 

mechanisms similar to the main shock. The larger events 

have mechanisms that span the full range of possibilities: 

right-lateral, left-lateral, reverse, and normal. Analysis of 

the state of stress implied by the aftershock mechanisms 

suggest near-total shear-stress release in the main shock 

(Michael and others, 1990; Beroza and Zoback, 1993). 

In addition to the main dipping zone, aftershocks define numerous other clusters, many of which are well removed from the primary rupture surface of the earthquake. 

Hypocenters within these clusters align with the nodal 

planes of their focal mechanisms, providing strong support for the argument that they represent coherent faults 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D 13 

\ 

Table 2.-Actual shot locations and misfits of their relocations using our velocity model 

Shot Location Misfit (Actual Ã‘ Relocation) 

Shot Ref Time Latitude Longitude At Epicentral Error Depth 

yrrnoda hrrnn s N W s azm km km 

* Mooney and Colburn (1985) 

T~ur~h~ and others (1992) 

(L. Seeber, written cornmun., 1992). While some of the 

secondary faults activated in the aftershock sequence were 

also active in the two decades preceding the main shock, 

most were not; thus they represent a triggered response to 

the stress redistribution of the main shock. 

The complexity of the sequence, in particular the large 

component of vertical displacement in the earthquake, 

Figure 7.-Map showing the epicenter of an aftershock (October 19, 

1989, at 0634 UTC) as determined from one-dimensional velocity 

model (diamond) and three-dimensional model (hexagon); positions 

of CALNET seismographic stations (triangles) and portable seismic 

instruments (dots); and contours of relative P-wave traveltimes of 

the event (dotted lines, contour interval 0.2 s). 

while surprising at first appearance, can be understood in 

terms of the local geometry of the San Andreas fault (Dietz 

and Ellsworth, 1990). Within the southern Santa Cruz 

Mountains, the San Andreas fault makes a prominent left 

Distance from Main Shock Epicenter (km) 

Figure 8.-Map of epicentral region showing epicenter of the Loma Prieta 

earthquake (dot), first motions at CALNET seismographic stations (U or 

D), contours of relative P-wave traveltimes of main shock (dotted lines, 

contour interval 1 s), and surface projection of main shock focal mechanism 

nodal lines from figure 9 (heavy lines). Note the field of compressional 

first-motions (U) in the dilatational quadrant to the north of the epicenter. 

Discrepant dilatational first-motions (D) may also be seen to the east of the 

epicenter. SJ, San Juan Bautista. 
D 14 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

(compressional) bend, connecting straighter subparallel 

segments to the north and south. Over geologic time, the 

compression within the bend must be relieved as the plates 

move either by lateral flow, subduction or mountainbuilding. This event contributed to the uplift of the mountain range (although Loma Prieta peak was downdropped 

about 10 cm), and the ratio of horizontal to vertical slip 

on the fault is predicted well by simple kinematic arguments (Dietz and Ellsworth, 1990). 

MAIN SHOCK 

The Loma Prieta Ms =7.1 earthquake initiated at 00:04 

15.28 UTC on October 18, 1989, at 37' 2.01'N 

121' 53.08W and 15.9 km depth. The 95-percent confidence ellipsoid for this hypocenter has semi-major axes 

with azimuth, plunge, and length (km) of 133O, 0Â° 0.29; 

43' 15' 0.58; and 226', 74', 0.89. Our preferred hypocenter is slightly southwest and significantly shallower 

than our previously published estimate (Dietz and 

Ellsworth, 1990). It is also slightly shallower than several 

other recent estimates (table 3). The revised depth results 

from slight changes in the crustal structure and station 

corrections. We prefer the present value principally because our revised velocity model fits the explosion traveltime data significantly better than our previous model for 

stations with epicentral distances of less than the main 

shock focal depth. 

The focal mechanism of the earthquake, or more properly, of the initial slip at the hypocenter, corresponds to a 

fault plane striking N50Â° and dipping 70Â°SW with a 

rake of 140' (Oppenheimer, 1990). This regional 

P- and 7-axis \-.*-" 

Figure 9.-Lower hemisphere projection of main shock focal mechanism showing dilatational (open circle) and compressional (+) first arrivals at CALNET stations. 

first-motion solution agrees well with teleseismic solutions (Choy and Boatwright, 1990; Kanamori and Satake, 

1990; Romanowicz and Lyon-Caen, 1990), and simple 

geodetic models of the static displacement field (Arnadottir 

and others, 1992; Marshall and others, 1991; Lisowski 

and others, 1990). 

FOCAL DEPTH 

A trade-off exists between origin time, earthquake depth, 

and model parameters for the deepest events in our catalog due to poor resolution of the deep velocity structure. 

Therefore, we conducted a series of experiments to test 

the stability of the main shock hypocenter. In one experiment, a range of starting depths between 1 and 22 km 

was used to test the convergence of the HYPOINVERSE 

soluton. All runs yielded main shock depths between 15.6 

and 15.9 km. When the depth of the shock was held 

fixed at the trial starting depth (again varying from 1 to 

22 km), the hypocenter at 16.5 km depth had the lowest 

rms (0.09 s; fig. 11). Fixed-depth solutions outside of the 

range from 15 to 17.5 km had significantly higher rms. 

Finally, we examined the performance of a suite of velocity models computed using the same 108 events used to 

develop the final model. Here we fixed the depth of the 

main shock and simultaneously solved for the aftershock 

hypocenters and station corrections that minimized the 

overall data misfit. The lowest overall data variances were 

obtained when the initial slip of the main shock was fixed 

between 15 and 17 km; the minimum data variance was 

obtained with the main shock depth fixed at 16 km (fig. 

12). 

NUCLEATION PROCESS 

The process by which the earthquake started is of interest, as it did not simply propagate uniformly away from 

the hypocenter beginning at the above origin time. As 

noted by Wald and others (1991, 1996), the CALNET 

origin time precedes the initiation of strong motion from 

the hypocenter by about 2 s, suggesting that our origin 

time corresponds either to a foreshock or to a more complex initiation process. Amplitudes of the high-frequency 

waves radiated in the first 1.5 s of the event were generally too weak to trigger strong-motion instruments (typically a vertical acceleration of 0.01 g), indicating relatively 

small slip at the initiation of rupture. Most strong-motion 

recorders, however, did trigger before the arrival of the 

shear waves radiated by the initial slip, and thus recorded 

them. Peak velocities of the shear waves corresponding to 

the S waves of the initial rupture were compared to peak 

velocities predicted in the frequency band from 2 to 5 Hz 

for events of various magnitudes using the relation of 

Joyner and Boore (1988): 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS 

where y is the ground motion parameter to be predicted, 

M is the earthquake magnitude, rn is the shortest distance 

from the recording site to the vertical projection of the 

fault rupture on the surface, h is the depth of the earthquake, and the parameters for predicting peak velocity 

are: a=2.17, b=0.49, c=0, d=-1.0, k=-0.0026, and 

s = 0.17. The observed peak velocities of the initial rupture 

generally fall within one standard deviation of the values 

predicted for a M=5.0 event at 16 km depth (fig. 13). 

Large-amplitude, high-frequency motions, corresponding to the start of the "main shock" as inferred from strong 

motion modeling studies (Beroza, 1991, 1996; Wald and 

others, 1991, 1996; Steidl and others, 1991, 1996) began 

1.6 s after the initial event, or at 00:04 16.9 UTC. P- and 

S-arrival times for both the foreshock and the main shock 

could be read at a number of stations (table 4). P arrivals 

were picked from vertical components and S arrivals were 

picked from horizontal components of acceleration after 

integration to velocity. Using the program VELEST to 

determine the separation between their hypocenters, we 

find the main shock locates 0.3 km to the southeast, 1.1 

krn above the initial event, and 1.61 s later. However, the 

uncertainty in this location due to the uncertainty in the 

identification and timing of the phases and in the shear 

velocity structure precludes us from rejecting the hypothesis that the two events nucleated at the same point. 

While it is possible to interpret the initial event as an 

immediate M=5 foreshock, examination of digital 

strong-motion records from instruments in the near-field 

of the event suggest otherwise. Integration of the observed 

accelerograms to displacement shows a continuous growth 

in displacement from the moment of initiation (fig. 14). 

To model the displacement seismograms, we use the complete synthetic seismogram for a point dislocation in a 

half-space (Johnson, 1974) and assume the fault orientation determined from the first-motion data. The seismograms are modeled by determining the seismic moment 

release as a function of time, a well-posed linear inverse 

problem. The best fit to the observed displacement seismograms shows a continuous growth in seismic moment 

beginning at the high-frequency origin time of 15.28 s 

(fig. 15A). The total moment release by 1.6 s into the 

event, just up to the time of the main shock, is equivalent 

to a ~=5'/2 earthquake (fig. 1 SB), significantly larger than 

the magnitude determined from 2 to 5 Hz waves (fig. 13). 

Thus, the initial event appears to mark the beginning of a 

continuous and relatively smooth process of moment release that only began to propagate away from the hypocenter after 1.6 s. 

DEVELOPMENT OF THE 

AFTERSHOCK ZONE 

INITIATION 

The divide separating the main shock from its aftershocks is no less ambiguous than the boundary between 

the foreshock and the main shock. Modeling studies of 

the main shock, summarized by Spudich (1996) suggest a 

duration of at least 6 s and not more than 15 s, with most 

of the slip occurring in the first 10 s of the event. If 

aftershocks are defined as distinct earthquakes removed 

in time but not necessarily in space from faulting during 

the propagation of the main shock, the earliest recognizable aftershock, M = 4.7, occurred 32 s after the start of 

the main shock (Simila and others, 1990). Earlier aftershocks of similar or larger magnitude, however, might 

have occurred but might have been obscured by the arrival of scattered (coda) waves from the main shock, as 

this first aftershock just emerges from them (see fig. 2-6 

of McNally and others, 1996). 

During the first 10 minutes of the sequence about 20 

aftershocks (Md=4) can be identified in the CALNET seismograms. Locations could be determined for eight of these 

aftershocks. McNally and others (1996) used records from 

their digital accelerometers to locate these and seven additional events. After 10 minutes, our catalog appears to 

be complete to Md=4.0. Using the frequency-magnitude 

density distribution (for example, in fig. 16) to estimate 

completeness at later times, hours 1 to 6 appear complete 

to Md=3.0, and hours 6 to 13 complete to Md=l.5. After 

hour 13 of the aftershock sequence, our catalog appears 

complete to Md=l.O. 

A cautionary note is in order, however, as systematic 

underreporting of earthquakes of all sizes will occur when 

smaller events are buried in the codas of larger events. At 

the end of the first day, the aftershock rate was down to 

one event (Md 21.0) every 2 minutes, with an average 

event duration of -10 s. Missing events after the first day 

should, therefore, be few in number, which agrees with 

the frequency-magnitude evidence for completeness discussed above. 

FORMATION 

The aftershock zone formed extremely rapidly, assuming essentially its final shape by the end of the first day 

(fig. 10). Even the incompletely recovered first 46 minutes of activity conforms well with the later activity, clearly 

mapping out the main dipping zone as extending approximately 45 km along strike. Some of the early aftershocks 

fall outside of the main zone and are associated with 
Figure 10.-Spatial and temporal development of the sequence. Earthquakes within (A) 0.001 to logarithmic time intervals which would contain equal numbers of events if the aftershock rate 

0.032 days, (B) 0.032 to 1.0 days, (Q 1.0 to 31.6 days, and (D) 31.6 to 730 days after the main decayed with time as dNIdt = t -I. Earthquake symbols are the same as in figure 3. 

shock. Maps, and longitudinal (2-Z') and transverse (Y-Y') cross sections correspond to equal 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 3.-Main shock hypocentral estimates 

Source Origin Time Latitude Longitude Depth 

yrrnoda hrmn s N W km 

This study 891018 0004 15.28 37' 2.01' 121' 53.08' 15.94 

Dietz and Ellsworth (1990) 15.21 37' 2.37' 121' 52.81' 17.85 

U.S. Geological Survey Staff (1990) 15.25 37' 2.19' 121' 52.98' 17.6 

Eberhart-Phillips and Stuart (1992) 15.34 37' 2.02' 121' 52.85' 17.16 

Pujol (1995) 15.26 37' 1.92' 121' 52.98' 16.3 

Roecker and Gupta (this volume) 15.45 37' 2.22' 121' 53.46' 15.60 

aftershock activity on secondary faults, principally to the 

northeast of the San Andreas fault. 

The aftershock zone apparently lengthens with time to 

the southeast where it overlaps the seismically active central segment of the San Andreas fault (figs. 3, 10). While 

some of this extension reflects the occurrence of earthquakes normally expected to occur along this part of the 

San Andreas fault, the analysis of Reasenberg and Simpson 

(1992, and this chapter) shows that these events occurred 

at a higher rate after the mainshock and decayed in an 

aftershock-like manner. Including them in the definition 

of the aftershock zone would increase its length to about 

60 km. 

Fixed Depth (km) 

Figure 11.-Final traveltime rms-residual versus depth for the Loma 

Prieta main shock location. In these HYPOINVERSE tests the epicenter 

was free to move, but the depth was fixed. The main shock's depth of 

15.9 km from an unconstrained relocation (arrow) falls within the 

fixed-depth range (15.0-17.5 km) yielding the lowest rms values. 

In general, the aftershocks are highly clustered and tend 

to populate the periphery of the main shock rupture as 

determined from studies of the main shock seismograms 

(Beroza, 1991, 1996; Wald and others, 1991, 1996). In 

nearly every case, these clusters contain at least one large 

aftershock (Md>4) from the first few hours' activity. The 

persistence of activity in these clusters is particularly striking (figs. 10, 17) and clearly cannot be explained as secondary aftershock activity spawned by a large aftershock 

from early in the sequence. One secondary aftershock sequence, however, can be recognized following the April 

18, 1990, ML=5.4 Chittenden earthquake (north of PG, 

fig. 17), which was also the largest event in the entire 

aftershock sequence. Note the rapid decay of its aftershocks, compared to the overall decline in activity (figs. 

12 14 16 18 20 

Fixed Depth of Main Shock (km) 

Figure 12.-Residual variance of traveltimes for 108 earthquakes used 

to develop final velocity model for models computed with the main 

shock depth held fixed. Each data point corresponds to the residual 

variance obtained by inverting the traveltime observations for optimum 

earthquake hypocenters and station delays, with the constraint that the 

main shock depth has the value shown. Results show a clear preference 

for the unconstrained model solution that places the main shock at a 

depth of 15.9 km. 
I 

AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS Dl9 

DECAY 

As a whole, the decay of the aftershock sequence 

conforms with the modified Omori's law (Utsu, 1969; 

Reasenberg and Simpson, this chapter). We determine a 

best-fit to Omori's law, dNldt = K/(t + c)P, by binning the 

aftershocks into equal log time intervals, starting 11.5 

minutes after the main shock (fig. 19). Using all events 

Md61.0 for the entire two-year catalog, we find: 

p = 0.94k0.01, c = 0.34k0.04, and K = 670125. Observed 

rates during the first 13 hours of the sequence (first 6 

intervals, figure 19) undoubtedly underestimate the true 

rate of Md21.0 events, due to grossly incomplete reporting at this magnitude threshold. When only events Md63 

are considered, p = 1.07k0.04, c = 0.04k0.03, and 

K = 2633.4. 

The end of aftershock activity may be defined as the 

intersection of the Omori rate dNldt with the rate preceding the earthquake. Reasenberg and Simpson (this chapter) estimate this time as 2.3k0.3 years, based on activity 

(Md61.5) within the broadly defined aftershock zone 

through April 17, 1990. In this study we define the "main 

aftershock zone" with a series of polygons and depth 

ranges (fig. 2 15). Because the pre-earthquake rate varies 

widely throughout the aftershock zone, we compare the 

aftershock decay to the background rate (average number 

of Md>l.O events per day from 1980 to October 17, 1989) 

in three different areas: the main, dipping part of the aftershock zone (portion of the main zone which lies in 

regions LPS, LPC, and LPN of figure 6; generally north 

of 36' 58' ); the part of the main zone along the 

microseismically active part of the San Andreas fault at 

the south end of the sequence (portion of the main zone in 

region CSA of figure 6); and the previously active part of 

the Castro fault (commonly thought to be the Sargent 

fault-see below) to the southeast of the main shock hypocenter (region SAR in figure 6). Within the main dipping zone, aftershock activity will not reach the very low 

pre-Loma Prieta rate of 0.02 eventslday until over 25 years 

have passed, or until 2015. This long time reflects the 

very low level of pre-Loma Prieta seismicity in this zone. 

Activity along the San Andreas fault immediately southeast of the earthquake would be expected to intersect the 

background rate of 0.16 eventslday after t=1.5 years if we 

model only the aftershocks prior to the April 18, 1990, 

Chittenden earthquake. However, this event further boosted 

the seismicity rate in this region with its own aftershock 

sequence, lengthening the actual time to decay to background rate. Activity along the Castro fault met the background rate after just one year. 

Epicentral Distance (km) 

Figure 13.-Peak horizontal velocity of 5-waves recorded by strong-motion accelerographs from the first 1.0 s of rupture. 

Amplitudes were measured from accelerograms after integration to velocity. Measurement window restricted to direct arrivals 

from the foreshock and excludes arrivals from the main shock. Curves are predicted mean (solid) and Â±l (dashed) peak 

horizontal velocity from Joyner and Boore (1988) for M=5 earthquake at 16 km depth. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 4.-Foreshock and main shock arrival times (seconds after 891018 0004 15.35) 

Station Latitude Longitude Foreshock Main Shock 

N W P wt S wt P wt S wt 

C563 

SAOC 

COO7 

BSRZ 

WAHO 

LGPC 

BRAN 

BKSB 

HBTM 

HSFM 

JBLM 

JMPM 

HPLM 

HQRM 

A Santa Teresa Hills 

c Walter's House 

WAHO 

Branciforte 

D U.C. Santa Cruz 

Seconds 

Figure 14.-Displacement seismograms of vertical component of motion 

from broad-band integration of digital accelerogaph stations located in the 

near-field of the earthquake (solid line) and synthetic seismograms (dot-dash 

line). Bandpass of displacement filtered with 1-pole Butterworth filter at 

0.01 Hz. Origin time for each seismogram is P-wave arrival time of the 

Seconds 

foreshock. Synthetic seismograms determined by inverting observed displacements for seismic moment release rate of point source double couple 

with orientation of main shock focal mechanism. Divergence of fit after 2 

s reflects breakdown of point source approximation and may reflect change 

in radiation pattern. 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D2 1 

0 1 2 3 

Seconds 

0 1 2 3 

Seconds 

Figure 15.-Cumulative seismic moment release versus time determined by inversion of near-field displacement seismograms (figure 14). Both the 

total seismic moment (A) and equivalent moment magnitude (B) are shown. Results indicate that the foreshock released more moment than would be 

predicted by either its amplitude magnitude of 5 (figure 13) or its triggering of strong-motion stations. 

01234567 

MAGNITUDE 

Figure 16.-Number of earthquakes in each 0.1 magnitude-unit bin for 

aftershocks (excluding activity on October 18, 1989) in our catalog. The 

aftershock catalog is complete at Md>l.O and the b-value for the sequence is 0.8310.01. 

STRUCTURE OF THE AFTERSHOCK 

ZONE 

The overall form of the aftershock zone is dominated 

by the central dipping zone, or main zone, that rises from 

the main shock hypocenter toward the San Andreas fault 

to the northeast. About 75 percent of the aftershocks fall 

in this zone which extends from 2 to 18 km depth. The 

coincidence between the N5 1 OWk2O strike and 6S0kS0SW 

dip of the zone and the main shock fault plane leaves 

little doubt about the existence of a primary, causal relationship between these events and the main shock. 

There is much more to the sequence, however, than 

continuing faulting in the style of the main shock dislocation. Both the occurrence of aftershocks within a much 

larger crustal volume, and the wide range of aftershock 

focal mechanism types demonstrate that numerous faults 

participated in the post-main shock process of crustal adjustment. As first noted by Oppenheimer (1990), almost 

none of the aftershock focal mechanisms resemble the 

main shock. Indeed, with the exception of aftershocks at 

the base of the zone near the main shock, it can be fairly 

AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS 

18 19 20 

OCT 1989 

Figure 18.-Cumulative number of events versus time for (A) the first 3 days of the sequence and (B) 2 years of the 

sequence. Heavy line shows all Md>3.0 events in our catalog and thin line shows all MdSl.O events. 

stated that no seismic displacements took place on the 

main shock fault once its dynamic motions stopped. Thus, 

the central dipping zone and the more removed parts of 

0.01 0.1 1 .o 10. 100. 1000. 

TIME AFTER MAIN SHOCK (DAYS) 

Figure 19.-Seismicity rate (number of earthquakes per day) versus time 

after the main shock (in days) for all MdS1.O aftershocks in our catalog 

(diamonds). The line shows the best fit to Omori's Law, dNIdt = 

K/(t + c) P, with parameter values given above. 

the sequence both speak to a complex process of adjustment that bears little resemblance to the typical San Andreas earthquake sequence in which the aftershocks 

continue the main shock faulting process (for example, 

Baton and others, 1970). 

To facilitate the examination of the aftershock zone we 

subdivide our discussion into five parts corresponding to 

the main dipping zone and four geographically distinct 

areas surrounding it. 

MAIN ZONE 

The main zone consists of a tabular volume of aftershock hypocenters that maintains a nearly-constant dip of 

60' to 65' for a distance of 42 km, centered approximately on the main shock (figs. 10, 20, 22,23). The depth 

at the base of the zone averages about 18 km to the northwest of the main shock and shoals to about 8 km to the 

southeast at the point where it joins the San Andreas fault. 

The top of the dipping zone lies about 6 km below the 

surface, on average, and locates directly beneath the surface trace of the San Andreas fault. Seismicity shallower 

than 6 km primarily locates northeast of the San Andreas 

trace and defines no single large-scale structure. Specifically, the main zone does not appear to cross the San 

Andreas fault as a continuous body. 

Viewed in longitudinal cross section (figs. 10, 23), the 

main zone is seen to be composed of numerous clusters of 
D24 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

activity, generally located on the periphery of 

the zone. The interior of the zone is relatively 

devoid of activity. As has been noted by numerous authors, the less seismic interior of 

the zone correlates well with the regions that 

slipped during the main shock (see Spudich, 

1996). The absence of aftershocks on the portions of the fault that slipped during the main 

shock is a general feature of the main shock1 

aftershock process, at least for California 

earthquakes (Mendoza and Hartzell, 1988). 

Transverse cross-sections show that the 

main zone has an average width of 1 to 2 km 

(fig. 22). Because the width is resolved by the 

hypocentral locations (fig. 5) and has relatively sharply defined edges, it implies that 

these aftershocks map out a volume of relatively weak rock triggered into activity by the 

main shock. We cannot tell, however, on the 

basis of these data, where the main shock rupture passes in relation to the main zone. It 

could locate equally well within or on either 

edge of the body. 

Aftershock focal mechanisms in the dipping 

zone do little to clarify this picture. With the 

exception of a few focal mechanisms located 

at the base of the zone just northwest of the 

main shock, none of the aftershock focal 

mechanisms in the main zone show rightlateral, oblique-thrust motion on the main 

shock plane. Rather, they display a very wide 

range of mechanism types and orientations 

(fig. 20), with the only unifying trend being a 

tendency for the P-axes to orient approximately normal to the fault plane (fig. 24). This 

fault-normal coordination of the P-axes is consistent with an approximately uniaxial stress 

field acting perpendicular to the main shock 

fault plane (Zoback and Beroza, 1993) and 

thus is incompatible with the stresses that 

drove the main shock faulting. 

It is possible that some aftershock mechanisms may have closely resembled the main 

shock mechanism while the grid search method 

of FPFIT selected a different mechanism as 

the optimal solution. To test this hypothesis, 

we re-examined each aftershock through December 1989 to determine if the main shock 

focal mechanism (initial slip) could provide a 

reasonable explanation of the aftershock's observed first-motion polarities. Using FPFIT, 

we constrained the solution to be the main 

shock mechanism and compared the resulting 

misfit to the confidence region of each 

I - LPC 

H0 10 

DISTANCE (KM) 

Figure 20.Ã‘Continue 

DISTANCE (KM) 

Figure 20.-Continued 


AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D3 1 

Figure 21.-Key maps showing selection criteria used in following figures. Labeled dots show landmarks from figure 6. Dashed line is the 

perimeter of the study region. (A) Selection boxes (solid outlines) and 

corresponding cross sections (labeled lines) for the series of transverse 

depth sections in figure 22: (B) Selection polygons and corresponding 

depth ranges used to define "main zone" seismicity for figure 23A,B. 

Events which are outside of the "main zone" regions but inside the Z-2' 

box (solid line) are plotted as "outside-main zone" events in figure 23C,D. 

unconstrained mechanism, only considering constrained 

solutions having misfits 50.15. We found that at the 80-percent confidence level only 6 percent of the aftershock 

mechanisms (47 of 730) matched the main shock. Next 

we freed the rake on the northwest-trending plane and 

found only 10 percent more events (74 of 730 within the 

80-percent limit) which fit any kind of motion on the 

main shock rupture plane. These results strongly imply 

that the post-main shock stress field fundamentally differs 

from the pre-main shock stress field. 

Although there is great diversity and heterogeneity in 

the distribution of aftershock focal mechanisms, several 

large-scale trends are noteworthy. Right-lateral strike-slip 

and right-lateral oblique-thrust events tend to cluster 

around the perimeter of the aftershock distribution, principally to the southeast of the main shock (figs, 20, 25). 

The line of events at the base of the zone to the southeast 

of the main shock display nearly uniform slip on planes 

slightly steeper than the main shock and with a slightly 

more horizontal rake. These mechanisms are consistent 

with the greater component of strike-slip displacement in 

the main shock to the southeast of the epicenter inferred 

from waveform modeling (see Spudich, 1996), and geodetic modeling (Lisowski and others, 1990; and Arnadottir 

and others, 1992). 

Strike-slip mechanisms also occur at the northwest end 

of the zone. These mechanisms, however, are rotated about 

45' to the trend of the zone, and have fault-normal P-axes. 

The largest aftershock in the early part of the sequence, 

the M=5.3 event of October 18 at 00:41 (event D, fig. 20 

D, E) has this mechanism. Strike-slip events with the same 

fault plane orientations, but the opposite sense of slip 

(fault-normal T-axes) are also observed, such as the M=4.7 

event within the interior of the main zone a few kilometers to the northwest and above the main shock (event X, 

fig. 20 G, H). Thus, the pattern of fault-normal compression is not universal, even within the interior of the zone. 

Normal faulting mechanisms principally occur below 7 

km depth and within the interior of the main zone (fig. 

25). These events also have fault-normal coordinated 

T-axes. Thrust mechanisms are scattered through the entire aftershock zone. Most of them occur northwest of the 

main shock, and many occur at depths above 6 km. Most 

thrust mechanisms strike sub-parallel to the main zone 

(and San Andreas surface trace) and have fault-normal 

P-axes. Some of the events toward the northern end of the 

zone, however, (events S and C, fig. 20 D, E) could represent pure reverse slip on the main fault plane. Nodal planes 

of the thrust events more commonly have too shallow a 

dip (-45') to match the more steeply dipping main shock 

fault plane and thus most probably occur on secondary 

structures. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Figure 22.-Series of transverse cross sections from northwest 

(A) to southeast (H) along the San Andreas fault. Well-located 

M&l.O events from (Al-HI) January 1, 1969, to October 17, 

1989; and (A2-H2) October 18, 1989, to October 17, 1991. 

Symbol sizes are scaled to rupture area assuming a 30-bar stress 

drop. Inverted solid triangle represents the surface trace of the 

San Andreas fault, open triangle shows the Sargent fault. See 

figure 21A for selection boxes and section endpoints chosen to 

minimize the projected width of the aftershock zone. 0 10 

DISTANCE (KM) 

c 

0 10 

DISTANCE (KM) 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS 

0 10 

DISTANCE (KM) 

Figure 22.-Continued 
- 0 

- <0 

- 0 

- in 

- 0 

- ^ 

- 0 

- CO 

- 0 

- CM 

- 0 

- TÃ 

- 0 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

P-axes 891018-911017 

0 - 5.99 km depth 

Figure 24.-Map view projections of P-axes from well-determined mechanisms of Md>l .5 earthquakes. Events from October 18, 1989, to October 

17, 1991, at (A) 0-6 km depth, (B) 6-10 km depth, (0 10-14 km depth, 

and (D) 14-20 km depth; and (E) January 1, 1969, to October 17, 1989, 

0-20 km depth. (F) a few focal mechanisms of small events near the San 

Andreas fault, 1969 to 1989. 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D37 

SAN ANDREAS FAULT 

At its southern end, the aftershock zone merges 

seamlessly into the activity of the creeping central segment of the San Andreas fault where fault slip occurs as a 

mix of earthquakes and fault creep (Wesson and others, 

1973b). Focal mechanisms of earthquakes on this section 

of the San Andreas fault are homogeneous and reflect 

dextral slip on the San Andreas fault. Slip in the main 

shock, as determined from both dynamic (Beroza, 1991; 

Steidl and others, 1991; Wald and others, 1991) and static 

models (Lisowski and others, 1990; Snay and others, 1991; 

and Marshall and others, 1991), extended for no more 

than 15 to 20 km to the southeast of the hypocenter. The 

seismicity rate in the southernmost section decays differently from the rate in the dipping portion of the aftershock zone (Reasenberg and Simpson, this chapter). Thus, 

activity along the southeastern 25 km of the aftershock 

zone probably reflects stress transfer to it from slip in the 

main shock. 

About 10 km southeast of the main shock epicenter, the 

main zone begins to steepen in dip and it warps toward a 

slightly more northerly strike. It continues to steepen until 

it merges with the San Andreas fault at a point about 8 

km northwest of Pajaro Gap (PG in figs. 6, 21, 23). As 

the zone steepens, it also narrows in width, until it has no 

measurable width (a.7 km) after merging with the San 

Andreas fault. 

10' 122' 50' 40' 30' 

Figure 24.Ã‘Continue 

The southern terminus of rupture in the main shock 

coincides with the point where the aftershock zone reaches 

the San Andreas fault. The end of the rupture may be 

marked by the clusters of aftershocks located 20 to 21 km 

southeast of the main shock hypocenter. The largest aftershock, the M=5.4 Chittenden earthquake of April 18, 1990, 

locates immediately to the southeast of this point at the 

top of the aftershock zone. This earthquake triggered a 

major secondary aftershock sequence that extended for 10 

km to the southeast along the San Andreas fault (fig. 17). 

This sequence filled a 7-km-long gap on the fault between 

3 and 8 km depth that was visible in the pre-Loma Prieta 

seismicity (fig. 23A). The mechanisms of these events 

show predominantly right-lateral strike-slip motion on 

northwest-trending planes. Any dip to these fault planes is 

believed to be an artifact caused by lateral refractions 

from unmodeled local velocity changes, as previously stated. 

CASTRO AND SOUTHERN SARGENT FAULTS 

Directly to the northeast of the San Andreas fault along 

the southeastern part of the aftershock zone (region SAR, 

fig. 20 N, O), aftershock activity defines a shallow (1-4 

km), near-vertical, right-lateral strike-slip fault. This fault 

also stands out as a prominent feature in the pre-Loma 

Prieta seismicity (fig. 3; Olson and Hill, 1993). This activity has been commonly assumed to be on the Sargent 

Focal Mechanisms 6901 01-891 01 7 
+ left-lateral oblique A 

A 0 right-lateral oblique normal V 

9, I I I I, 

- 

thrust 

normal - 

DISTANCE (KM) 

Qepth sections of aftershocks within the 2-2' box (see figure 21B) having SW-dipping (open circle) and NE-dipping (square) planes. (B) Right-lateral oblique-normal 

"a1 mechanisms with one nodal plane sub-parallel (1 30Â°Â±45 to the events (- 154' 5 rake 5 - 1 15', square); thrust events (65' 5 rake <: 1 14', triangle-point-up); 

Symbol type depends on rake, symbol size is proportional to normal events (-1 14' 5 rake < -65O, triangle-point-down); and left-lateral and oblique events 

;ral strike-slip events ( I rake I> 155') on SW-dipping (x) and (-64' 5 rake 5 64', +). 

jlanes; Right-lateral oblique-thrust events (1 15' 5 rake 5 154') on 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D39 

fault (for example, Baton and others, 1970). Our locations, as well as those of Olson and Hill, however, lie 

distinctly east of and at an acute angle to the southern 

Sargent fault. They coincide, however, with the Castro 

fault, first described by Allen (1946). We believe that this 

seismically defined fault is most probably the subsurface 

expression of the Castro fault and is not associated with 

the Sargent fault. 

A few aftershocks are also scattered between the traces 

of the Sargent and San Andreas faults. Most of these events 

also show right-lateral strike-slip on northwest planes, but 

a few shallow thrust events occur as well. Only 4 of the 

events in this region are Md23.0, the largest a Md=3.2 on 

December 31, 1989. As noted previously, the rate of activity along the Castro fault underwent a step-like increase 

at the time of the Loma Prieta earthquake, then declined 

to the pre-event rate within about one year. 

FAULTS NORTH OF THE MAIN ZONE AND 

EAST OF THE SAN ANDREAS 

The Loma Prieta earthquake triggered activity within a 

plexus of faults in the footwall block, principally within 

the southern Santa Cruz Mountains to the east of the San 

Andreas fault. Aside from the activity along the Castro 

fault (described above) the aftershock activity in this region occurred on faults that were virtually aseismic during the two decades before the earthquake (figs. 3, 22, 

23). Areas that had been active before the main shock 

were stilled by the earthquake. 

The overall pattern of aftershocks in this region reflects 

release of fault-normal compression, principally on reverse faults. These events are generally deepest adjacent 

to the San Andreas fault, shallow to the northeast, and are 

consistent with movement on the Shannon, Monte Vista, 

and Berrocal faults (fig. 22). Within this general framework, however, the case can be made for reverse faulting 

on steeply northeast-dipping planes, in addition to the 

named northwest-dipping faults. 

Strike-slip events, although rare, are also observed. 

These occur at shallow depth (1-2 km), and directly underlie the northern part of the Sargent fault (fig. 20 F, H). 

At only slightly greater depth, reverse faulting predominates within the same area. 

One of the clearest examples of a distinct secondary 

fault outlined by aftershock activity appears updip of the 

main shock hypocenter. This activity clusters in a 

near-vertical plane between 6 and 9 km depth lying between the surface traces of the San Andreas and Sargent 

faults. This cluster was initiated by a M=4.3 aftershock on 

October 25, 1989, at 0127 UTC. The focal mechanism 

implies pure reverse motion (SW side down) on a vertical 

plane parallel to the San Andreas (event a, fig. 20 G, H; 

appendix A). 

The total amount of aftershock activity in the northernmost portion of the foot wall of the Loma Prieta earthquake was relatively modest, with only four well-located 

aftershocks Md23 (region SCV, fig. 20 A, B). The largest 

of these, Md=4.7, occurred 3 minutes after the main shock 

along the northeastern perimeter of the entire sequence, 

just to the west of the Shannon fault. Because P-wave 

arrivals for this event were obscured by the main shock 

coda waves, a reliable focal mechanism could not be determined. We strongly suspect, however, that this event 

was a shallow reverse-faulting event associated with the 

Shannon fault. 

Field investigations along the Monte Vista and Shannon faults revealed systematic evidence of tectonic shortening caused by the Loma Prieta earthquake (Schmidt and 

others, 1996). The events along this northernmost extension of the aftershock zone (fig. 20 A, B) are consistent 

with minor displacement along these faults during the main 

shock or in the earliest moments of the aftershock sequence. The prompt occurrence of the Md=4.7 event at 

shallow depth just west of the trace of the Shannon fault 

supports this interpretation. 

HANGING WALL ACTIVITY 

Although the majority of all aftershocks locate southwest of the San Andreas fault, almost all of these lie 

within the main zone, and comparatively few locate in the 

hanging-wall block of the earthquake. The aftershock activity in the hanging wall (region MON, fig. 20 P, Q) 

concentrate in one large and several smaller clusters. The 

large cluster contains the Md=4.5 event of October 19, 

1989, at 1014 UTC (event 2, fig. 20 P, Q; appendix A) 

and its aftershocks. This sub-vertical cluster trends 143' 

and is centered about 10 km south-southeast of the main 

shock. It is about 5 km long and lies between 5 and 8.5 

km depth. The focal mechanism indicates right-lateral 

strike-slip on a 170' trending vertical plane, with the P-axis 

in fault-normal coordination with the main zone. This secondary fault lies near the projected position of the Zayante 

fault (R. Jachens and A. Griscom, written commun., 1992), 

but has focal mechanisms and fault plane strikes that are 

incompatible with their occurrence on this reverse fault. 

A second cluster locates 9 km northwest of the first, at 

a depth between 1 and 4 km. The largest event in this 

group also has a north-striking, right-slip fault plane solution (event 11, fig. 20 P, Q; appendix A). Ten kilometers 

further to the northwest, strike-slip faulting in the 

hanging-wall activity appears again, this time with a northwest strike and reverse component (event I, fig. 20 D, E; 

appendix A). The fourth minor cluster locates in Monterey 

Bay, a few km seaward of the shoreline and directly south 

of the main shock epicenter. The largest event in this 

cluster (event 2b, fig. 20 P, Q; appendix A) shows right 
D40 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

slip on a N50Â°W near-vertical plane. No young faults 

have been identified at the surface in the vicinity of these 

other clusters. 

RELATIONSHIP OF PRESHOCKS TO 

THE MAIN SHOCK AND AFTERSHOCKS 

The detailed record of seismic activity afforded by the 

CALNET during the 20+ years before the Loma Prieta 

earthquake bears little resemblance to the distribution of 

aftershocks (fig. 3; Olson and Hill, 1993). In only three 

peripheral areas does it appear that the aftershocks occurred on the same faults displaying activity in the two 

decades before the shock. These are the San Andreas fault 

southeast of the main shock, the Castro fault to the south 

and east of the earthquake, and the zone of reverse faulting at the extreme northwestern end of the zone, generally 

in the area near Stevens Creek Reservior (SC, fig. 6). As 

noted above, all of these areas continued to be active after 

the Loma Prieta earthquake and each experienced an accelerated rate of earthquake production following the main 

shock (Reasenberg and Simpson, 1992, and this chapter). 

As discussed by Olson and Hill (1993), the main zone 

had very little activity for at least the 20 years leading to 

the main shock (figs. 3, 22, 23). While the deep activity 

forming the prominent "U-shaped" band of activity that 

was subsequently filled by the main shock (U.S. Geological Survey Staff, 1990) included the sparse main zone 

activity, it was principally constructed from activity to the 

northeast of the main zone. These northeastern areas, however, were virtually silent during the aftershock sequence, 

as can be seen by comparing the pre- and post-earthquake 

event distributions (fig. 23). In general, the pre-earthquake 

seismicity reflects the release of a north-south compressive stress, as would be expected for the San Andreas 

fault system (fig. 24 E, F). 

Several important sequences of ML=4-5 earthquakes 

preceded the main shock. The two events near Lake Elsman 

in 1988 and 1989, ML=5.0 and ML=5.2 respectively, are 

well-known to be the largest events near the Loma Prieta 

earthquake in over 20 years (Olson and Hill, 1993; Sykes 

and Jaum6, 1990). These events ruptured adjacent parts of 

deep fault at the intersection of the Sargent fault with the 

San Andreas fault near what was to become the northwest 

end of the Loma Prieta rupture. As with the other zones 

noted above, seismicity in the Lake Elsman zone was not 

activated by the main shock. 

In addition to the 1988 and 1989 events, several notable sequences of M=4-5 events also occurred near the 

main zone in the 1960's. These include the events of November 16, 1964 (ML=5.0), October 14, 1966 (ML=4.0), 

September 7, 1967 (ML=4.7), December 18, 1967 

(ML=5.3), and March 21, 1968 (ML=4.3). Although the 

catalog locations of these events place them in the general 

region, they lack the precision to draw any firm conclusions regarding their laskdciation to the Loma Prieta earthquake. Using the method of joint hypocentral determination 

(Dewey, 1972), we can accurately locate these events with 

respect to the modern data. Furthermore, the 1967 and 

1968 events were recorded by the CALNET and can be 

accurately located without resort to JHD. 

Relocation of these events using both JHD and absolute 

methods shows that they are significantly more tightly 

clustered than the catalog locations would suggest (fig. 

26, table 5). The 1964 and 1966 events locate quite near 

the San Andreas fault, and are plausibly on it. The activity in 1967 and 1968, however, form a clear northwest 

trending lineation, subparallel to and removed 2 krn to the 

northeast of the San Andreas fault at the latitude of the 

main shock hypocenter. This trend also contains the 

ML=4.5 event of August 19, 1982. A reliable focal mechanism was determined for the largest event (December 18, 

1967), and it shows right-lateral strike slip on a 

N50Â°W-strikin fault. We conclude that these events do 

indeed fall on a single fault, possibly the Sargent fault, 

but probably not the San Andreas fault. As with most of 

the other significant concentrations of pre-Loma Prieta 

seismicity, activity along this zone was curtailed by the 

main shock. 

Finally, we address the question of whether or not the 

San Andreas fault was active adjacent to the Loma Prieta 

Figure 26.-Map of ML>4.0 events from 1964 through October 17, 1989; 

star = main shock, circle = ML>5.0, diamond = MG4.0. Earthquakes 

occurring in the 1960's were located relative to a subset of later events 

(solid symbols, table 5) by a joint hypocentral determination (JHD). In 

the JHD, events after 1970 were fixed at their locations from this study, 

but their origin times were free to move. 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS 

Table 5.-Earthquakes ML 24.0 located by joint hypocentral determination 

Origin Time Latitude Longitude Depth M, 

yrrnoda hrrnn s N W krn 

rupture and main zone of aftershocks between 1969 and 

the earthquake. A critical examination of the pre-Loma 

Prieta activity shows that there are a few events in the 

central part of the zone that lie vertically below the trace 

of the San Andreas fault (fig. 22 Dl; fig. 3A). They suggest the possibility of some activity on a vertical San 

Andreas fault. Although focal mechanisms for these small 

events are difficult to obtain, a number of them indicate 

right-lateral slip on planes parallel to the San Andreas 

fault (fig. 24F). If these events locate on the primary fault 

that moved in the 1906 earthquake, they represent only 

very minor subsequent movement. 

DISCUSSION 

Perhaps the most intriguing feature of the Loma Prieta 

aftershock sequence is its almost complete disassociation 

from the preceding seismicity. Only at the extreme ends 

of the sequence, well removed in space from the main 

shock slip surface, do we observe aftershock activity of 

the same style and at the same location as the preshocks. 

Nearer the source, the aftershocks activated numerous 

faults not seen before the earthquake, and their focal 

mechanisms reflect a first-order change from release of 

north-south compression expected for the San Andreas 

system to release of fault-normal compression. Indeed, it 

appears probable that the main shock effectively released 

all of the shear stress acting across its fault plane (Michael 

and others, 1990), leaving a residual stress field in an 

average state of uniaxial compression (Zoback and Beroza, 

1993). 

Activation of the full aftershock zone proceeded very 

rapidly. In fact, we detected no significant change in the 

form of the zone from its earliest observable moments 

(fig. 10). A high degree of spatial clustering of aftershocks, 

particularly within the main zone and nearby areas northeast of the San Andreas fault, is a primary feature of the 

distribution. Many of these clusters represent individual 

faults and have been studied in detail by Seeber (written 

commun., 1992). These clusters commonly produced their 

largest locatable event very early in the sequence and thus 

may be related to short-wavelength variations in main 

shock displacements such as at the ends of the rupture. 

The rapid spatial variations in the style and orientation 

of aftershock focal mechanisms between the interior clusters, such as changes over short distances from fault-normal 

compression to fault-normal extension (fig. 25), further 

suggest a causal association with localized stresses created by the main shock. The general lack of correlation of 

the aftershock clusters with the main shock slip zones 

(see Spudich, 1996) thus supports Oppenheimer's assertion (1990) that abrupt spatial variations in main shock 

displacement is required to account for the observed short 

distance mechanism variability, particularly within the interior of the main zone. 

THE EARTHQUAKE AND THE SAN ANDREAS 

FAULT 

Finally, we examine the question of the relationship 

between the Loma Prieta rupture and the San Andreas 

fault in the southern Santa Cruz Mountains. The geologic 

investigations of Prentice and Schwartz (1991) and detailed trenching studies by Schwartz and others (1991) 

leave little doubt that the 1906 earthquake ruptured the 

main trace of the San Andreas fault adjacent to the 1989 

earthquake. Furthermore, both geodetic and strong-motion 

analyses of the Loma Prieta main shock (Spudich, 1996) 

require a large component of right-lateral strike-slip motion in the earthquake to the southeast of the main shock 
D42 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

epicenter. However, the upper extent of this slip did not 

reach the surface. We also know that the main zone of 

aftershocks merges seamlessly with the San Andreas fault 

near Pajaro Gap, virtually at the southern terminus of the 

1906 fault break. 

If the main zone of aftershocks faithfully maps the main 

shock rupture, as we hypothesize, then the association 

between it and the San Andreas fault could have several 

resolutions. In one, the Loma Prieta fault truncates a vertical San Andreas fault at -9 km depth. Alternatively, 

Olson (1990) suggests that the San Andreas truncates the 

Loma Prieta rupture plane. A third possibility is that they 

are one in the same, with the San Andreas fault following 

the main zone below 10 km and steepening to vertical 

above that depth. 

Neither the aftershocks nor the geodetic models of the 

earthquake (Lisowski and others, 1990; Snay and others, 

199 1; Marshall and others, 1991 ; Horton and others, 1996) 

imply that main shock faulting continued to the northeast 

of the San Andreas surface trace. The main zone, like the 

geodetic models, terminates updip below the trace of the 

San Andreas fault (sections B2 through F2, fig. 22). Thus, 

evidence for the San Andreas being cut by the Loma Prieta 

fault is lacking. The opposite case of the San Andreas 

truncating the Loma Prieta fault is at least consistent with 

these data. Extending the San Andreas to depths below 

the hypothesized truncation point at -9 km depth, however, has little to recommend it. A deep extension of the 

fault might be argued for on the basis of a single earthquake with a right-lateral strike-slip focal mechanism seen 

before the earthquake (fig. 22 Dl, deepest event). Regarding the third possibility, slip on both the deeper, inclined part of the fault in the Loma Prieta earthquake and 

on the vertical part in the 1906 earthquake is kinematically possible only along the southeastern part of the 1989 

rupture. North of the epicenter, the 1989 slip vector for 

the inclined part of the fault (predominately reverse), and 

the 1906 slip vector for the shallower vertical part (horizontal) are incompatible for a single structure and cannot 

be sustained in ongoing movement. 

Indeed, a common shortcoming of all three of these 

possibilities is that they fail to address the kinematics of 

the crust traversing the prominent left (compressional) bend 

in the San Andreas fault within the southern Santa Cruz 

Mountains. Over geologic time, flow of the crust around 

the bend must be accommodated by either lateral flow, 

subduction or mountain-building (Anderson, 1990). The 

simple kinematic model of Dietz and Ellsworth (1990) for 

the present-day geometry successfully predicts the ratio 

of horizontal to vertical slip on the Loma Prieta fault (fig. 

27). While this model is not unique, it suggests that we 

are observing a rapidly evolving fault geometry which 

must continually adjust as new crust enters the compressional bend. 

The smooth connection between the Loma Prieta rupture and the San Andreas fault to the southeast of the 

bend raises the possibility that the Loma Prieta earthquake followed what was once the main strand of the 

fault, but which is now in the process of being abandoned. Along the southern half of the earthquake where 

motion was predominately strike slip, the strand continues to carry some of the San Andreas motion. Farther to 

the north, where motion was dominantly reverse, it has 

been abandoned as a transform fault, but it still released 

the compressional component of the kinematically prescribed slip, with the translational part presumably being 

accommodated on the San Andreas. 

Ultimately, the resolution of this problem must await a 

more complete understanding of the faults within the southern Santa Cruz Mountains. The complete absence of seismicity on the fault that ruptured in this Ms=7. 1 earthquake 

over at least the 20 years preceding the event and the 

absence of an obvious imprint left by its rupture in the 

geologic record should remind us that other large active 

faults, including the San Andreas itself, may be "hidden" 

within the southern Santa Cruz Mountains and will only 

be fully revealed when they slip in major earthquakes. 

ACKNOWLEDGEMENTS 

We thank L. Jones and R. Page for their helpful reviews of this manuscript. 

9 = change in fault strike 

s = slip vector on inclined plane 

H = horizontal slip 

D = dip-slip motion 

Figure 27.-Simple block model to describe the slip vector (5) on a 

dipping fault segment within a bend of a vertical strike-slip fault. This 

model successfully predicts the ratio of horizontal (H) to vertical (D) 

slip observed on the Loma Prieta fault. 
AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D43 

Table 6.-Event locations and nodal planes from focal mechanisms in figure 20 and appendix A 

EQ Origin Time Latitude Longitude Depth M, Nodal Plane 1 Nodal Plane 2 

yrmoda hrrnn N W krn dip-azm dip rake dip-azm dip rake 

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D44 

Table 6.-Continued 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

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yrrnoda hrrnn N W krn dip-azm dip rake dip-azm dip rake 

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AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Appendix A.-Focal mechanisms (lower hemisphere projections) from figure 20 showing compressional (+) and dilatational (open circle) 

first arrivals. Lettered events locate in regions LPN, LPC, LPS, and CSA while numbered events locate in SAR, SCV, and MON. Event 

locations and nodal planes are listed in table 6. 
AFTEiRSHOCKS OF THE LOMA PRIETA EARTHQUAKE AND THEIR TECTONIC IMPLICATIONS D47 

Appendix A.-Continued 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKE OCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS 

CHANGE PRODUCED BY THE LOMA PRIETA EARTHQUAKE 

CONTENTS 

By Paul A. Reasenberg and Robert W. Simpson, 

U.S. Geological Survey 

Abstract ............................................................... Introduction ............................................................... 

Regional stress changes caused by the Lorna Prieta earthquake 

Regional seismicity changes after the Lorna Prieta earthquake 

Earthquake data .................................................... 

Quantifying seismicity rate changes ........................... Pattern of seismicity rate changes .............................. 

Comparison of regional stress and seismicity patterns ----------- 

Temporal changes in seismicity caused by the Lorna Prieta 

earthqu& ....................................................... 

Discussion and c~~~~us~o~~ ............................................ 

References cited ......................................................... 

Page 

D49 

49 

50 

5 3 

53 

53 

59 

59 

64 

67 

70 

ABSTRACT 

The Loma Prieta earthquake perturbed the static stress 

field over a large area of central California and greatly 

altered the levels of regional microseismicity. We have 

calculated the static stress changes produced on central 

California faults using elastic dislocation models of the 

Loma Prieta rupture. We used a Coulomb failure function 

(CFF) to describe the proximity of these faults to failure, 

and compared the calculated changes in CFF after the 

earthquake to observed changes in microseismicity rates 

on these faults. The pattern of stress changes on the major 

faults in the region that were predicted by our models of 

the earthquake agree closely with the pattern of changes 

in the regional seismicity rate observed after the earthquake. The agreement between stress change and seismicity change was detected at epicentral distances of up to 

100 km. The agreement is best for models with low assumed values of apparent coefficient of friction 

(0.15p'50.3) on San Francisco Bay area faults. Both the 

stress models and the microseismicity observations suggest that stresses were increased on the San Francisco 

peninsula segments of the San Andreas fault northwest of 

the Loma Prieta rupture and on the creeping segments of 

the San Andreas fault southeast of the rupture. These results are consistent with an increase in the long-term probability of a large earthquake on the San Francisco peninsula segments of the San Andreas fault and with geodetic 

and surface creep-rate observations suggesting accelerated 

slip on the creeping section of the San Andreas fault southeast of the Loma Prieta earthquake. 

The microseismicity observations suggest that the 

Hayward fault relaxed slightly after the Loma Prieta earthquake. The onset of the observed decrease in seismicity 

rate on the Hayward fault is poorly resolved, however, 

because the overall seismicity rate on the Hayward fault 

is low. In our stress models, the Hayward fault relaxes 

after the Loma Prieta earthquake because of the addition 

of small amounts of left-lateral shear to this right-lateral 

fault. This relaxation is small (a few bars or tenths of 

bars) and may already have been eroded by continuing 

tectonic loading. In our models, this relaxation could be 

countered at the southern end of the Hayward fault by the 

calculated increase in extensional normal stress that would 

tend to unclamp this part of the fault if friction is important. In one scenario, the inferred low values of apparent 

coefficient of friction and calculated stress changes could 

be the result of pore-fluid pressure changes in the fault 

zone that would work to diminish the effective normal 

stress changes initially. With time, as the pore fluids have 

a chance to re-equilibrate, these normal stress changes 

may become more effective, and the south end of the 

Hayward fault could be brought closer to failure. This 

scenario does provide a mechanism for the cross-bay triggering of a large earthquake on the Hayward fault. However, given the small magnitude of the calculated stress 

changes, such an earthquake on the Hayward fault would 

need to be nearly ready to happen in any event, and the 

changes in the already high long-term probability of a 

large earthquake on the Hayward fault have probably not 

been altered greatly. 

INTRODUCTION 

The Loma Prieta earthquake (Mz7.1) was the largest 

earthquake to strike the Bay Area since the 1906 San 
D50 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Francisco (M=8.3) earthquake (U.S. Geological Survey 

Staff, 1990). With an average slip of 2 meters over a 40- 

km-long rupture surface, the Loma Prieta earthquake profoundly altered the stresses in the crust throughout the 

Bay Area. An immediate concern after the earthquake was 

the possible effect this stress perturbation might have on 

other major faults in the Bay Area. Seventeen days after 

the Loma Prieta main shock, a magnitude 3.7 earthquake 

occurred on the Hayward fault about 85 km from the Loma 

Prieta epicenter. Although this was not a large event, it 

did cause some concern because on at least two occasions 

in the 1800's a large earthquake on one side of San Francisco Bay had been followed within 3 years by a second 

large earthquake on the opposite side of the bay (for example, Louderback, 1947, Ellsworth, 1990, Sykes and 

Jaum6, 1990)~~ Since earthquakes on the San Andreas and 

Hayward faults may have been involved in both of these 

pairings, there was a possibility that this magnitude 3.7 

earthquake might be a precursor to a larger shock, and the 

question arose as to how the Loma Prieta rupture might 

have affected stress levels on other Bay Area faults. Could 

stress changes produced by the Loma Prieta earthquake 

trigger another large earthquake east of San Francisco Bay 

or on the San Francisco peninsula? 

A working group convened by the U.S. Geological Survey considered in detail the effect of an increase in rightlateral shear stress on segments of the San Andreas fault 

north of the Loma Prieta earthquake rupture (Working 

Group, 1990). They concluded that the effect was to 

shorten the time to the next earthquake on these segments. 

Because this result was based largely on theoretical models and assumptions about fault behavior, it raised the 

following additional questions: Were these models and 

assumptions realistic and consistent with seismicity and 

other observations? What was the effect of the Loma Prieta 

earthquake on the Hayward fault? Did a scenario of crossbay triggering of earthquakes seem plausible? 

The location of the Lorna Prieta earthquake within a 

complex section of the San Andreas fault system in central California and in a region monitored by a dense seismograph network was fortuitous. The network provided 

an enormous quantity of high-quality observations of the 

aftershock sequence from an area that includes several 

major active faults in the San Francisco Bay and Monterey 

suddenly changed, most by increasing. The changes in 

seismicity provided a natural, in situ indication of the 

stress changes acting at seismogenic depths on these faults. 

Recently, other studies have suggested that changes in 

the regional stress produced by an earthquake may affect 

the hypocentral locations or focal mechanisms of its aftershocks. These include Stein and Lisowski (1983), Mavko 

and others (1985), Poley and others (1987), Oppenheimer 

and others (1988), Simpson and others (1988), Michael 

and others (1990), Rydelek and Sacks (1990), Seeber and 

Armbruster (1990), Michael (1991), and Reasenberg and 

Simpson (1992). Still others have suggested, through similar analyses, the possibility that stress changes produced 

by one or a series of earthquakes may be capable of triggering a comparable or larger earthquake on a neighboring fault (Hudnut and others, 1989, Sykes and Jaurnk, 

1990, Sykes and Jaum6, 1992, Harris and Simpson, 1992, 

Stein and others, 1992). 

In this paper we examine the relationship between the 

regional changes in static stress produced by the Loma 

Prieta earthquake and the changes in seismic activity on 

the nearby faults in central California. To estimate the 

stress changes, we calculated the stresses produced within 

an elastic halfspace by slip on a set of model dislocation 

surfaces. (See Simpson and Reasenberg (1994) for a detailed description of the stress modeling procedures and 

results for this region using various possible fault structures and source models.) We examine the spatial pattern 

of the seismicity change and its relation to patterns of 

stress change predicted by our fault and dislocation models. Because several different faults are involved, we find 

that the .seismicity observations are capable of partially 

constraining the stress and faulting models. Then, we examine the implications of the favored stress models for 

the potential for triggering another large earthquake in the 

Bay Area. Finally, we consider the aftershock-like temporal characteristics of the seismicity after Loma Prieta and 

find considerable variation among the many faults in the 

region. 

REGIONAL STRESS CHANGES CAUSED 

BY THE LOMA PRIETA EARTHQUAKE 

Bay regi0nse These currentl~ loaded stresses A detailed discussion of models used to calculate stress 

associated with the relative motion of the North American changes on other faults after the prieta earthquake and pacific plates, and are cOnstantl~ producing is given in Simpson and Reasenberg (1994). For corn- micrOearthquakese when the prieta earthquake OC- pleteness, we will briefly present the method of calcula- curred, the level seismic it^ On these tion, the slip model used, and a definition of the coulomb 

failure function (CFF). 

l~he 1836 (Mz6.8) earthquake on the northern Hayward fault was our calculations were based on dislocation theory in an 

f0110~ed in 1838 by an (Mz7.0) earthquake on the sm Francisco elastic half-space (Erickson 1986, Okada, 1992). bentral 

Peninsula section of the San Andreas fault, and the 1865 (Mz6.5) ca1ifornia faults were represented by alignments of verti- earthquake, possibly on the Lorna Prieta segment of the San Andreas 

fault, was followed in 1868 by m (Mz7.0) eaahquake on the s~utt,crn cal (and m me times dipping) rectangular dislocation surHayward fault. faces under the traces of major known faults (fig. 1). The 
HORIZONTAL SHEAR STRESS, IN BARS 

Rgure l.-Change in horizontal shear stress on each fault segment, calculated using Lisowski and others' (1990) model of the Lorna Frieta rupture. Red colors indicate increased fight-laterd 

shear; blue colors indicate decreased right-lateral shear. 
D52 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

vertical rectangles were 13 km tall and approximately 10 

km long, and the stress changes were calculated at their 

centers, which lie at about the mean depth of seismicity in 

the Bay Area. Stress tensors were resolved into three components relative to the rectangular fault patches: a horizontal shear component, an up-dip shear component, and 

a normal component perpendicular to the surface. 

To calculate the stress changes described in this paper, 

we used a slip model for the Loma Prieta rupture inferred 

by Lisowski and others (1990) from geodetic observations. Their slip model is a single, uniformly slipping, 

rectangular dislocation (shown in yellow in fig. 1) striking at N44OW, dipping 70Â°SW and extending 37 km horizontally. It has its top at a depth of 5 km and its bottom at 

a depth of 17.5 km. Geodetically inferred right-lateral slip 

of 1-66 m and reverse slip of 1-19 m yield a moment of 

3.0~10~~ Nm. More complex models of the rupture slip 

distribution were also explored but do not substantially 

change the results of the analysis presented here. 

In order to measure the change in the proximity to failure caused by changes in the shear and normal stress 

components, we defined a CFF based on the CoulombAmmonton failure criterion (for example, Jaeger and Cook, 

1979): 

CFF = + p[oP + p) - S, (1) 

where Itp[ is the magnitude of the shear stress acting on the 

surface, op is the normal stress acting on the surface (positive for extension), p is the pore-fluid pressure in the medium, p is the coefficient of friction, and S is the cohesion. 

If we assume that the cohesion S and the coefficient of 

friction p do not change before and after the Loma Prieta 

rupture, then the change in CFF becomes 

quantity of interest. Certain dipping fault patches discussed 

below have had ACFF calculated using assumed rake directions. 

The quantity (Aop+Ap), which is the change in normal 

stress modulated by the change in pore-fluid pressure, is 

also referred to as the change in the effective normal stress. 

We cannot calculate Ap without some additional assumptions. If the medium is homogeneous, then immediately after the earthquake (in the "undrained state" before 

the pore fluids have had time to re-equilibrate) the change 

in pore-fluid pressure can be represented as (Rice and 

Cleary, 1976) 

where B is Skempton's coefficient with value between 0 

and 1, so that 

ACFF = AT + p[AoP - BAokk 13). (4) 

If fault zones are assumed to be formed of more compliant material, then Rice (1992) has shown that c~~=o~~=o

in the fault zone, so that 

and 

ACFF = AT + p(1- B)Aop (6) 

For our analysis, we simplify the above possibilities by 

defining an apparent coefficient of friction p' and writing 

ACFF as 

ACFF = AT + ptAoP. (7) 

ACFF = Altp 1 + p[AoP + Ap). (2) 

For the homogeneous case, 

If the ambient shear stress acting on a surface is considerably larger than the increment added by the Loma Prieta 

earthquake, then the quantity ~17~1 can be approximated 

by the component of Loma Prieta induced shear, denoted 

in the following as just AT, in the direction of the ambient shear field. Armbruster and Seeber (1991) have used a 

similar quantity in their calculation of encouraging and 

discouraging stresses. Their assumption is that faults will 

generally slip in the same directions in which they have 

slipped previously, so that the appropriate component of 

shear stress to consider is the one in the direction of the 

rakes of past earthquakes. Because most of the vertical 

faults in the Bay Area are thought to be strike-slip faults, 

we have used the change in horizontal shear stress as the 

and for the compliant fault zone case of Rice, 

p' = p(1- B). (9) 

Values of B ranging from 0.23 to 0.99 have been reported 

for various rock types (Rice and Cleary, 1976, Roeloffs, 

1988). A value of BzO.73 and a value of p=0.75 close to 

values observed in the laboratory yield an apparent coefficient of friction pt=O.2, consistent with correlations described below. 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE D53 

We explored correlations between ACFF and seismicity rate changes as a function of p' by using equation 7.* 

Figures 1 and 2 show changes in shear stress and normal 

stress on the model fault segments. The normal stress 

changes will be most effective for high values of apparent 

coefficient of friction. Figure 3 shows the change in CFF 

for an assumed value of pt=0.2. Additional plots for other 

values of p' are shown in the paper by Simpson and 

Reasenberg (1 994). 

REGIONAL SEISMICITY CHANGES 

AFTER THE LOMA PRIETA 

EARTHQUAKE 

EARTHQUAKE DATA 

To accurately measure the fluctuations in seismicity rate 

in the San Francisco Bay region we must first account for 

possible artificial rate changes in the catalog. Such artifacts may be produced by changes in the detection capability of the seismograph network or by changes in the 

way earthquake magnitudes are estimated. While significant growth in the network during the 1970's increased 

its sensitivity to the smallest earthquakes, the detection 

level in the study area since 1980 for earthquakes with 

magnitude 1.5 and larger has remained relatively constant. 

Owing to the exponential distribution of earthquake 

magnitudes (Richter, 1958), the seismicity rate estimated 

from a magnitude-truncated sample is very sensitive to a 

systematic error in magnitude. For seismicity with a b 

value of 1, a systematic reduction of 0.1 magnitude units 

produces approximately a 2 1 -percent apparent reduction 

in rate. A change in the network occurred in May 1977 

when the amplification of certain elements in the Calnet 

network had been reduced. This change is believed to 

have introduced an artificial shift in the calculated magnitudes of earthquakes, which were determined from measurements of the coda waves that did not account for the 

amplification change. Systematic magnitude decreases of 

between 0.1 and 0.3 units, relative to the magnitudes listed 

in the Berkeley catalog, were detected by Habermann and 

Craig (1988) at around this time on the southernmost sec2~e also attempted to explore the degree of correlation between 

ACFF as defined by equation 4 and seismicity rate changes. In this 

attempt the degree of correlation was a function of both p and B. The 

results were inconclusive, suggesting that either the data did not have 

enough resolving power to yield statistically significant values for both 

p and B (which could happen because of the approximate parallelism 

of most Bay Area faults), or the hypothesis being tested in the form of 

equation 4, with constant values of p and B everywhere, was not 

correct. Indeed, if equation 9 holds, then p and B cannot be determined 

separately. 

tion of the Calaveras fault, the Sargent fault, and the San 

Juan Bautista section of the San Andreas fault. Our measurements show that the average rate of earthquakes located on these fault sections and listed in the catalog with 

magnitude M21.5 decreased 44 percent. This decrease in 

rate is equivalent to an artificial magnitude shift of 0.25 

units when the b-value is 1, and is consistent with 

Habermann and Craig's estimate. No change in network 

operation affecting magnitude determination is known to 

have occurred since 1977. 

To avoid these potential problems of stability with magnitude estimation and network sensitivity, we shall focus 

on the M2l.5 earthquakes during the 10-year period 1979 

to 1989 for our comparisons with the post-earthquake activity. These data represent the best years of the USGS 

network and provide a stable yardstick for assessing the 

Loma Prieta-induced changes in the regional seismicity. 

On the whole, the pattern of earthquake activity after 

the earthquake resembles the pattern for the 10 years of 

seismicity before the earthquake (fig. 4). The major faults 

are clearly visible in both patterns as narrow alignments 

of epicenters. Other features that are clearly apparent in 

the 10-year pattern are also apparent, though perhaps not 

so obviously, in the shorter-duration postseismic period. 

Some areas appear to have reduced activity after the Loma 

Prieta earthquake, but it is difficult to assess this from 

figure 4 because the time intervals represented in the figure are unequal. In the next section we try to quantify 

these comparisons. 

QUANTIFYING SEISMICITY RATE CHANGES 

Because earthquakes occur at discrete locations and 

times, the concept of a continuously varying "level" of 

seismicity that parallels our concept of the stress field 

presents some problems of definition, especially in regions where seismicity is sparse. To compare the seismicity to the regional coseismic stress change we have 

calculated, we sought a quantitative measure of the seismic activity that would be sensitive to coseismic changes 

while remaining free from assumptions about the spatial 

nature and extent of the changes. We compared the average rate of earthquakes in small, overlapping subregions 

during two specified time intervals-one before and the 

other after the earthquake. Using the time of the Loma 

Prieta earthquake to define the intervals assures that the 

result will be sensitive to the average or DC changes occurring at that time. By finely subdividing the region with 

a uniform grid, we avoid imposing artificial spatial constraints, as would be introduced by the selection of a few 

arbitrary "boxes" or subregions, and we allow the spatial 

pattern of the seismicity change to reveal itself naturally. 

We measured the regional coseismic change in seismicity rate using the statistic (3, which is sensitive to a 
NORMAL STRESS, IN BARS 

Figure 2.Ã‘Chang in normal stress on each fault segment, calculated using Lisowski and others' (1990) model of the Lorna Prieta rupture. Red colors indicate increased normal extensional 

stress; blue colors indicate increased normal compressional stress. 
-5 -4 -3 -2 -1 0 

RELAXED, BLUE 

12345 

- LOADED, RED 

Segments with increased CFF (red) are more likely to produce earthquakes; segments with 

decreased CFF (blue), less likely. This oblique view is from the southwest and down at an 

angle of 45O. Inset map shows study area. Segment numbers are shown for reference in fig. 9. 

Figure 3.Ã‘Chang in Coulomb failure function (CFF) produced by the Loma Prieta earthquake on rectangular segments representing faults in central California, calculated using 

Lisowski and others' (1990) model for the Lorna Prieta rupture (yellow) derived from geodetic measurrnents. ACFF (in bars) on each fault segment for p'=0.2 is indicated by color. 
' D56 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

contrast of average seismicity rate between two time intervals in a specified area (Matthews and Reasenberg, 

1988). We compared the rate r in the postseismic interval of duration ta with the rate rb in the preseismic interval of duration tb, where ra=na/ta, rbznb/ti, and n, nb are 

the numbers of earthquakes occurring in the respective 

intervals. The rate change is expressed as 

(nap "b ' ) - 

"'I - Â£(" ) 

'Ib Jwr(i!,)' 

where var denotes variance and E(na)=rbta is the value of 

nu expected under the null hypothesis of stationary random occurrence. The variance was represented by that of 

a binomial process: ~ar(n~)=n~~. 

We calculated 0 for fixed ta and tb in overlapping 10- 

km-square cells covering an area 140 by 390 krn for the 

preseismic (background) period between 17 October 1979 

and 17 October 1989 and postseismic period between 18 

October 1989 and 3 1 May 1991 (fig. 5). Gaussian spatial 

smoothing with halfwidth 5 km was applied to the data in 

each cell to minimize artifacts introduced by the square 

gridding. Positive values of 0 indicate that the postseismic 

rate was higher than the background rate; negative values, 

lower than the background rate. Significance levels for 101 

estimated from its asymptotic (Gaussian) distribution are 

1.96 (^=0.05) and 2.57 (^=0.01). 

Figure 4.-Earthquakes (ml.5) located in the study area 10 years before and 20 months after the Loma Prieta earthquake. 

(A) Earthquakes occurring between 18 October 1979 and 17 October 1989. Some 21,000 events are represented. (B) 

Earthquakes occurring between 18 October 1989 and 31 May 1991. Approximately 5,000 events are represented. 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE D57 

An asymmetry in the definition of fl introduces an ambiguity in its interpretation: a relatively low postseismic 

rate cannot be distinguished from an abnormally high 

preseismic rate, as would result, for example, from an 

earthquake swarm or aftershock sequence in the preseismic 

(background) period. This limitation stems from our use 

of the empirical background seismicity level as a reference level and presents an inherent difficulty in discerning rate changes in a finite sample of a point process. 

In our measurements of San Francisco Bay region seismicity, several artifacts corresponding to the aftershocks 

of moderate earthquakes in the pre-Loma Prieta background period are apparent (fig. 5). These artifacts appear 

as dark blue patches and correspond to the 1979 Coyote 

Lake (M=5.9), 1980 Livermore (M=5.9), 1983 Coalinga 

(M=6.7), 1984 Morgan Hill (M=6.2), 1985 Kettleman Hills 

(M=5.5), 1986 Mt. Lewis (M=5.7), 1986 Quien Sabe 

(M=5.7) and 1988 Alum Rock (M=5.1) (Reasenberg and 

Ellsworth, 1982, Oppenheimer and others, 1990, 

Oppenheimer and others, 1988, Baton, 1990). 

We tried varying the choice of background period to 

avoid these artifacts. When the background is taken as 

1969-1979, many of the aftershock zones that were blue 

in figure 5 became red (fig. 6), indicating that these aftershock sequences may not have fully decayed at the time 

of the Loma Prieta earthquake. New artifacts, including 

those associated with earthquake sequences near Danville 

in 1972, Bear Valley in 1972, and on the Busch fault in 

1974, were introduced (Oppenheimer and others, 1990, 

Ellsworth, 1975, Lee and others, 197 1). 

We tried to reduce these artifacts by removing aftershocks from the catalog with a computer algorithm 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

(Reasenberg, 1985). Removal of aftershocks apparently postseismic rate (-3Spc0) remaining along the Calaveras 

eliminated some of the artifacts (for example, the 1983 and Mt. Lewis faults in figure 7 may reflect incomplete 

Coalinga aftershock sequence), but may have only par- removal of aftershocks, and this possibility hampers intertially eliminated others (fig. 7). The low apparent pretation of these features as postseismic effects. 

BETA 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE D59 

PATTERN OF SEISMICITY RATE CHANGES 

Overall, the level of regional seismic activity after the 

Loma Prieta earthquake increased. The pattern of the increase was highly heterogeneous, with intense postseismic 

activity occurring on some parts of some faults, and little 

change or a decrease at comparable distance on others. 

Areas of high seismicity rate after the earthquake include 

the immediate aftershock zone (from Los Gatos to 

Watsonville); the San Francisco peninsula in an area of 

thrust faults near Los Gatos and along the San Andreas 

and San Gregorio faults near Daly City; along the San 

Andreas and Sargent faults from Watsonville to Bear Valley; a partially off-shore segment of the San Gregorio 

fault near Pt. Ano Nuevo; and along the Alamo fault, near 

Livermore. With the exception of the activity near Pt. 

Ano Nuevo, these areas generally coincide with fault segments having increased CFF. At Pt. Ano Nuevo, segments 129, 130, and 13 1 experienced sizable seismicity 

rate increases and calculated decreases in CFF(p') for all 

values of apparent friction. 

Increases in seismicity relative to background activity 

levels were greatest close to the Loma Prieta rupture, along 

the San Andreas and Sargent faults (fig. 8). The apparent 

"S" shape of the area of most intense increase reflects 

relatively deep activity (below 10 km depth) along the 

center and northwest parts of the rupture and shallower 

activity on the southeast part of the rupture and on the 

Sargent fault (see Dietz and Ellsworth, this volume, for 

detailed views of the spatial distribution of the aftershocks). 

At somewhat greater distances, the apparent disagreement between the seismicity and calculated stress changes 

near Pt. Ano Nuevo suggested to us that our model for the 

San Gregorio fault may be inaccurate and prompted us to 

explore alternate fault models. Because this seismic activity is largely offshore, the ability of the network observations to constrain the fault plane solutions for these 

earthquakes is somewhat limited. Nevertheless, the set of 

focal mechanisms obtained for 133 earthquakes occurring 

between 1984 and 1991 shows a trend of oblique rightlateral thrust solutions (median rake of 120' aligned 

subparallel to the San Gregorio fault (with approximate 

strike N20Â°W and dipping northeast between 50' and 

80'. As illustrated by the dipping segment (200) adjacent 

to segment 130, such thrust mechanisms would be favored in this vicinity (fig. 3). R. Jachens and A. Griscom 

(written commun., 1992) report that aeromagnetic data 

suggest a 50-60 degree northeastward dip for the onland 

part of the San Gregorio fault and its offshore extension 

to the south. 

Areas that experienced an apparent decrease in seismicity rate after the Loma Prieta earthquake include the 

southern Calaveras, Hayward, southern Rodgers Creek, 

and Mt. Lewis faults. These fault segments also have relaxed stress (ACFF<O). As discussed in the previous section, however, the effects of other strong aftershock 

sequences before 1989 preclude unambiguous interpretation of the negative values of (3 obtained along the southern Calaveras and Mt. Lewis faults in terms of postseismic 

effects. The Hayward fault, however, was essentially free 

of significant swarm and aftershock activity during the 

1980's (compare fig. 5 and fig. 6). Consequently, the apparent low postseismic rate observed there (strongest on 

the southern Hayward fault), is not believed to be an artifact, and may be related to the Loma Prieta earthquake. 

This low postseismic rate is also consistent with an observed slowdown in creep rates on the Hayward fault reported by Galehouse (this chapter) and by Lienkaemper 

and others (1989, 1991). 

Figure 5.-Changes in seismicity rate statistic p in central California occurring at the time of the Loma Prieta earthquake. Values of p 

(color) compare average rates in the intervals 17 October 1979 to 17 

October 1989 and 18 October 1989 to 31 May 1991. Areas that experienced a co-seismic increase or decrease in seismicity rate are shown as 

red or blue, respectively; areas with insufficient earthquakes for determination of p are white. Absolute value of 13 indicates the statistical 

significance of the rate change, taking into account the variance of the 

background seismicity. Major faults and coastline are indicated by solid 

lines. Main shock rupture and fault segments 198 and 199 appear as 

rectangles due to projection of dipping planes onto surface (see fig. 3). 

The large bright red area includes the Loma Prieta aftershock zone. 

Values of p in the area close to the rupture greatly exceed 5, but are 

truncated by color scale (see fig. 8). Bright blue patches along and 

northeast of the Calaveras fault are artifacts corresponding to strong 

aftershock sequences that occurred during the background interval, and 

do not indicate Loma Prieta-related postseismic effects. A, 1979 (M=5.9) 

Coyote Lake earthquake; B, 1980 (M=5.9) Livermore earthquake; C, 

1983 (M=6.7) Coalinga earthquake; D, 1984 (M=6.2) Morgan Hill earthquake; E, 1985 (M=5.5) Kettleman Hills earthquake; F, 1986 (M=5.7) 

Mt. Lewis earthquake; G, 1986 (M=5.7) Quien Sabe earthquake; H, 

1988 (M=5.1) Alum Rock earthquake. 

COMPARISON OF REGIONAL STRESS 

AND SEISMICITY PATTERNS 

To investigate the overall dependence between the modeled stress changes and the observed seismicity, we compared, for each fault segment Sj the mean seismicity 

rate change index % corresponding to the subset of cells 

within 5 km of segment Sj to the coseismic static stress 

change ACFFj (p')calculated on that segment, for various 

assumed values of the apparent coefficient of friction p'. 

We used values of (3 obtained with aftershocks removed 

for the background period 1979-1989. We show one of 

the better fitting comparisons (for p'=0.2) in figure 9. For 

this value of the apparent coefficient of friction, concordant changes in stress and seismicity (both increase or 

both decrease) occurred on 87 fault segments, discordant 

changes occurred on 54 segments, and a generally 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

BETA 

Figure 6.Ã‘Change in seismicity rate, represented as in figure 5, except that the preseismic interval is 

chosen here to be between 18 October 1969 and 17 October 1979. Switching to this earlier background 

interval avoids many energetic aftershock sequences in the 1980's. Apparently, however, many of 

these same areas remain at elevated seismicity levels (red patches), compared to the levels of the 

1970's, indicating that these aftershock sequences are not yet over. 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE 

BETA 

Figure 7.Ã‘Change in seismicity rate, represented as in figure 5 (with background period 1979 to 

1989) except that identified earthquake clusters have been removed from the catalog. While removing 

the clusters reduces the effects of aftershock sequences in the background period, some residual 

aftershock activity apparently remains, preventing unambiguous interpretation of the postseismic rates 

in these areas. The depressed rate along the southern part of the Hayward fault, however, is apparently 

not affected by such aftershock-related artifacts; it is blue in all three representations of the postseismic 

rate change (figs. 5 to 7). 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

positive correspondance between and ACFF is visually for example, Sachs, 1982) allows a comparison of four 

apparent. relative frequencies, fl, f2, f,, fs, each representing a subTo quantify the agreement between stress change and set of the fault segments in our model. In the table, fl is 

seismicity rate change we used the chi-squared statistic in the number of fault segments for which ACF-0 and p>0, 

conjunction with a fourfold table. The fourfold table (see, f4 is the number of fault segments for which ACFFe0 and 

Figure 8.Ã‘Change in seismicity rate in the Loma Prieta aftershock 

zone relative to the background rate, calculated as in figure 5 but with 

the full positive range of values of beta shown. Intense increases in 

seismicity (yellow) near the center and northwest end of the rupture are 

BETA 

relatively deep (10 to 18 krn depth). Intense aftershock activity is 

shallower on the Sargent fault and near the southeast end of the Lorna 

Prieta rupture, near the site of the 18 April 1990, M=5.2 Chitenden 

aftershock. 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE D63 

fk0, and f2 and f3 correspond to the numbers of segments 

with opposing (different sign) changes of stress and seismicity. High values of y2 correspond to high values of f1 

and f4, low values of f2 and f3 and positive correlation 

between the stress and seismicity rate change. Significance levels for y2 are 3.84 (p= 0.05), 6.64 (p=0.01), and 

10.83 (p=0.001). An advantage of this formulation of chisquared is that it makes no assumptions about the functional relationship betweeen the stress changes and 

seismicity changes, except that they be of the same sign 

to be in agreement. 

Application of the y2 test to all 141 fault segments represented in figure 9 rejects the null hypothesis that and 

ACFF(p') are independent (p<0.001 for 0.2$'S0.3; p<0.0 1 

for O.lSp'SO.4) (fig. 10A). The correlation coefficient p 

attains its maximum value for p'zO.2 and exceeds 0.5 for 

O.lSp'SO.3 (fig. 105). 

Many of the points plotted in figure 9 indicate absolute 

changes in stress or seismicity on individual segments 

comparable to the uncertainties associated with modeling 

errors (stress) and stochastic variance (seismicity). When 

the y2 and correlation tests were applied to a subset of the 

fault segments with significant changes in stress and seismicity (numbered points in fig. 9), similar results were 

obtained (fig. 10). y2: p< 0.01 for O.lSp'SO.4; correlation: 

p exceeds 0.5 for O.OSptS 0.3. 

Static stress changes as small as a few tenths of a bar 

apparently produced detectable changes in seismicity (fig. 

11). This level of stress change is about one order of 

magnitude larger than that of tidally induced stress changes 

(Heaton, 1982, McNutt and Heaton, 1981, Stacey, 1977) 

and is comparable to stress changes produced at 

seismogenic depths by the filling of water reservoirs 

(Roeloffs, 1988). 

The response of seismicity to static stress changes produced by the Loma Prieta earthquake decreases with distance from the rupture and is statistically undetectable (at 

the p=0.05 level) at distances beyond 80 to 100 km; at 

this distance the maximum absolute value of ACFF is 

approximately 0.1 bar. To determine this, we calculated 

y2, taking p'=0.2, for subsets of segments in various ranges 

of distance from the center of the earthquake dislocation. 

y2 is maximum (y^ll) for the 80 segments between 0 

and 100 km from the source; x2 just fails to exceed the 

p=0.05 critical point for the farthest 70 segments located 

80 or more km from the earthquake. For comparison of 

this distance range to that of other earthquakes, we calculate that static stress changes of 0.1 bar are produced by 

M=3 earthquakes at hypocentral distances of approximately 

0.4 km, and by M=5 earthquakes at distances of approximately 7 km on favorably oriented planes. Thus, small 

earthquakes close to the fault segment may have an equally 

-2 -1 0 1 2 3 4 

Change in Coulomb Failure Function (bar) 

Figure 9.-Seismicity change and stress change AClT(u/) calculated on 141 model fault 

segments, assuming p'=0.2. Numbers (refer to fig. 3) indicate segments experiencing the 

largest changes (I Bj I22 and 1ACFF)20.2 bar). 
D64 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

intense, but spatially limited, effect on seismicity rate as 

do large, distant earthquakes. 

In a speculative vein, the apparent sensitivity of the 

central California seismicity to small stress changes suggests that, in general, fluctuations in regional seismicity 

might also reveal regional, aseismic slip events such as 

earthquake afterslip, slow earthquakes, slip on the ductile 

portion of vertical faults, and slip on horizontal detachment surfaces. The feasibility of monitoring the seismicity to infer aseismic strain events obviously would depend 

on both the sensitivity of the seismic network and the 

overall level of regional seismicity, and it would certainly 

be an enormous challenge to future modelers to constrain 

the location and source parameters of such events solely 

from seismicity data. There are some suggestions of largely 

aseismic strain events in the year before Loma Prieta. 

Galehouse (this volume), in his creep data from the 

Hayward fault, finds evidence of slowdowns at various 

sites ranging from 0.3 to 3.6 years before the Loma Prieta 

earthquake. Gladwin and others (this volume) report the 

beginning of a shear strain change at a tensor strainmeter 

near San Juan Bautista in mid- 1988. 

TEMPORAL CHANGES IN SEISMICITY 

CAUSED BY THE LOMA PRIETA 

EARTHQUAKE 

Up to this point we have considered the change in regional seismicity as a coseismic step change in rate in 

order to facilitate comparison to our calculated step 

changes in stress. Having seen that the sense of change in 

seismicity and stress agree in many areas, we now focus 

on the time dependence of the seismicity rate in some of 

these areas. The time-dependent part of the response of 

seismicity to the regional stress change may help constrain present and future models of fault behavior and 

regional deformation. 

In this section we examine the rate of M>1.5 earthquakes in several geographical zones as a function of time. 

We defined the zones to be large enough to contain sufficient earthquakes for a statistically significant analysis, 

yet small enough to resolve areas that appear to have 

responded coherently to the Loma Prieta earthquake. The 

0.0 0.2 0.4 0.6 

Coefficient of Friction 

0.0 0.2 0.4 0.6 

Coefficient of Friction 

Figure 10.-Measures of agreement between stress changes and seismicity changes on fault segments, shown as a function of the apparent coefficient of friction 1' assumed for the stress calculations. 

Solid lines, all 141 fault segments; broken lines, fault segments experiencing significantly large changes (numbered points in fig. 9). 

(A) Chi-squared vs. friction. Chi-squared confidence levels are 6.64 

(p=0.01) and 10.83 (D=O.OOl). (B) Correlation vs. friction. 

J .- 

c 

en 0 50 100 150 200 250 

(0 

s Distance from Earthquake (km) 

w .- 

a 0 50 100 150 200 250 

cn Distance from Earthquake (km) 

Figure 1 1.-Absolute levels of stress change and seismicity rate 

change on individual segments, shown as a function of segment 

distance from the earthquake epicenter. (A) Calculated stress 

changes (for assumed value of apparent coefficient of friction 

p'=0.2). (B) Observed seismicity rate change statistic, B. Significant correlation between the stress changes and seismicity rate 

changes was observed at distances up to 100 km. 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE D65 

zones are defined in terms of subsets of the model fault 

segments shown in figure 3 and include all M>1.5 earthquakes within 5 krn of any segment in the subset (table 

1). There is a small amount of geographical overlap among 

the zones. For each zone, we calculated the rate of earthquakes in 30-day, non-overlapping intervals between January 1979 and March 1992 (fig. 12). The choice of 30-day 

intervals for our rate estimates was subjectively made to 

provide a balance between resolution and accuracy. 

Seismicity during the preseismic period in all zones 

appears to occur at more or less constant rates punctuated 

by bursts of activity. For example, the largest four 

bursts of seismicity along the creeping section of the 

San Andreas fault (Zones 2N and 2s) before the Loma 

Prieta earthquake are associated with moderate (M=4.4 

to M=4.7) earthquakes near Fremont Peak in April 

1980 and near Bear Valley in August and September 1982, 

May 1986 and July 1988 (fig. 12, table 2). The lower 

average preseismic rates in the other zones are similarly 

punctuated by smaller clusters involving smaller earthquakes. 

The change in seismicity rate after the Loma Prieta 

earthquake varied significantly from zone to zone (table 

1). For a rough measure of these differences we compared, for each zone, the change in seismicity rate between the 10-year preseismic and 20-month postseismic 

periods. Of course, the seismicity rates varied greatly during this postseismic period, so our results depend strongly 

on the choice of intervals, but they serve for comparison. 

Along the southern portion of the creeping section of the 

San Andreas fault (zone 2s) the relatively high background 

rate of 0.58 eventslday increased 50 percent after the Loma 

Prieta earthquake, while farther north and closer to the 

rupture (zone 2N) the rate increased sevenfold. Within 

the aftershock zone proper (zone 1) seismicity rate increased nearly twentyfold, and on the southern San Francisco peninsula (zone 3), seismicity rate increased 

threefold. 

We modeled the seismicity in each zone during the 

background period as a homogeneous (constant rate) Poisson process. While these models generally are not very 

well fit (in a least-squares sense) by the seismicity rate 

data due to the presence of bursts of aftershocks, they 

provide convenient and well-defined reference levels for 

comparison to the postseismic period. Seismicity rate immediately after the Loma Prieta earthquake in zones 1 to 

4 abruptly increased and then decayed in an aftershocklike manner. Accordingly, we modeled the seismicity rate 

for the postseismic period in zones 1 through 4 as a 

nonhomogeneous Poisson process with rate N(t) obeying 

the modified Omori's law (Utsu, 1971): 

K N (t) = - 

(t + c)' ' 

where t is time after the main shock, and K, c, and p are 

positive constants. We estimated these model parameters 

separately for the seismicity in each zone with a maximum likelihood method (Ogata, 1983) (fig. 13). These 

models generally provided excellent fits to the first 100 

days of data but were incapable of modeling the complexity in seismicity rate in zones 1 and 2N resulting from the 

strong secondary aftershock sequence that followed the 

M=5.6 aftershock near Chittenden on 18 April 1990. For 

these zones, we fit Omori models for the intervals 18 

October 1989 to 17 April 1990 (table 1); in zones 3 and 4, 

data between 18 October 1989 and 31 March 1992 were 

used. 

Our power-law models for the aftershock zone, the 

southern San Francisco peninsula, and the northern part 

of the creeping San Andreas fault all have p-values close 

to the mean value of p previously found for California 

aftershock sequences of 1.07M.2 (Â±lo (Reasenberg and 

Jones, 1989) (fig. 13, table 1). Our model for the decay of 

seismicity rate on the San Gregorio fault based on earthquakes mostly occurring near segments 131 to 134 is not 

well constrained, owing to the overall low level of the 

seismicity response there, and the absence of recorded 

aftershocks during the first day of the sequence. While 

the estimate for p there (0.62k0.15) is lower than the California average, these data are too sparse to justify interpreting this model. 

From the Omori models and observations of the long 

term background rates we estimated times at which the 

seismicity rate in each zone will return to the pre-Loma 

Prieta level. We defined the end of elevated-rate period as 

the time when the Omori model rate equals the 1979- 

1989 background rate. We found that in the aftershock 

zone (Zone 1) elevated activity will have ended between 

1991.3 and 1992.9. The activity on the northern creeping 

section of the San Andreas fault will have subsided to 

normal between 1990.4 to 1992.0, and activity on the San 

Francisco peninsula will have returned to normal between 

1990.5 and 1991.3 (table 1). Because the levels of both 

aftershock activity and background activity vary spatially, 

our models (and, consequently, these estimates) depend 

on the particular zones we have used. The zones group 

together shallow and deep earthquake activity and include 

multiple fault strands. As a consequence, our results represent subregional averages. In contrast, Dietz and 

Ellsworth (this chapter) have applied a similar analysis to 

the seismicity located within approximately 30 km from 

the main shock epicenter, using M2l earthquakes in 

smaller and more complexely defined zones, and find significant variations in temporal decay of the aftershock 

activity associated with specific fault structures. 

The post-Loma Prieta seismicity rate along the Hayward 

fault (zone 5) is the lowest among the regions we have 

considered, and the coseismic change there was a rate 
D66 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 1 .-Summary of seismicity in selected zones 

fine Region 

l1 I Aftershock Zone 

2N Creeping SA fault 

(Northern Section) 

2S Creeping SA fault 

(Southern Section) 

3 Southern S.F. Peninsula 

4 San Gregorio fault 

Rate (2) 

(Eventsmay) 

1. Earthquakes within 5 km of any of the fault segments were used. 

-- - 

Postseismic 

Rate (3) 

(Eventsmay) 

2. Mean rate in the 10-year interval 17 October 1979 to 17 October 1989. 

3. Mean rate in the 20-month interval 18 October 1989 to 31 March 1992. 

4. Model fit to the interval 18 October 1989 to 31 March 1992. 

5. Model fit to the interval 18 October 1989 to 17 April 1990. 

6. Mean rate in the interval 17 October 1979 to 1 January 1988. 

7. Mean rate in the interval 1 January 1988 to 31 March 1992. 

decrease, which is inherently more difficult to detect than 

a rate increase. As a result of these differences, the 

Hayward fault seismicity rate record is particularly difficult to interpret, and the changes we infer for the Hayward 

fault are less certain than those for the other zones. After 

the Loma Prieta earthquake, the seismicity rate on the 

Hayward fault is about half the rate during the 10-year 

background period (fig. 12, table 1). In a previous section 

we interpreted this reduction in rate as support for our 

stress models. However, the Hayward activity appears to 

have undergone its most prominent rate decrease around 

the beginning of 1988, after which time the rate averages 

0.07 eventslday and does not exceed 0.17 eventslday. This 

decrease, also on the order of 50 percent, coincides roughly 

with the times of observed decreases in right lateral creep 

rates between May 1988 and June 1989 at four of five 

alignment arrays situated along the Hayward fault 

(Galehouse, 1990, 1992; also, this chapter). While there is 

considerable uncertainty in the onset times of both our 

seismicity rate decrease and Galehouse's creep rate reductions, their approximate coincidence suggests a conK 

(Eventsmay) 

75 + 7 

nection and raises the possibility that a separate episode 

of relaxation on part of the Hayward fault may have begun in this period before the Loma Prieta earthquake. 

When compared to the relatively low-rate period starting 

in 1988, the post-Loma Prieta seismicity still shows a 

decrease, although the size of the decrease is naturally 

smaller than it is in comparison to the 10-year background 

period. A possible interpretation of the Hayward fault seismicity is that two relaxation events may have occurredone related to the 1988-1989 creep retardation and the 

other related to the Loma Prieta earthquake. 

During the period since the Loma Prieta earthquake, 

the seismicity rate on the Hayward fault has increased 

slightly (fig. 12). We attempted to model this increase as 

a linear rate change, suggestive of a constant-rate process 

of reloading by tectonic forces. We had hoped to be able 

to estimate the loading rate from the ratio of the average 

coseismic stress drop calculated for the Hayward fault to 

the estimated time at which the seismicity rate returns to 

its preseismic level. This was not possible, however, because of the large uncertainties in the regressed time of 

P 

0.90 k 0.02 

c 

(day) 

0.21 k 0.07 

Return to 

Preseismic 

Rate (Years) 

6.9 k 1.5 (4) 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE D67 

return to the preseismic rate (treturn=1994.7k112 years) 

that result from the low seismicity rates on the Hayward 

fault. 

DISCUSSION AND CONCLUSIONS 

The pattern of coseismic changes in seismicity rate in 

the San Francisco Bay region are positively correlated 

with the static stress changes calculated for central California faults at distances up to 100 km from the Loma 

Prieta rupture. The largest increases in seismicity were 

near the ends of the earthquake rupture on the San Francisco peninsula to the northwest and along the creeping 

section of the San Andreas fault to the southeast. The 

seismicity increase on the San Francscio peninsula is consistent with elastic dislocation models for the earthquake, 

which predict increases in the Coulomb failure function 

there and support a conclusion of the Working Group 

(1990) that the probability of a large earthquake on the 

San Francisco peninsula was increased by the Loma Prieta 

earthquake. 

The correlation between the observed seismicity rate 

changes and modeled stress changes is best for models 

involving low apparent friction (0.1<p':<0.3) on central 

California faults. In contrast, laboratory experiments on 

frictional slip in rock typically find high coefficients of 

friction (0.5SptS0.8) (Byerlee, 1978, Dieterich, 1981). Our 

result is consistent with the idea that low apparent friction 

is required to explain seismological and other evidence 

for fault-normal compression in central California (Zoback 

Zone 3 - Southern S.F. Peninsula 1 

Zone 4 - San Gregorio fault 

I I I I I I I 

80 82 84 86 88 90 92 

Year 

Figure 12.-Rate of MS1.5 earthquakes in selected zones (see table 1) 

between 1 January 1979 and 31 March 1992. Letter symbols "a" through 

"1" significant aftershock sequences in the region (see table 2). 

Rates were calculated in nonoverlapping, 30-day intervals. Average rates 

in the 10-year preseismic and 20-month postseismic periods are indicated by horizontal lines. In the postseismic interval in zones 1, 2N, 3 

and 4, modified Omori's law (power law) fits to the data are shown by 

smooth curves. In zones 1 and 2N models representing all data (dashed 

curve) and data up to the time of the M=5.6 aftershock on 18 April 1990 

(solid curve) are shown (See fig. 13). In zones 3 and 4 models representing all data are shown by solid curve. Along the Hayward fault (zone 5), 

an apparent decrease in rate occurs near the beginning of 1988. In addition, a decrease in seismicity rate at the time of the Lorna Prieta earthquake may be present. This decrease and the subsequent gradual increase 

in rate are modeled by a linear regression (solid increasing straight 

line). 
D68 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 2.Ã‘Significan regional aftershock sequences (1979-1992) in regions shown in figure 12 

1. Refer to Figure 12. 

Index (1) 

a 

b 

c 

d 

e 

f 

g 

h 

i 

j 

k 

1 

Time 

0452 

0004 

1354 

0616 

0747 

0848 

03 27 

0813 

0202 

1324 

1624 

0646 

Date 

80061 8 

891018 

90041 8 

8004 1 3 

82081 1 

86053 1 

880726 

900908 

790123 

820326 

860329 

87051 1 

Latitude 

- 

Longitude 

and others, 1987, and references therein; Oppenheimer 

and others, 1988) and the lack of an observable heat flow 

anomaly associated with the San Andreas fault 

(Lachenbruch and McGarr, 1990). Our low apparent-friction value can be explained by a high coefficient of friction (consistent with laboratory studies) combined with 

pore-pressure changes in the fault zone (Byerlee, 1990, 

Rice, 1992) induced by the Loma Prieta stress changes. 

Our stress models suggest that for low values of apparent coefficient of friction, most parts of the Hayward fault, 

especially the southern end, underwent a slight relaxation 

of the expected right-lateral shear load. While the observed decrease in seismicity rate on the Hayward fault is 

consistent with such a relaxation, we are unable to determine the precise onset time of the low-rate period; it may 

have begun before the Loma Prieta earthquake, possibly 

as early as the beginning of 1988. Measured surface-creep 

rates at four of five sites along the Hayward fault 

(Galehouse, 1992) apparently began to decrease during 

1988 and 1989, leaving open the possibility that the creep 

and seismicity on the Hayward fault may have responded 

to a fault-relaxing event a year or two before the Loma 

Prieta earthquake, in addition to the coseismic elastic relaxation associated with the Loma Prieta earthquake itself. Because these possible relaxation events might be 

relatively small compared to stress changes produced by 

moderate-size earthquakes close to the fault and may already have been eroded by tectonic loading, they are not 

considered to have significantly affected the long-term 

estimates for the probability of a large earthquake on the 

Hayward fault. 

For low assumed values of the coefficient of friction, 

our models do not offer any ready mechanism for the 

triggering of a large earthquake on the Hayward fault by 

Depth (km) Location 

6.0 

17.8 

5.9 

8.8 

11.0 

5.9 

4.1 

7.5 

3 .O 

5.0 

13.4 

4.8 

Pajaro Gap 

Lorna Prieta 

Watsonville 

Freemont Peak 

Bear Valley 

Bear Valley 

Bear Valley 

Bear Valley 

Calaveras Reservoir 

Piedmont 

Piedmont 

Piedmont 

Magnitude 

3.9 

7.1 

5.2 

4.4 

4.5 

4.7 

4.5 

4.0 

2.4 

3.3 

4.0 

3.5 

the Loma Prieta earthquake. The increase in extension 

calculated for models along the southern part of the 

Hayward fault would be expected to bring that part of the 

fault closer to failure if friction were important, and there 

is one scenario, discussed in more detail by Simpson and 

Reasenberg (1994), that might be consistent with the inferred low apparent coefficients of friction and yet could 

offer a mechanism for triggering. This scenario would 

appeal to the pore-fluid pressure changes induced by the 

Loma Prieta earthquake as the cause for the low apparent 

coefficient of friction observed immediately after the earthquake. If these pore fluids are able to re-equilibrate with 

time (and are not trapped by impermeable barriers in the 

fault zone as suggested by Byerlee, 1992, Blanpied and 

others, 1992), then the value for the apparent of coefficient of friction would change with time, eventually reaching high laboratory values when the pore fluids had 

completely re-equilibrated. At that time the full effect of 

the unclamping of the southern end of the Hayward fault 

might be felt, which might advance the time of the next 

large earthquake on that fault by 5 to 10 years. 

We have not considered how increased stresses might 

produce higher seismicity on a fault, nor have we attempted 

to quantify the relationship between stress changes and 

seismicity changes. We used only the fact that the Coulomb stress and seismicity apparently increased and decreased in unison, with their agreement most prevalent 

for low-apparent-friction faults within about 100 km of 

the Loma Prieta epicenter. The fact that the seismicity 

changes correlate with the stresses estimated from homogeneous elastic stress models suggests that the stresses 

predicted by these models are probably correct in their 

broad pattern, though not necessarily in detail. The actual 

stress changes could differ greatly in magnitude (and could 
RESPONSE OF REGIONAL SEISMICITY TO THE STATIC STRESS CHANGE D69 

possibly differ in sign in some places) if more complex 

and geologically realistic mechanisms of mechanical coupling, possibly involving weak fault zones, rigid blocks, 

fluids, and time-dependent effects, are at work. 

Our work raises new questions and suggests some directions for future work. Concerning the sensitivity of 

San Francisco Bay Area seismicity to small stress changes: 

Have other large or moderate size earthquakes in the Bay 

Area and California caused seismicity rate changes on 

A Aftershock Zone 

- 

> 

a 

a 

0 9 Days After Mainshock 1 c k 

c San Francisco Peninsula A 

0.01 0.10 1.00 10.00 100.00 1000.00 

Days After Mainshock 

Figure 13.-Rate of m1.5 earthquakes after the Loma Prieta earthquake in zones 1,2N, 3, and 4 (see table 1) is indicated by triangles. The 

postseismic decay in seismicity rate in each area is described by Omori's 

law (curved lines). Horizontal lines represent average seismicity rates in 

the 10 years preceding the Loma Prieta earthquake. Vertical lines mark 

time of the 18 April 1990 (Mz5.6) aftershock near Chittenden. (A) Aftershock zone (Zone 1). Model fit to data between 18 October 1989 and 

17 April 1990 is shown by solid curve; model fit to all of the data is 

surrounding faults that can be similarly modeled? Using 

methods similar to the one reported here, one can calculate the static stress changes produced on Bay Area faults 

by all regional earthquakes for which reliable rupture models can be inferred. In principle, such static models can be 

combined into a single, time-dependent model that may 

be capable of explaining a significant portion of the regional background seismicity. After the stress-model-induced seismicity is accounted for, patterns and changes in 

1 6 Creeping San Andreas fault 

Days After Mainshock 

D San Gregorio fault 

Days After Mainshock 

0.01 0.10 I .oo 10.00 100-.00 1000.00 

shown by dashed curve. (B) Northern creeping segments of the San 

Andreas fault immediately southeast of the Loma Prieta rupture (Zone 

2N). Model fit to data between 18 October 1989 and 17 April 1990 is 

shown by solid curve; model fit to all of the data is shown by dashed 

curve. (C) Southern San Francisco Peninsula (Zone 3). Model reflects 

data between 18 October 1989 and 31 March 1992. (D) San Gregorio 

fault (Zone 4). Model reflects data between 18 October 1989 and 31 

March 1992. 
D70 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

the residua1 seismicity may provide information on other 

possible sources of stress, including aseismic slip on or 

below Bay Area faults. 

Concerning the apparent low friction of the faults: Do 

all faults in the Bay Area have a low effective coefficient 

of friction, or is this true of only some of the faults? Our 

result, which relies on earthquakes along all Bay Area 

faults, does not distinguish among individual faults. Is the 

low apparent friction on Bay Area faults suggested by our 

study controlled by the presence of pore fluids in the 

seismogenic crust? If so, on which faults, and over what 

characteristic time, will the effective friction change as 

fluids re-equilibrate in response to the postseismic stresses? 

Continued close monitoring of the southern Hayward fault 

seismicity and creep rates may help answer this question. 

In particular, an increase in the microseismicity and creep 

rates on the southern Hayward fault might indicate that 

pore pressure has returned (increased) to the pre-Lorna 

Prieta level and may signal the possible occurrence of a 

delayed, triggered large earthquake on the Hayward fault. 

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THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

SPATIAL VARIATIONS IN STRESS FROM THE FIRST SIX WEEKS OF 

AFTERSHOCKS OF THE LOMA PRIETA EARTHQUAKE 

CONTENTS 

By John W. Gephart, 

Cprnell University 

Page 

D73 

7 3 

74 

75 

77 

77 

80 

8 1 

8 8 

89 

89 

ABSTRACT 

The spatial variation in the state of stress along the 

main shock rupture plane during the first six weeks after 

the Loma Prieta earthquake is evaluated by estimating the 

stress tensor from groups of aftershock focal mechanisms 

in each of six subregions. Among the best-fitting stresses, 

two systematic patterns of stress variation are observed: 

(1) northwest of the main shock hypocenter the inferred 

stresses are only moderately well constrained but everywhere indicate low shear stress on the main shock fault 

plane, with ol essentially perpendicular to the plane in 

most places. The directions of o2 and o3 appear to vary 

systematically with position on the fault plane. The mechanisms are generally not consistent with north-south compression and east-west tension, as might be predicted from 

strike-slip motion on the San Andreas fault if the fault is 

assumed to be a locus of high shear stress; (2) southeast 

of the hypocenter the inferred stresses are well constrained 

and suggest that significant shear stress may have remained 

on the main shock fault plane after it failed, although the 

absolute magnitude of the shear traction cannot be determined (and may be very low if the fault is weak, as commonly assumed of the adjacent creeping section of the 

San Andreas fault). The best-fitting principal stress directions are coincident with the main shock kinematic axes 

in the region close to the hypocenter but deviate significantly from these with distance to the southeast. The direction of shear stress on the main shock fault plane varies from: (1) alignment with the overall main shock slip 

vector near the hypocenter to (2) shallow northwest-plunging at the southeast end of the aftershock zone; this variation in shear stress direction (counter-clockwise to the 

southeast when viewed in the direction of the downward 

fault pole) is of the same sense as the variation in main 

shock slip distribution determined by others, but of much 

smaller magnitude. The mechanisms southeast of the main 

shock hypocenter generally are not consistent with faultnormal compression, as proposed by some other workers 

and as might occur if the main shock slip relieved most of 

the ambient stress in the region. The pattern of stresses 

observed in this study apparently reflects the distribution 

of load on the main fault plane shortly after failure. It is 

possible that the two regions (northwest and southeast of 

the hypocenter) sample different spatial domains around 

the main shock fault plane: the diverse events to the northwest reflect the crust adjacent the fault zone, while the 

well-aligned events to the southeast reflect the fault zone 

itself. If this is so, then the observations above may support the model of Rice (1992), which predicts spatial heterogeneity in stress around a well-developed fault zone 

owing to differences in pore pressure and material properties between the fault zone and the adjoining crust. 

ANTRODUCTION 

Aftershocks occurring in the first six weeks following 

the Loma Prieta earthquake exhibit diverse mechanisms 

which vary widely with location around the main shock 

fault plane (Oppenheimer, 1990; Dietz and Ellsworth, 

1990). Whereas each event indicates a partial measure of 

the state of stress at the time and place of its occurrence, 

the well-constrained focal mechanisms of these data can 

be used to map out variations in stress in the period immediately following the Mz7.1 earthquake. Preliminary 

investigations of this kind, based on qualitative or approximate quantitative means, infer strongly heterogeneous 
D74 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

postseismic stresses, issuing from the observation of 

extreme local diversity of mechanisms, which appear to 

be mutually inconsistent (Oppenheimer, 1990; Michael and 

others, 1990). Moreover, these studies suggest that the 

aftershocks may reflect a condition in which the ambient 

tectonic stress was relieved almost entirely during main 

shock slip (Michael and others, 1990; Zoback and Beroza, 

1993; and Beroza and Zoback, 1993). This study searches 

for possible spatial heterogeneity of the stress field indicated by 164 well-constrained focal mechanisms of early 

aftershocks (Mz0.9-4.0) from around the rupture zone (fig. 

I), applying a technique that may resolve subtle stress 

variations. It appears that while a portion of the main 

shock fault plane is largely free of shear stress at this 

time, another part of the fault plane remains loaded, although the absolute magnitude of the deviatoric stress cannot be determined. This work may improve our 

understanding of the way in which stress is redistributed 

during large earthquakes. 

METHOD 

In this study the stress field is inferred from the directions of slip on fault planes, as indicated by the focal 

mechanisms of earthquakes, following the procedure of 

Gephart and Forsyth (1984) and Gephart (1990a, 1990b) 

and based largely on the earlier principles of Wallace 

(1951), Bott (1959), and McKenzie (1969). The information embodied by focal mechanisms-the orientations of 

fault planes and associated slip directions-constrains four 

of the six numbers of the stress tensor. As discussed by 

Gephart (1990a), these four dimensionless quantities can 

be expressed as three Euler angles-fixing the principal 

stress orientations, and one stress magnitude term, Rdescribing the magnitude of o2 relative to ol and 03; 

R=(02-ol)/(03-ol), such that 05 R51. The two numbers 

of the stress tensor that cannot be determined by this approach reflect the magnitudes of a regional normal and 

shear stress, om=(ol ++)I2 and ~~=(o~-o~)/2; as these 

numbers have the dimensions of stress, they cannot be 

estimated from (dimensionless) orientation data alone. By 

combining the information expressed in multiple focal 

mechanism observations, which are assumed to reflect a 

common stress field, best-fitting values of the four resolvable stress parameters can be determined. In these experiments all the data are grouped in spatially distinct 

subregions, and independent inversions are performed on 

each set. The variations in the best-fitting stress fields 

determined for all the subregions should approximate the 

continuous variation in stress across the region. 

The constraint on the four resolvable stress parameters 

is derived from the requirements that (1) there be no component of shear stress on any fault plane perpendicular to 

the direction of slip and (2) the component of shear stress 

parallel to the direction of slip have the same polarity as 

the slip. Thus, if s is the (unit) slip vector, b is the (unit) 

B-axis vector (b=nxs, where n is the unit fault normal), 
SPATIAL VARIATIONS IN STRESS FROM THE FIRST SIX WEEKS OF AFTERSHOCKS D75 

and T is the shear stress vector in the fault plane, then it is 

required that ~*b=0 and T*s>O. The inverse procedure effects a transformation of the stress tensor from these fault 

coordinates (unique to each datum) to regional coordinates (common to all data). Following Gephart (1990a), 

this is performed based on the conventional stress transformation equation, oij'=Pikhlokl (applying the Einstein 

summation convention of summing over repeated subscripts), where oij and oij are the stress tensor in the two 

different reference frames, and Pij is the transformation 

matrix of angle cosines between the two coordinates sets. 

It is important to note the extreme non-linearity of tensor 

mechanics that is bound up in the stress transformation 

equation; this implies that the shear stress direction on a 

fault plane may vary widely with small variations in the 

plane or stress tensor, and that there are always many 

stress tensors that are consistent with any given fault datum based on the two conditions above. As a practical 

matter, this precludes the possibility of substituting averaged kinematic (P, B, and T) axes for principal stress 

directions. That is, the maximum and minimum stress (oi 

and o3) directions may not be at large oblique angles 

(-45') to the fault plane, and the intermediate stress (G~) 

direction may not be at a small angle to the plane, as is 

conventionally assumed for simplicity. Thus, as only the 

direction of shear stress on the fault plane is indicated by 

the focal mechanisms, and no measure of the magnitude 

of shear stress, the approach implicitly allows that any of 

the fault planes may be either strong or weak. 

Regarding the extreme non-linearity of the problem, a 

simple and reliable approach for inverting the fault data is 

by a grid search over the four resolvable stress parameters 

(Gephart and Forsyth, 1984). A range of combinations of 

four stress parameters are tested one-by-one against all of 

the data. For each set of four parameters (a stress model), 

the misfit of each datum is determined. Here the misfit is 

defined as the smallest difference in orientation between 

the observed mechanism and any one that fits the stress 

model-in which the slip direction matches the predicted 

shear stress direction on at least one of the two nodal 

planes. The difference in orientation between any such 

observed and predicted mechanisms is equivalent to a rotation in space about some arbitrary axis. The fitness of a 

particular model relative to all the data is given by the 

average misfit rotation for the whole data set. The optimum stress model is the one which has the smallest average misfit, and confidence limits can be defined based on 

the values of the mean misfits using statistics for onenorm procedures (Gephart and Forsyth, 1984). In the 

course of the inversion, it is necessary to select the fault 

plane from among the two nodal planes of each focal 

mechanism. In the absence of an objective means for identifying the true fault plane based on observation, this is 

done by testing each nodal plane independently against 

each stress model and selecting the one with the smaller 

rotation misfit as the preferred fault plane relative to that 

stress model (the result for the alternate plane is neglected). 

To regard the physical properties of particular fault 

planes (for example, "strong" versus "weak"), it is useful 

to consider the stress vector that acts on the fault plane. 

The relation between a stress tensor and a fault geometry 

can be succinctly illustrated using the Mohr Sphere diagram presented by Gephart (1990a) (fig. 2). This is a 

generalization of the conventional Mohr Circle diagram 

for stress (for example, Jaeger and Cook, 1979), which is 

a plot of normal-versus-shear stress. The Mohr Sphere 

diagram differs from the Mohr Circle diagram in that it 

regards two orthogonal components of shear stress, parallel and perpendicular to the observed slip direction, rather 

than a single component. In both the Mohr Circle and 

Sphere figures, admissible fault data are represented as 

points which indicate the magnitudes of normal and shear 

stresses acting on a fault plane. Whereas the locus of all 

possible fault orientations in the Mohr Circle diagram is 

the area between the one large and two small circles which 

intersect at three crossing points on the normal stress axis 

(marking the principal stress magnitudes), the locus of all 

possible fault geometries in the Mohr Sphere diagram is 

the volume between the corresponding one large and two 

small spheres. (That no admissible fault geometries are 

represented by points in the regions within the two smaller 

circles/spheres or outside the largest one is an intrinsic 

property of the stress tensor.) Projections of the Mohr 

Sphere onto planes parallel to the normal stress axis resemble the Mohr Circle diagram. Two orthogonal projections of the Mohr Sphere, each plotting normal stress 

versus one component of shear stress, are sufficient to 

illustrate the distribution of fault geometries in this threedimensional figure. Each fault geometry appears as a single 

point in each of the two projections. For a particular fault 

geometry to be consistent with the stress tensor in question, based on matching shear stress and slip directions, 

the corresponding point must plot above the abscissa in 

the upper figure (T*s>~) and on the abscissa in the lower 

one (wb = O), indicating that the shear stress and slip 

vectors are aligned and of the same polarity. 

TEST 

To test for spatial variations in stress, the focal mechanisms of 164 well-constrained Loma Prieta aftershocks 

were divided among several subsets based on location 

(fig. 1). To the extent possible, divisions between the subregions were chosen to correspond with distinct spatial 

gaps among the aftershocks. Some experimentation was 

applied to establish appropriate domain sizes and boundaries, such that meaningful stress variations could be 
D76 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

determined; however, the separation was defined solely 

by event location-no consideration was given to the orientations of the mechanisms (and none were rejected as 

inconsistent, in spite of remarkable local heterogeneity). 

Three subregions-A, B, and C-were defined along the 

strike of the main shock rupture zone in the region below 

and to the southeast of the main shock auxiliary plane 

(see fig. 1). Subregion A essentially coincides with the 

northwestern end the central creeping section of the San 

Andreas fault (U. S. Geological Survey Staff, 1990). The 

aftershock zone appears to steepen, and its maximum depth 

shallows, with increased distance to the southeast from a 

point near the main shock hypocenter; thus the main shock 

fault plane may vary in orientation and vertical dimension 

from subregions C to A. Three subregions-I, J, and Kwere defined above and to the northwest of the main shock 

auxiliary plane. Subregion I comprises a group of events 

immediately northwest of the main shock hypocenter at 

Mohr Sphere 

Two orthogonal projections 

admissi bleorientations 

Figure 2.-Mohr Sphere construction: a plot of normal versus two components of shear stress (here parallel to the slip vector, s, 

and B axis, b, on any fault plane). Fault geometries falling on the [T*s>~, ~@b=0] half-plane relative to any given stress model are 

admissible, in that shear stress and slip vectors agree in both direction and polarity. 
SPATIAL VARIATIONS IN STRESS FROM THE FIRST SIX WEEKS OF AFTERSHOCKS D77 

depths >10 km, subregion J comprises a cluster of relatively deep events farther to the northwest, and subregion 

K comprises all shallow events. 

The orientations of aftershock focal mechanisms are 

illustrated by the P and T axes at the left side of figure 3. 

The P and T axes of the main shock mechanism are shown 

for reference on each plot; however, the main shock datum was not included in the inversions of the aftershock 

data sets. These figures indicate generally modest variation in the character of mechanisms in adjacent subareas, 

except across the main shock hypocenter where the variation is marked. As noted by Oppenheimer (1990), many 

of the events southeast of the main shock hypocenter (subareas A, B, and C) are aligned along a steeply-dipping 

plane and have mechanisms indicating right-lateral strikeslip on a northwest-trending fault, consistent with slip on 

the San Andreas fault. In contrast, aftershocks to the northwest of the main shock hypocenter (subareas I, J, and K) 

are distributed widely around the main shock fault plane 

with mechanisms that vary widely and generally do not 

resemble that of the main shock, suggesting that these 

events occurred on a variety of secondary structures 

(Oppenheimer, 1990). 

Independent tests were performed on each of these data 

subsets in order to observe variations in the stress field 

across the region. For each data set, an exhaustive search 

over the ranges of four stress parameters was made to 

find the model that is most nearly consistent with the 

data, based on matching observed and predicted slip directions on one of the two nodal planes of each focal 

mechanism. 

RESULTS 

The results of the experiments for each of the six subregions are illustrated in the right side of figure 3, which 

shows the best-fitting stress models, as indicated by the 

orientations of ol and o3 axes on a stereonet, a histogram 

of R values, and the associated uncertainties of each. The 

main shock focal mechanism is superimposed on each 

stereonet for reference. The principal stress directions and 

value of R of the best-fitting models of each subregion are 

listed in table 1. Figures 4 and 5 illustrate the variations 

of the best-fitting stresses among the subregions and emphasize the relation between the inferred stresses and the 

main shock mechanism. The misfit rotations and the observed and predicted (nearest perfectly fitting) focal mechanisms of each event relative to the best-fitting stress model 

in each region are given in table 2; this listing indicates 

the chosen (preferred) fault plane from among the two 

nodal planes of each mechanism. Based on these results, 

there appears to be a marked difference in the nature of 

the stresses on either side of the main shock hypocenter. 

Thus, in the following discussion the results of the three 

southeastern and three northwestern subsets are presented 

separately. 

SOUTHEAST OF THE MAIN SHOCK 

HYPOCENTER 

To the southeast of the main shock, the best-fitting 

stresses fit the data very well, with average rotation misfits of 4'; this suggests that the stresses may be reasonably uniform within each group, as assumed in performing 

the analysis. Although the aftershock mechanisms of each 

of the three subsets appear to be qualitatively similar to 

one another, based on P- and T-axis orientations, the ranges 

of acceptable stresses vary systematically from one subregion to the next with distance along the main shock fault 

trace (fig. 3). In subarea C, near the main shock hypocenter, ol is nearly horizontal north-south and o3 plunges 

moderately to the east. In subarea A, at some distance 

from the main shock hypocenter, ol plunges moderately 

to the north and o3 is nearly horizontal east-west. In subarea B, located between subareas C and A, the principal 

stresses are intermediate in orientation to those on either 

side. Although the formal uncertainties in stress orientations are relatively small, the results do not absolutely 

require that the stresses are smoothly varying from subareas C to A, as there is some overlap at the 95-percent 

confidence level; however, as the variation in the bestfitting stresses are systematic, the following discussion 

assumes that this observed trend is meaningful. In all three 

cases, the acceptable o1 directions (within the 95-percent 

confidence limit) are nearly confined to the main shock 

dilatational quadrants, and the 03 directions to the compressional quadrants. The preferred value of R varies 

among the subregions from 0.3 to 0.7, but shows no clear 

systematic spatial trend, as with the stress directions. 

The orientations of the best-fitting stresses across the 

aftershock region appear to be systematically related to 

the focal mechanism of the main shock, as illustrated in 

figure 4. In subarea C (adjacent to the main shock hypocenter) the inferred ol, 02, and <y3 directions are very 

nearly aligned with the P, B, and T axes, respectively, of 

the main shock focal mechanism. Progressively to the 

southeast of the main shock hypocenter, in subareas B 

and A, the axis remains parallel to the main shock 

fault plane, but of increasing rake from a northwest strike 

line. Concomitant variations in the the other two principal 

stresses are also systematic in space: o1 steepens progressively northward, and o3 shallows progressively eastward, 

from subareas C to B to A. Based on these inferred stress 

models, the predicted slip (shear stress) direction on the 

main shock fault plane varies from alignment with the 

observed main shock slip direction in subarea C, to shallow southeast-plunging (with a small reverse dip-slip 

SPATIAL VARIATIONS IN STRESS FROM THE FIRST SIX WEEKS OF AFTERSHOCKS D79 

Table 1 .-Best-fitting stress models of each subarea 

[Principal stress directions indicated by plunge and 

azimuth] 

Subarea Principal Stress Directions - R 

component) in subarea B, to shallow northwest-plunging 

(with a small normal dip-slip component) in subarea A. It 

can be shown that if the main shock fault plane in subareas B and A is steeper than shown in figure 4 (but with a 

slip vector of the same rake), as suggested by the alignment of aftershocks in that region, then the mismatch between the observed and predicted main shock slip is 

increased in that region; thus we cannot easily associate 

the along-strike differences among the inferred stresses 

with variations in main shock fault orientation. 

Relative to the best-fitting stress tensor in each subregion, the traction vector on each inferred fault plane is 

illustrated in the Mohr Sphere diagrams of figure 5 (again, 

aftershock fault planes were selected from among the two 

nodal planes of each event in the course of the inversion, 

so as to minimize the difference between observation and 

model). That the observed and predicted slip directions 

are nearly in agreement on most of the aftershock faults is 

indicated by a clustering of poles on the [T~s>O, ~eb=O] 

half-plane. For most of the aftershocks in all three subregions, the magnitude of shear stress approaches the maximum for the region (near the top of the largest circle in 

the plot of T*S versus 0); this suggests that the aftershock 

fault planes may be relatively strong, as they apparently 

fail only under relatively high shear stress. Of course as 

described above, the absolute shear stress magnitude is 

unknown, as the magnitude of T (the radius of the Mohr 

Sphere) is unconstrained in the present analysis. 

Also shown in figure 5 is the traction on the main shock 

fault plane for the best-fitting stress models of each of the 

subregions. In all three southeastern subregions the total 

relative magnitude of shear stress on the main shock is 

high (nearly the regional maximum), but the magnitude of 

the components parallel and perpendicular to the observed 

slip direction vary considerably with location. The main 

shock fault geometry is consistent with the inferred stresses 

in subregion C, immediately to the southeast of the main 

shock hypocenter, but it is progressively less consistent 

with distance to the southeast in regions B and A, as 

indicated by the increasing divergence of the main shock 

datum from the abscissa, [~eb=O], in the lower figures. 

This is consistent with a spatially varying inferred shear 

stress direction on the main shock fault plane, as shown 

in figure 4. It is interesting that, while in all three southeastern subregions the main shock P and T axes fall within 

the scatter of those of the aftershocks (fig. 3), apparently 

the main shock and aftershock stress tensors are consistent only in subregion C; that is, in terms of the inferred 

stress tensors, the main shock mechanism is clearly distinct from the aftershocks in subregions B and (especially) 

A, as indicated in the Mohr Sphere diagrams. 

NW of mainshock 

Subareas J K I 

% J (R = 0.4) 

K (R = 0.4) 

QI I (R = 0.8) 

SE of mainshock 

Subareas C B A 

C (R = 0.6) 

B (R = 0.3) 

~v, A (R = 0.7) 

Figure 4.-Summary of best-fitting stress models for each subregion 

and comparison with main shock focal mechanism. Top, Three northwestern subregions: principal stress axes (oi, 02, 03) tend to align parallel and perpendicular with the main shock fault plane. In all cases the 

resolved shear stress on the main shock fault plane is very small relative 

to the regional maximum. Bottom, Three southeastern subregions: principal stress directions in subregion C are aligned with main shock kinematic axes, but drift away in subregions B and A; in all cases the % axis 

is nearly parallel to the main shock fault plane. The predicted shear 

stress direction on the main shock fault plane varies smoothly between 

subareas C, B, and A, as shown. 
D80 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

NORTHWEST OF THE MAIN SHOCK 

HYPOCENTER 

To the northwest of the main shock hypocenter the average rotation misfits among the best-fitting stress models 

in each subarea are on the order of 6-9' (fig. 3)-greater 

than to the southeast but less than the uncertainty in the 

focal mechanism determinations; thus, the analysis does 

not preclude the possibility that stresses are essentially 

uniform within each subregion. Moreover, the stresses 

appear to be less uniform between subareas I, J, and K 

than among the subareas to the southeast. 

As shown in figure 4, the stresses northwest of the 

main shock hypocenter appear to bear a very different 

relationship to the main shock fault plane than to the southeast; here the principal stress directions are nearly parallel 

and perpendicular to the fault. In subarea I, near the main 

shock hypocenter, the stresses are especially poorly resolved, although the best-fitting oi and o3 directions are 

nearly aligned with the B and T axes of the main shock, 

respectively. In an experiment similar to the present one 

(using a different algorithm to invert for stress), Michael 

and others (1990) were unable to identify acceptable 

homogeneous stresses in the region of subarea I, and cone main shock aftershocks 

Figure 5.-Two orthogonal projections of the Mohr Sphere construction for the six subregions, corresponding to the best-fitting 

stresses in each case. Fault geometries that are consistent with the stress model fall on the abscissa in the lower projection and 

above the abscissa in the top one. Both aftershock (small open circles) and main shock (large filled circles) mechanisms are shown. 
SPATIAL VARIATIONS IN STRESS FROM THE FIRST SIX WEEKS OF AFTERSHOCKS D8 1 

eluded that the stresses there were extremely heterogeneous. 

In subareas J and K the inferred stress orientations appear to be similar at high confidence levels (for example, 

95-percent), with acceptable ol axes clustering about the 

pole of the fault plane and o3 axes distributed in orientations parallel to the fault plane; among the best-fitting 

stresses, the ol directions of the two subareas are nearly 

identical but the o3 directions are nearly orthogonal to 

one another. This is illustrated more fully in figure 6, 

which shows the results of tests based on only a very 

limited set of stress models-those with q perpendicular 

to the main shock fault plane. Among these models, figure 6A shows the range of acceptable 03 directions parallel to the main shock fault plane, in the same format as in 

figure 3; the best o3 directions are oriented differently 

along the fault plane for subareas J versus K. This description involves some inherent ambiguity, as each 03 

direction is associated with several models, spanning a 

range of values of R (and each value of R is associated 

with a range of o3 directions). Figure 6B is an unambiguous illustration of the full range of models with oi perpendicular to the main shock fault plane, with each point 

indicating a single model (a unique combination of stress 

directions and R). At the 95-percent confidence limit, there 

is nearly a complementary distribution of acceptable models for the two subregions-generally, models acceptable 

for J are not acceptable for K, and vice versa. Thus, while 

certain elements of the ranges of acceptable stresses are 

common to the two regions (q direction), other elements 

appear to be mutually exclusive (the combination of 

direction and R). 

In the Mohr Sphere plots in figure 5, there is considerable scatter of the aftershock data about the [T*S>O, T*~=O] 

half-plane among subareas I, J, and K, indicating that the 

fit to the optimum stress models are only moderately good. 

In the [T*S versus o] projections, most aftershocks plot at 

relatively low shear stress magnitudes (at considerable 

distances from the top of the largest circle), suggesting 

that the aftershock fault planes may be weak. In all three 

of these subregions, the main shock fault geometry plots 

very near to a point of intersection of two of the spheres, 

indicating that the fault plane is nearly a principal planeand thus experiences a very low relative shear stress magnitude at the time of the aftershocks. In such cases, the 

direction of shear stress on the fault plane is of little significance (and so is not shown at the top of figure 4). 

DISCUSSION 

The above results indicate a systematic spatial variation 

in stress along the main shock rupture zone, with markedly different expressions on the northwest and southeast 

sides of the main shock hypocenter. To the northwest of 

the hypocenter, principal stress axes exhibit large differences in orientations between subregions (an exchange of 

axes among the principal stresses) but in a manner that 

everywhere yields relatively low shear stress on the main 

shock fault plane. To the southeast of the hypocenter, 

principal stress axes exhibit only minor differences between subregions (small deflections of axes) but in all 

areas so as to yield relatively high shear stress on the 

main shock fault plane, and with a shear stress vector of 

smoothly varying rake with position along the fault. This 

distinct spatial and geometric association of the main shock 

and stress inferred from aftershocks implies a physical 

connection between the main shock failure and the 

postseismic state of stress. 

Presumably, during the main shock the ambient 

(preseismic) stress was reduced in magnitude and redistributed, although perhaps not completely or evenly; the 

stress determined from the early aftershocks must reflect 

the stress state that exists following the main shock stress 

drop. The marked difference in stress observed on either 

side of the hypocenter reflects the way that stress was 

reorganized on the rupture plane. To the northwest, the 

main shock fault plane appears to be mostly relieved of 

stress, suggesting that the stress drop was nearly complete 

(as suggested by Zoback and Beroza, 1993, and Beroza 

and Zoback, 1993), while to the southeast the fault plane 

appears to remain loaded, possibly owing to incomplete 

stress drop on that portion of the fault (in apparent conflict with the interpretation of these workers). 

It is interesting that the pattern of the shear stress direction on the main shock fault plane inferred between subareas C and A (fig. 4) bears some resemblance to the 

interpreted variations in the main shock slip distribution 

from geodetic-leveling and strong-motion data. The latter, 

with minor differences among various workers (for example, Beroza, 199 1 ; Hartzell and others, 199 1 ; Marshall 

and others, 1991; Steidl and others, 1991), indicate predominant thrust slip to the northwest of the main shock 

hypocenter, and predominant right-lateral strike slip to 

the southeast (although Wald and others, 1991, find a 

more uniform slip rake). The inferred postseismic stresses 

from this study indicate a similar counter-clockwise rotation of predicted slip (shear stress) on the main shock 

fault plane with distance to the southeast, though of much 

smaller magnitude than observed in coseismic slip, and 

limited to a region only southeast of the hypocenter. Assuming that the slip distribution reflects shear stress directions (as implicit in the stress inversion procedure), it 

is possible that a large spatial variation in preseismic stress, 

which is reflected in the two disparate domains of main 

shock slip, evolved into a more subdued and localized 

variation in postseismic stress observed among the aftershocks. 
Table 2.-Data and results for best-fitting stress models of each subarea 

[All angles in degrees] 

Misfit Rotation Vector Observed Focal Mechanism Rotated Focal Mechanism 

Event Misfit Rot. Rot. Obs. Obs. Obs. Obs. Rot. Rot. Rot. Rot. 

# ~otations Axis Axis Fault Fault Aux. Aux. slip4 Fault Fault Aux. Aux. slip4 

Pref. (Alt.) ~zim.~ plunge2 Strike Dip3 Strike ~i~~ Strike Dip3 Strike Dip3 

Subarea A: 1 

2 

3 

4 

5 

6 

7 

8 

9 

10 

11 

12 

13 

14 

15 

16 

17 

18 

19 

Subarea B: 1 

2 

3 

4 

5 

6 

7 

8 

9 

10 

11 

12 

13 

14 

15 

16 
1 Table 2.Ã‘Continue 

Subarea C: 1 0.4 

2 9.3 

3 1.3 

4 1.6 

5 3.2 

6 1.5 

7 0.8 

8 1.2 

9 5.1 

10 0.5 

11 2.1 

12 6.8 

13 7.3 

14 10.9 

15 1.5 

16 2.3 

17 9.6 

18 5.6 

19 4.1 

20 17.6 

2 1 1.9 

22 9.0 

23 7.3 

24 16.8 

25 1 .O 

26 0.4 

27 1.8 

28 0.3 

Subarea I: 1 2.0 
Table 2.Ã‘Continue 

Subarea J: 1 

2 

3 

4 

5 

6 

7 

8 

9 

10 

11 

12 

13 

14 

15 

16 

17 

18 

19 

20 

2 1 

22 

Subarea K: 1 

2 

3 

4 

5 

6 

7 
Table 2.Ã‘Continue 

l~inimum rotation between the observed mechanism and any one consistent with the model for: preferred (alternate) nodal plane 

rotation axis azimuth and plunge-applies to preferred nodal plane (the one with smaller misfit) 

^ip angle; dip direction is 90Â clockwise from strike direction 

4Predominant slip component of preferred nodal plane: N = normal, T = thrust, R = right-lateral, L = left-lateral 
D86 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Thus, the final stress state may reflect an overall reduced and redistributed pattern of loading on the main 

rupture plane; the fact that the principal stress axes are 

not parallel and perpendicular to the plane in subareas AC suggests that the shear stress is not completely relieved 

in that region. This differs from the general conclusions 

of Zoback and Beroza (1993) and Beroza and Zoback 

(1993), although they included in their analysis all aftershocks from a 21-month period following the main shock, 

and they did not explore for spatial variations by subdividing the area, as done here. That the southeast end of 

the aftershock zone approaches the creeping section of 

the San Andreas might suggest that shear stress magnitudes are low in that region-that is, where the fault is 

creeping we might surmise that it is weak, as it does not 

support sufficient ambient stress to allow it to fail in large 

events. Thus the aftershocks in this region might be induced by even a small static stress change caused by the 

main shock (for example, as inferred to have occurred on 

the Calaveras fault in response to the Morgan Hill earthquake, by Oppenheimer and others, 1988); however Beroza 

and Zoback (1993) found that the Loma Prieta aftershock 

geometries are inconsistent with this hypothesis. 

In subareas J and K, northwest of the main shock hypocenter, the inferred stresses have a common oi direction, 

essentially normal to the main shock fault plane, indicating that virtually no shear stress remains on that plane in 

either area (consistent with Zoback and Beroza, 1993, and 

Beroza and Zoback, 1993). However, the other two principal stress directions, which are parallel to the main shock 

Figure 6.-Details of results for subareas J and K. Top, Results of grid and 9 directions vary by azimuth, R varies radially, such that each point 

search of stress models with a, direction normal to the main shock fault represents a unique stress model. (Stress directions are indicated by plunge 

plane. The optimum % directions in the two cases are oriented in comple- and azimuth.) There is a strong negative correlation between the ranges 

mentary directions within the fault plane. Bottom, A more detailed illus- of acceptable models for the two subregions. 

tration of all stress models with q perpendicular to the fault plane~a~ 
SPATIAL VARIATIONS IN STRESS FROM THE FIRST SIX WEEKS OF AFTERSHOCKS D87 

fault plane and therefore do not contribute to shear stress 

on it, exhibit different orientations in the two subareas 

(fig. 6). These observations suggest that at least a small 

amount of deviatoric stress remains in the region, but that 

it is of an orientation that cannot be relieved by slip on 

the main shock fault plane. This might account for the 

observation that aftershocks in the region exhibit diverse 

focal mechanisms that are generally dissimilar to that of 

the main shock (Oppenheimer, 1990). 

It is possible that the systematic and complementary 

distributions of acceptable stresses in subareas J and K 

shown in figure 6 may reflect variations owing to differences in relative locations along the fault plane, although 

the mechanical basis for this is not entirely clear. Subarea 

J is located essentially up the slip vector from the hypocenter (fig. I), in a region through which the main rupture 

propagated as a Mode I1 crack (in-plane shear; Atkinson, 

1987); in contrast, subarea K is everywhere significantly 

up-dip of the hypocenter, where the main rupture had a 

greater Mode I11 nature (anti-plane shear). It is unclear 

why a difference in failure mode should give rise to the 

observed expression among the aftershock stresses, but 

the observed symmetry and systematic pattern of inferred 

stresses in figure 6 is intriguing. In any case, the differences observed between subareas J and K apparently reflect minor heterogeneities in the postseismic stress, of a 

kind that do not significantly contribute to the shear stress 

on the main shock fault plane. 

An explanation for the difference in stress behavior observed on either side of the main shock hypocenter may 

emerge from the model of Rice (1992). According to Rice, 

maturely deformed fault zones such as the San Andreas 

may exhibit two distinct stress domains, which arise due 

to spatial variation in pore pressure or physical properties 

of the region, or both. In his model, the narrow region 

within the fault zone indicates stress consistent with laboratory-derived frictional properties of rocks (with oi at 

about 25-30' to the fault plane), while the surrounding 

region indicates near fault normal compression; this may 

resolve the long-standing apparent conflict between the 

inferred weakness of some fault zones, based on the lack 

of a significant heat flow anomaly, and experimental and 

theoretical estimates of fault strength. In the present case, 

it is possible that the heterogeneous aftershocks northwest 

of the main shock hypocenter record stress off the fault 

plane, while the more uniform aftershocks southeast of 

the hypocenter indicate stress inside a well-developed fault 

zone. If this is so, then the stresses inferred from the early 

aftershocks, while reflecting real and meaningful spatial 

variations in stress, may not fully sample all the variation 

throughout the region. 

The determinations of stress in this study may be of 

some value in assessing the seismic hazard of known faults 

in the region. If the four stress parameters of interest here 

are known in a region, together with the orientations of a 

particular planar zone of weakness, then a forward procedure can be applied to evaluate the relative potential for 

failure of the prospective faults, as shown in figure 7. As 

discussed above, Mohr Sphere diagrams can be constructed 

for each region in which the four stress parameters are 

known: three principal stress directions and the value of 

R. These diagrams illustrate the magnitudes of relative 

normal and shear stress on planes of fixed orientation, 

normalized by the unknown quantities im and on,. Figure 

7 shows the projections of the Mohr Spheres on the [~ob=O] 

plane, such that shear stress and slip vectors are aligned. 

The Mohr Spheres for five of the subareas are superimposed; for each of these, the sizes of the two inner spheres 

are different from those of the other subareas, so none are 

shown. 

On figure 7, the pole of a single plane of orientation, 

N35OW, vertical, roughly characteristic of the regional 

San Andreas fault system, is shown relative to the stresses 

in each region-the pole plots in different places-on the 

Mohr Sphere for each region because the stresses for which 

the sphere is constructed are different in each case. By 

requiring that the slip vector match the shear stress direction on the fault plane, the focal mechanisms predicted 

for each subarea can be determined, as shown. Conventional Mohr or Coulomb fault mechanics (for example, 

Jaeger and Cook, 1979) suggests that failure should occur 

on planes of relatively high shear stress and relatively low 

normal stress. Thus, neglecting the uncertainty of the absolute stress magnitudes, failure is most likely on a fault 

of this orientation in the vicinity of subarea A, for which 

the pole falls at the upper left of the Mohr Sphere, and 

less likely near subareas J and K, with poles at the middle 

right side of the figure. A number of considerations may 

limit the utility of this notion, as only local stress and 

fault geometries are considered; this discussion neglects 

other relevant physical parameters which are very poorly 

known, such as possible variation in fault properties and 

local stress heterogeneities. Perhaps most importantly, the 

values of T and (7, which cannot be estimated in the 

present analysis of fault geometries (because they are not 

dimensionless), are of particular significance for the relative magnitudes of shear and normal stress on all planes 

and thus for the likelihood of failure. 

A different experiment regarding the relation between 

stress and earthquakes, applied by Reasenberg and Simpson 

(1992), models the Loma Prieta main shock as a dislocation in an elastic half-space and evaluates seismic risk by 

predicting a spatially-varying pattern of static stress 

changes on known vertical fault planes of the San Andreas fault system in central California. This analysis, although based on simplified fault orientation data, explicitly 

considers changes in absolute normal and shear stress 

magnitudes-quantities that are certainly relevant to the 

occurrence of fault rupture. The predicted stress changes 

and consequent seismic risk depend strongly on fault trend 
D88 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

and location; these generally agree well with observed 

changes in rates of seismicity from before to after the 

main shock. 

The forward procedure employed by Reasenberg and 

Simpson (1992) and the inverse procedure of this study 

may be complementary in some respects. Not only do the 

two analyses apply information in opposite directions (forward versus inverse) and presume different associations 

between earthquakes with the stress field (active versus 

passive), but they treat substantially different aspects of 

the stress tensor. Of the two experiments, the dislocation 

modeling addresses the aspect of the stress tensor that 

affects the occurrence of failure on fault planes of known 

(or assumed) orientation and physical properties, while 

the inverse approach regards the part of the stress tensor 

that governs the geometry of fault motion. 

CONCLUSIONS 

Analysis of the focal mechanisms of groups of aftershocks from the first six weeks after the Loma Prieta 

earthquake indicates a systematic pattern of postseismic 

stresses along the main rupture plane. To a first order, the 

main shock fault plane appears to be nearly completely 

relieved of shear stress in the region northwest of the 

main shock hypocenter but may remain loaded by shear 

stress to the southeast; the direction of shear stress on the 

fault plane is consistent with the overall main shock slip 

in the region immediately to the southeast of the hypocenter but deviates from it considerably further to the 

southeast near the end of the aftershock zone. Within each 

of these two apparently independent stress domains (northwest and southeast of the main shock hypocenter), some 

variation in stress is observed at length scales of about 10 

km or less. From these observed patterns of stress variation, it appears that steep northwest- or north-northwesttrending faults in the region may be at greater risk of 

failure near the southeast end of the aftershock zone than 

to the northwest of the main shock hypocenter. It is important to note that these interpretations are based only on 

the estimates of the four dimensionless stress parameters 

derived from the present inverse procedure, including the 

three principal stress directions and just one measure of 

stress magnitude; the absolute magnitudes of both normal 

and deviatoric shear stresses are unconstrained in this 

analysis. Thus it is possible that, while in some places the 

main shock fault plane occurs in an orientation that is 

favorable for renewed slip (in other words, relatively high 

shear stress and low normal stress compared to other orientations), the absolute shear stress magnitude may be 

low and the normal stress magnitude may be high, such 

that the plane might not be likely to fail. The difference in 

stress inferred on either side of the main shock hypocenter (along strike variation) may indicate a difference 

in the nature of the two aftershock regions relative to the 

width of the main shock fault zone (across strike variation), as in the model of Rice (1992). In this case, the 

aftershocks to the northwest of the hypocenter may reflect 

the volume surrounding the fault zone, while the ones to 

the southeast may reflect the condition of the fault zone 

itself. 

Figure 7.-Projection of Mohr Sphere diagram for 

five of the subregions (as in fig. 5). The diagram 

was constructed using the best-fitting stress model 

found in each region (each with a unique value of 

R, which would prescribe the size of the two inner circles, but is not shown). Plotted on the projection is the pole of a fault of orientation N35OW, 

vertical, together with the corresponding focal 

mechanism that would yield [TW-O, ~ob=O]. A 

weak plane of this orientation may be at risk of 

failure under the stress at subarea A, as it experiences relatively high shear stress and low normal 

stress, and thus may exceed conventional failure 

envelopes. 
SPATIAL VARIATIONS IN STRESS FROM THE FIRST SIX WEEKS OF AFTERSHOCKS D89 

ACKNOWLEDGMENTS 

I am grateful to David Oppenheimer and others at the 

U. S. Geological Survey for compiling and making available the focal mechanism data used in this study and for 

helpful discussions. David, Andy Michael, and Paul 

Reasenberg provided thoughtful reviews. This work was 

supported by USGS Grant 14-08-0001-G184 1. INSTOC 

Contribution 193. 

REFERENCES CITED 

Atkinson, B.K., 1987, Introduction to fracture mechanics and its geophysical applications, in Atkinson (ed.), B.K., Fracture mechanics 

of rock: London, Academic Press, p. 1-26. 

Beroza, G.C., 1991, Near-source modeling of the Loma Prieta earthquake: Evidence for heterogeneous slip and implications for earthquake harzard: Bulletin of the Seismological Society of America, 

v. 81, no. 5, p. 1603-1621. 

Beroza, G.C., and Zoback, M.D., 1993, Mechanism diversity in the 

Loma Prieta aftershocks and the mechanics of main shock-aftershock interaction: Science, v. 259, p. 210-213. 

Bott, M.H.P., 1959, The mechanics of oblique slip faulting: Geological 

Magazine, v. 96, p. 109-117. 

Dietz, L.D., and Ellsworth, W.L., 1990, The October 17, 1989, Loma 

Prieta, California, earthquake and its aftershocks: Geometry of the 

sequence from high-resolution locations: Geophysical Research 

Letters, v. 17, no. 9, p. 1417-1420. 

Gephart, J.W., 1990a, Stress and the direction of slip on fault planes, 

Tectonics, v. 9, p. 845-858. 

1990b, FMSI: A FORTRAN program for inverting faultlslickenside and earthquake focal mechanism data to obtain the regional 

stress tensor, Computers & Geosciences, v. 16, p. 953-989. 

Gephart, J.W., and Forsyth, D.W., 1984, An improved method for determining the regional stress tensor using earthquake focal mechanism data: Application to the San Fernando earthquake sequence, 

Journal of Geophysical Research, v. 89, p. 9305-9320. 

Hartzell, S.H., Stewart, G.S., and Mendoza, C., 1991, Comparixon of 

L and L, norms in a teleseismic waveform inversion for the slip 

history of the Loma Prieta, California, earthquake: Bulletin of the 

Seismological Society of America, v. 81, no. 5, p. 1518-1539. 

Jaeger, J.C., and Cook, N.G.W., 1979, Fundamentals of rock mechanics: Chapman and Hall, New York, 593 p. 

Marshall, G.A., Stein, R.S., and Thatcher, W., 1991, Faulting geometry and slip from co-seismic elevation changes: The 18 October 

1989, Loma Prieta, California, earthquake: Bulletin of the Seismological Society of America, v. 81, no. 5, p. 1660-1693. 

McKenzie, D.P, 1969, The relationship between fault plane solutions 

for earthquakes and the directions of the principal stresses: 

Bulletin of the Seismological Society of America, v. 59, p. 591- 

601. 

Michael, A.J., Ellsworth, W.L., and Oppenheimer, D.H., 1990, 

Coseismic stress changes induced by the 1989 Loma Prieta, California earthquake: Geophysical Research Letters, v. 17, no. 9, p. 

1441-1444. 

Oppenheimer, D.H., 1990, Aftershock slip behavior or the 1989, Loma 

Prieta, California, earthquake: Geophysical Research Letters, v. 

17, no. 8, p. 1199-1202. 

Oppenheimer, D.H., Reasenberg, P.A., and Simpson, R.W., 1988, 

Fault plane solutions for the 1984 Morgan Hill, California, earthquake sequence: Evidence for the state of stress on the Calaveras 

fault: Journal of Geophysical Research, v. 93, p. 9007-9026. 

Reasenberg, P.A., and Simpson, R.W., 1992, Response of regional 

seismicity to the static stress change produced by the Loma Prieta 

earthquake: Science, v. 255, p. 1687-1690. 

Rice, J.R., 1992, Fault stress states, pore pressure distributions, and the 

weakness of the San Andreas fault, in Evans, B., and Wong, T.-F., 

eds., Fault mechanics and transport properties of rock: London, 

Academic Press, p. 475-503. 

Steidl, J.H., Archuleta, R.J., and Hartzell, S.H., 1991, Rupture history 

of the 1989 Loma Prieta, California, earthquake: Bulletin of the 

Seismological Society of America, v. 81, no. 5, p. 1573-1602. 

U.S. Geological Survey Staff, 1990, The Loma Prieta event: An anticipated event: Science, v. 247, p. 286-293. 

Wald, D.J., Helmberger, D.V., and Heaton, T.H., 199 1, Rupture model 

of the 1989 Loma Prieta earthquake from inversion of strong-motion and broadband teleseismic data: Bulletin of the Seismological 

Society of America, v. 81, no. 5, p. 1540-1572. 

Wallace, R.E., 1951, Geometry of shearing stress and relation to faulting: Journal of Geology, v. 59, p. 118-130. 

Zoback, M.D., and Beroza, G.C., 1993, Evidence for near-frictionless 

faulting in the 1989 (M 6.9) Loma Prieta, California, earthquake 

and its aftershocks: Geology, v. 21, p. 181-185. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKE OCCURRENCE 

- -- 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

LOMA PRIETA AFTERSHOCK RELOCATION WITH S-P TRAVEL 

TIMES FROM A PORTABLE ARRAY 

CONTENTS 

By Susan Y. Schwartz and Glenn D. Nelson, 

University of California, Santa Cruz 

Abstract ............................................................... 

Introduction ............................................................... Method and data ......................................................... Synthetic examples ...................................................... 

Loma Prieta aftershock locations ..................................... 

Discussion and conclusions ............................................ 

Acknowledgments ....................................................... 

References cited ......................................................... 

Page 

D9 1 

9 1 

92 

9 3 

96 

101 

102 

102 

ABSTRACT 

Aftershocks of the Loma Prieta earthquake are located 

using S-P arrival time measurements from stations of the 

PASSCAL aftershock deployment. We demonstrate the 

effectiveness of using S-P arrival time data in locating 

earthquakes recorded by a sparse, three-component network. Events are located using the program QUAKE3D 

(Nelson and Vidale, 1990) and a three-dimensional Pwave velocity model developed independently for this region. Both a constant and a variable vplvs ratio model are 

used to generate S-wave velocities. The dense coverage of 

the area around the Loma Prieta rupture zone by instruments of the California Network (CALNET) has allowed 

the U.S. Geological Survey (USGS) to find P-wave earthquake locations, which we compare with our solutions. 

We also perform synthetic calculations to estimate realistic location errors resulting from uncertainties in both the 

three-dimensional velocity structure and the timing of arrivals. These calculations provide a comparison of location accuracies obtained using S-P arrival times, S and P 

arrival times and P times alone. We estimate average absolute errors in epicentral location and in depth for the 

Loma Prieta aftershocks to be between 1 and 2 km using 

S-P phase data and the sparse PASSCAL instrument coverage. The synthetic tests show that these errors are much 

smaller than those predicted using P-wave data alone and 

are nearly the same as those predicted using S and P 

phase data separately. This suggests that future aftershock 

recording deployments with sparse networks of threecomponent data can retrieve accurate event locations even 

if absolute timing is problematic using S-P times. 

INTRODUCTION 

Seismic studies using local networks have revealed important information about earthquake dynamics, fault zone 

structure, and the effects of near-surface geology on the 

amplification of strong ground motion. Determination of 

accurate hypocentral locations is an essential step in analysis of local earthquake data; however, accurate locations 

are difficult to obtain in the common situation of sparse 

instrument coverage. Location difficulties arise from several sources: the mathematical instability of solutions due 

to the interdependence of origin time and depth, inadequacies in the velocity model used in the location process, and uncertainties in arrival times caused by both 

reading and clock errors. In this paper, we demonstrate 

the effectiveness of utilizing S-P differential arrival-time 

data in locating earthquakes recorded by a sparse threecomponent network. The use of S-P times for earthquake 

location removes the trade-off between depth and origin 

time as well as arrival-time uncertainties arising from clock 

errors. Furthermore, locations derived from S-P arrival 

times are less sensitive to errors in the velocity model, 

provided that the Poisson ratio remains relatively constant 

in the region under consideration. 

We use the finite-difference location program 

QUAKE3D (Nelson and Vidale, 1990) to locate aftershocks of the Loma Prieta earthquake using data recorded 

by stations of the PASSCAL aftershock deployment 

(Lerner-Lam and others, 1990; Schwartz and others, 1990). 

The program QUAKE3D locates earthquakes in regions 

with complex three-dimensional velocity heterogeneity. 

The P-wave velocity structure in the source region of the 

Loma Prieta earthquake has been well studied, and both 

two-dimensional (Dietz and Ellsworth, 1990) and threedimensional (Eberhart-Phillips and others, 1990a; Lees, 

1990) velocity models have been computed. The Loma 

Prieta earthquake occurred in a region of dense CALNET 

D9 1 
D92 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

instrument coverage, providing accurate locations for most 

of the aftershocks that we compare with our solutions. 

Therefore, the aftershock sequence of the Loma Prieta 

earthquake provides an excellent opportunity to evaluate 

the effectiveness of using S-P arrival times to locate earthquakes in a realistic three-dimensional velocity structure. 

In this study, we are constrained to use differential arrival 

times due to uncertainties in the absolute timing of the 

seismic phases recorded by the PASSCAL instruments. 

Because such timing uncertainties are often encountered 

in seismic field deployments, earthquake location techniques that require only the relative timing of phases are 

useful. The application of an S-P location scheme can 

also substantially reduce the time required to locate earthquakes for studies involving large volumes of data, since 

time-consuming clock calibrations and corrections are not 

necessary. 

METHOD AND DATA 

The location algorithm QUAKE3D has been described 

elsewhere (Nelson and Vidale, 1990) and will only be 

summarized briefly here. QUAKE3D is a two-step location procedure where the first step is the discretization of 

the source volume and the computation of travel times 

from all recording stations to all points in the grid using 

the finite-difference method of Vidale (1990). S-wave 

travel times are produced from P-wave travel times by 

assuming a constant Poisson ratio throughout the source 

volume or by using a three-dimensional model of vplvs 

ratios determined for this region by Thurber and Atre 

(1993). Next, the earthquake hypocenters are determined 

by finding the position within the source volume that yields 

the smallest travel time residuals for a set of arrival times 

for each earthquake. Since QUAKE3D is a grid searching 

algorithm rather than a least-squares iterative inversion, 

as are most earthquake location programs, the Ll or L2 

norms can be employed with equal ease to determine earthquake locations. Our solutions are computed using the Ll 

norm, since it is less sensitive to data outliers. 

In this study, we use S-P times to locate aftershocks of 

the 1989 Loma Prieta earthquake rather than only P times, 

as in Nelson and Vidale (1990). The S-P arrival times are 

picked by hand from recordings at stations of the Loma 

Prieta IRIS-PASSCAL aftershock deployment. The recording instruments consisted of Reftek 16-bit digitizers with 

L-22 (2 Hz) three-component geophones. They were deployed in the northern section of the Loma Prieta aftershock zone from October 20 through November 21, 1989. 

We locate 50 aftershocks with magnitudes between 1.5 

and 3.0 that occurred in the region of dense instrument 

coverage between October 22 and November 4. In this 

period, between 10 and 18 widely distributed stations were 

in operation. Between 4 and 10 stations yielded highquality P- and s-wave arrivals for each event. This number is significantly lower than the total number of stations 

deployed (21) for several reasons: (1) only 10 stations 

were established during the first week, with instrument 

coverage expanding southward to a total of 20 stations by 

the end of the second week; (2) not all stations recorded 

all events; (3) some stations had instrument problems; 

and (4) some arrivals were not impulsive enough to accurately read P- and S-wave times. Figure 1 shows a map of 

the PASSCAL station locations and the USGS locations 

(Dietz and Ellsworth, 1990) of the events we analyze. 

We locate these events using a three-dimensional velocity model derived through inversion of CALNET Pwave arrival-time data from Loma Prieta aftershocks 

(Eberhart-Phillips and others, 1990b). The three-dimensional velocity model was interpolated from the published 

model to produce a uniform Cartesian grid (horizontal 

directions perpendicular and parallel to the strike of the 

San Andreas Fault and the vertical direction) consisting 

of 58 x 100 x 56 points with a 0.5-km spacing (29 x 50 x 

28 km) covering the source area. The approximate borders of this travel-time grid are indicated in figure 1. Pand S-wave travel times were then computed from each 

station to every one of the 324,800 grid points to be evaluated as potential earthquake sites during the location procedure. S-wave travel times were computed either using a 

constant vplvs ratio of 1.73 or the three-dimensional vpl 

vs ratio model of Thurber and Atre (1993). 

The PASSCAL instruments recorded both high- and 

low-gain channels of three-component data that triggered 

when the ratio of the short term average amplitude (I s) 

to the long-term average amplitude (30 s) exceeded a value 

of 4.5. Once triggered, 60 s of data were recorded, including a pre-trigger length of 10 s. The data sample rate was 

200 samples per second during the first week of operation, after which it was reduced to 100 samples per second. We read P- and S-wave arrival times from the vertical 

and horizontal components respectively of the high-gain 

channel. When data recorded on this channel were clipped, 

we read arrival times from the low-gain channel. In general, data quality was very high; however, for several 

events, S-wave arrival times could not be determined reliably and these events were not located. We estimate our 

reading errors to be less than 0.05 s and 0.1s for the P and 

S waves respectively. Figure 2 shows examples of P and 

S waveforms from different earthquakes recorded by several of the PASSCAL stations. In many cases, both P and 

S waves are quite complicated; however, initial arrival 

times can be read very accurately from the vertical and 

horizontal components, respectively. It is apparent from 

figure 2 that attempts to read S-wave arrival times from 

the vertical component seismograms would introduce large 

errors, sometimes exceeding 0.2 s. 
LOMA PRIETA AFTERSHOCK RELOCATION WITH S-P TRAVEL TIMES FROM A PORTABLE ARRAY D93 

SYNTHETIC EXAMPLES 

1 The advantages of using S phases to help constrain earth- 

~ quake location have been recognized for some time (for 

example, James and others, 1969; Buland, 1976; and 

Gomberg and others, 1990). Much of the benefit of including S phases in the determination of earthquake hypocenters arises from the substantial increase in the range of 

partial derivatives of travel time with respect to the three 

spatial parameters of the hypocenter compared with P 

I 

I waves. The expanded range of the partial derivative with 

respect to depth can result in a reduction of the trade-off 

between depth and origin time, which plagues location 

schemes that use only P waves (Gomberg and others 1990). 

In situations where exact origin times of earthquakes are 

unnecessary, the use of S-P times for locating earthquakes 

removes the trade-off between depth and origin time, allowing depth determinations to be better constrained. Also, 

we show that S-P arrival times are less sensitive to uncertainties in velocity models than are P and S times separately. However, because uncertainties in both velocity 

structure and timing of arrivals are greater for S waves 

than for P waves, the actual benefit of incorporating S 

waves into earthquake location schemes may be substantially diminished compared with noise-free theoretical expectations. 

To evaluate the effectiveness of using S-P times for 

earthquake location and to approximate the errors expected 

from our earthquake locations, we first performed comprehensive synthetic experiments. The first experiment is 

designed to estimate location errors resulting from uncertainties in the assumed velocity model. Using the same 

receiver geometry as our data and the three-dimensional 

velocity model of Eberhart-Phillips and others (1990b), 

we computed exact P - and 5-wave arrival times from the 

locations we determined for the 50 events shown in figure 

1. We then located these events in a model generated by 

adding Gaussian-distributed random-velocity perturbations 

with an RMS value of 5 percent with scale lengths of 

about 5 km to the model of Eberhart-Phillips and others 

(1990b) and using the following combinations of phases: 

S and P, S-P, and P alone. When using both P- and Swave data to locate earthquakes, we divided the P- and Swave residuals by weights of 1 and vplvs, respectively, 

prior to choosing the location that minimizes the Ll norm 

of the residuals. The weighting was applied to equalize 

Figure 1 .-Map showing sites of PASSCAL stations (triangles) and USGS (Dietz and Ellsworth, 1990) locations 

of the earthquakes analyzed (solid circles). The boundaries of the finite-difference travel-time grid are indicated 

by the outer box. SAF is the San Andreas fault and SF is the Sargent fault. Earthquakes within the inner box are 

the subset of events included in the synthetic calculations. 
D94 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

the constraint imposed on the solution by the S and P 

phases. Figure 3 shows cross sections through the original 

and perturbed three-dimensional velocity models. The inclusion of 5-percent random-velocity perturbations to the 

original velocity model is expected to adequately reflect 

realistic uncertainties in crustal velocities. If the true crustal 

structure contains sharp lateral-velocity contrasts, perturbations to a smoothly varying three-dimensional model 

may not adequately reflect velocity uncertainties. Further 

investigation of this possibility is beyond the scope of this 

paper. 

For several of the simulations, P- and S-wave station 

delays were computed for each station by averaging the 

station residuals from all events, as is often done in real 

earthquake location. These station delays were then subtracted from the arrival-time data and the events were 

located again. The station delays were applied to account 

for imperfections in the assumed velocity model. Although 

the application of station delays had a small effect on the 

event locations and slightly reduced the Ll residuals, it 

did not significantly change the estimates of the location 

errors. Therefore, the results of the synthetic experiments 

do not include station delays. Since the 50 earthquakes 

are distributed over a large area, we also compute location errors from a subset of 19 closely spaced events located near the center of our travel-time grid (figure 1). In 

addition to calculating absolute errors in location resulting from erroneous assumptions about the velocity model 

we also estimate relative errors. Relative errors are determined by calculating the epicentral distances and depth 

DBWN 

WVRD 

10125189 00:27 

hz14.0 krn mag=1.8 

WSWD 

WAWA 

10/30/89 14:04 

h=3.0 krn mag=2.0 

SAOS 

ANNS 

- SEC 

0 2 

Figure 2-Vertical (V) and horizontal (N and E) component seismograms for three typical earthquakes recorded by six PASSCAL stations. P- and 

S-wave arrival times are marked on the vertical components. 5-wave arrival times can be accurately determined from the horizontal components. 

h=earthquake depth and mag=magnitude, both determined by the USGS (Dietz and Ellsworth, 1990). 
LOMA PRIETA AFTERSHOCK RELOCATION WITH S-P TRAVEL TIMES FROM A PORTABLE ARRAY D95 

differences between all events in the original synthetic 

configuration of earthquake locations, subtracting them 

from the distances and depth differences between events 

after new locations are computed and averaging these values. If the relative earthquake pattern does not change 

after the location procedure, then the average of these 

differences will be zero. The average lengths of the absolute and relative horizontal and depth mislocation vectors 

for all 50 events and for the subset of nineteen events are 

summarized in table 1 for the different combinations of 

phases. 

A small component of the error listed in table 1 is due 

to errors in the finite-difference calculation of travel times 

and in the interpolation of travel times prior to earthquake 

location. From a synthetic example with a similar size 

grid, Nelson and Vidale (1990) estimated average location errors of 0.2 km in the horizontal plane and 0.3 km in 

depth arising from the computation of finite-difference 

travel times. 

The use of either S and P or S-P phase data yield average lengths of the horizontal and depth mislocation vectors of about 1.2 km and 0.5 km, respectively. The average 

mislocation increases only slightly when computed for 

events having observations from only four or five stations 

and decreases only slightly when computed for events 

with observations from six or more stations. Only three 

events have both a horizontal mislocation greater than 2 

km and a depth mislocation greater than 1 km. These 

SW 3-D Velocity Model NE 

-15 -10 -5 0 5 I0 

Distance (krn) 

Figure 3-Cross sections through the 3-D model of Eberhart-Phillips and 

others (1990b) and the velocity model generated by introducing Gaussiandistributed random-velocity perturbations, with an average value of 5 

percent over scale lengths of about 5 km, to this model. The cross secevents are recorded by four stations having a highly unfavorable distribution (either lying along a single line or 

clustered into two groups). Thus it appears that given reasonable station coverage, nearly comparable location accuracy can be obtained for events recorded by four or five 

stations as for events recorded by six or more stations 

using either S and P or S-P phase data. Locations calculated using P phases alone produce much larger errors, 

especially in depth. This is not surprising considering the 

locations using only P arrivals have half the number of 

observations as locations using S and P data. The average 

horizontal and depth mislocation for events located with 

P waves from four or five stations is over twice the average of events located using P arrivals from six or more 

stations. As noted by previous workers, the inclusion of 

S-wave phase data drastically reduces the number of stations necessary to obtain accurate earthquake locations. 

The location errors are slightly reduced when only a 

subset of events with fairly uniform station coverage are 

considered (table 1). The average lengths of the relative 

error vectors are larger than the lengths of the absolute 

error vectors when all earthquakes are considered. The 

lack of improvement in the relative errors over the absolute errors is due to the rather large region containing the 

earthquake hypocenters as well as the lack of a consistent 

set of stations recording each event. The average lengths 

of the relative horizontal error vectors are smaller than 

the absolute error vectors for the more closely spaced 

Perturbed Model 

Distance (km) 

tions were taken near the intersection of the San Andreas fault with tne 

Sargent fault. Contours are P-wave velocities in km/s and distances are 

measured from the San Andreas fault (0 km). 
Table 1 .-Synthei 

Phases Error 

Source 

S and P V 

s-P v 

P v 

SandP V&T 

S-P V&T 

AFTERSHOCKS AND POSTSEISMIC EFTECTS 

; location errors 

All Events 

Absolute Errors Relative Errors 

xy sd z sd xy sd z sd 

1.2 0.9 0.5 0.6 1.5 1.3 0.7 0.8 

1.2 0.8 0.4 0.5 1.4 1.2 0.6 0.6 

1.8 1.5 1.8 2.0 2.2 2.2 2.4 2.3 

1.2 1.0 0.5 0.6 1.5 1.4 0.8 0.8 

1.8 2.4 0.8 1.2 1.9 2.2 1.1 1.3 

Subset of Events 

Absolute Errors Relative Errors 

xy sd z sd 1 xy sd z sd 

I 

velocitv model from the one used to I I I 

Error sources result from locating earthquakes in a different 

calculate the phase data (V) and from the addition of noise to the akival time data proportional to 

our estimated reading errors (T). xy=average length of the horizontal error vector, zzaverage 

length of the depth error vector and sd=standard deviation of the error. All errors are in km. 

subset of events; however, the absolute depth errors are 

still smaller than the relative depth errors, 

Although we expected the inclusion of S phases in the 

location determination to greatly improve their accuracy, 

the equal success of S and P and S-P arrival times is 

somewhat surprising considering that the range in the partial derivative of the s-P travel time with respect to the 

hypocentral parameters is substantially decreased compared 

with S or P arrival times (Gomberg and others, 1990). 

The success of S-P arrival times is partly due to the diminished sensitivity of this differential measure to the 

absolute velocity model assumed, Although we located 

the synthetic earthquakes in a velocity model perturbed 

from the one used to compute the arrival-time data, both 

P- and S-wave velocities were perturbed in the same manner using a constant vp/vs ratio of 1.73. If we locate the 

synthetic earthquakes using an S-wave velocity structure 

having a vplvs ratio that differs from that used to compute 

the arrival time data, the error estimates increase in proportion to the difference between the correct and assumed 

ratios. In general, the vplvs ratio is better known than the 

absolute velocity structure, so the use of S-P differential 

times can provide nearly as accurate locations as those 

obtained using S and P arrival times. The magnitude of 

the location errors resulting from a 3-percent difference 

between the assumed and actual vplvs ratios, using S-P 

arrival times, is on the order of 1 km both horizontally 

and in depth. This error does not increase when 5-percent 

random velocity perturbations are added to the original 

model. 

To estimate realistic errors to be expected in our aftershock locations, we repeated the first experiment after 

adding random time delays with an average amplitude 

determined from our estimated reading errors (0.05 s for 

P and 0.1s for S) to the arrival-time data. The results, 

averaging location errors determined for three separate 

simulations of random reading errors, are listed in table 1. 

The inclusion of reading errors had little effect on locations determined using S and P data but did increase the 

errors, as expected, when using S-P or P data. These synthetic calculations indicate that if absolute timing of S 

phases is unavailable, S-P arrival times can locate earthquakes with only slightly less accuracy than the use of S 

and P phases separately. Previous experiments on synthetic datasets including S phases indicated that epicentral 

parameters were far less sensitive to data inadequacies 

than depth (Gomberg and others, 1990). Here, we find 

that the inclusion of S phases in earthquake location results in consistently smaller depth mislocation than epicentral mislocations. 

LOMA PRIETA AFTERSHOCK 

LOCATIONS 

The synthetic calculations indicated that we can expect 

average absolute errors in QUAKE3D locations of the 

Loma Prieta aftershocks to be about 1.8 km in the horizontal plane and 0.8 km in depth using S-P phase data 

and the sparse PASSCAL instrument coverage, provided 

that the average vp/vs ratio is reasonably well known. We 

use the three-dimensional velocity model of EberhartPhillips and others (1990b) to relocate 50 aftershocks of 

the Loma Prieta earthquake. We compare QUAKE3D solutions obtained using this velocity model and a constant 

vplvs ratio to locations obtained by Eberhart-Phillips and 

others (1990b) using only P-wave arrivals. Associated with 

all seismicity maps, we show three cross-sections; one 

parallel to the strike of the San Andreas Fault and two 

perpendicular to it. 

Figures 4 and 5 show locations of 50 aftershocks determined with QUAKE3D using s-P arrival time data from 
LOMA PRIETA AFI'ERSHOCK RELOCATION WITH S-P TRAVEL TIMES FROM A PORTABLE ARRAY D97 

10 20 

DISTANCE [KMI 

Figure 4.-(A) Map view comparing locations determined for the 3-D the two events with very large location discrepancies. Lines X-X', A-A' 

velocity model using S-P finite-difference travel times from the and B-B' indicate the location of cross sections. (B) Cross section X-X' 

PASSCAL stations (x) and locations from Eberhart-Phillips and others along the length of the San Andreas fault. 

(1990b) (solid dots). The events marked by the boxes and triangles are 
D98 AFTERSHOCKS AND POSTSEISMIC EFFEiC'I'S 

the PASSCAL aftershock deployment (x's) and determined work (CALNET). The QUAKE3D locations are indicated 

by Eberhart-Phillips and others (1990b) from P-waves re- in table 2. The average difference in epicentral location 

corded by stations of the dense California Seismic Net- and in the absolute value of depth between the two loca10 20 

D I STANCE [KMI 

Figure 5.-Cross section A-A' (A) and B-B' (B) perpendicular to the San Andreas fault (SAF) showing the seismicity in figure 

4. Sections A-A' and B-B' include all events north and south of the intersection of the SAF with the Sargent fault (SF), 

respectively. 
LOMA PRIETA AFTERSHOCK RELOCATION WITH s-P TRAVEL TIMES FROM A PORTABLE ARRAY D99 

Table 2.-Loma Prieta aftershock relocations 

Event Latitude Longitude Depth Mag N* 

- 

hr mn deg min deg - min km 

1320 37 11.61 121 58.01 2.0 2.0 

891 103 0257 37 10.23 122 0.68 1 I .O I .7 

*N is number of S-P phases used in location 

tEvents with large loiation discrepancies compared with USGS location^. 

tion sets is 2.3k1.4 km and 1.3k1.7 km, if two events with 

large location differences are excluded (these events are 

indicated by triangles and boxes in figure 4). EberhartPhillips and others (199Ob) determined locations for both 

of these events using arrival time data determined with a 

real time processor (RTP), and it is possible that the RTP 

missed an emergent first arrival. One of the events (box 

in figure 4A) was located with P-wave arrival times from 

only nine stations; the smallest number of arrivals for any 

of the 50 events. This event also had the largest rms residual and the largest error estimates of the 50 events. We 

located these two events with observations from only a 

small number of stations with a less-than-ideal distribution. Although our synthetic tests indicated that locations 

using only four or five S-P observations were, in general, 

as accurate as those obtained using many more observations, the largest mislocation vectors determined for our 

tests were obtained for events with observations from few, 

poorly distributed stations. 

Although there are significant differences in the earthquake locations determined in this study and those determined by Eberhart-Phillips and others (1990b), it is 

difficult in such a geologically complex region to apply 

any simple criterion to decide which locations are more 

accurate. Our synthetic experiments predict that locations 

determined with S-P times should be better constrained 

than those obtained from P-wave data if the event-station 

geometries are constant and our assumption and assignment of a constant average vplvs ratio is fairly accurate. 

However, the locations determined by Eberhart-Phillips 

and others (1990b) from P-wave data used many more Pwave observations from more distant stations than the number of nearby S-P observations that we used. Also, recent 

work suggests that our assumption of a constant vplvs 

ratio for this region may not be valid. Thurber and Atre 

(1993) inverted S-P arrival times from Loma Prieta aftershocks recorded by stations of the PASSCAL aftershock 

array to determine a three-dimensional model of vplvs 

variations in the Loma Prieta rupture zone. Their results 

indicate that variations in the vplvs ratio may be as large 

as 10 percent in this region. To determine how a variable 

vplvs ratio would effect our locations, we calculate earthquake locations using the three-dimensional P-wave velocity model of Eberhart-Phillips and others (1990b) and 

the vplvs model of Thurber and Atre (1993). A comparison of these locations with those determined using a constant vplvs ratio of 1-73 are shown in figures 6 and 7. The 

average difference in position of earthquakes located 

using the variable vplvs ratio compared with the constant 

ratio is 1.5k 0.9 and 1.2k 0.9 km in epicentral location 

and in depth, respectively. The depth of events located 

using the variable vplvs ratio is systematically shallower 

than those located using a constant ratio. This is because 

the vplvs ratio (1.73) of the constant model is smaller 

than the average ratio in the variable model. Increasing 

the constant vplvs ratio would increase the focal depths 

and eradicate this depth bias. The use of Thurber and 

Atre's (1993) variable vplvs ratio to locate Loma Prieta 

aftershocks does not reduce the travel time residuals 

compared with a constant vplvs ratio and results in 

earthquake locations that are only slightly closer in location to those determined by Eberhart-Phillips and others 

(1990b). 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

D ISTANCE tKV1 
I 

LOMA PRIETA AFTERSHOCK RELOCATION WITH S-P TRAVEL TIMES FROM A PORTABLE ARRAY Dl01 

4 Figure 6.-(A) Map view comparing locations determined for the 3- 

D velocity model using S-P finite-difference travel times from the 

DISCUSSION AND CONCLUSIONS 

PASSCAL stations and a variable vplvs ratio model (x) with locations 

computed using a constant vplvs ratio of 1.73 (solid dots). (B) Cross Although several previous studies have ddressed the 

section X-X' along the length of the San Andreas fault. sensitivity of earthquake locations to uncertainties in 

assumed velocity models and in the timing of phases, 

these studies have almost exclusively considered layered 

velocity structures only. Changes in the lateral velocity 

A SAF SF 

1111111111111, 

A ' 

10 20 

DISTANCE WW 

Figure 7.-Cross sections A-A' (A) and B-B' (B) perpendicular to the San Andreas fault (SAF) showing the 

seismicity in figure 6 in the same manner as in figure 5. 
D 102 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

gradients will cause raypaths to curve. The effects of ray 

curvature on travel times are diminished when lateral gradients are not allowed, as in the case of propagation 

through a layered structure. This can result in an underestimation of location errors when they are assessed in planelayered media. Here we estimate location error that results 

from random perturbations in a three-dimensional velocity model, although the model we consider possesses only 

smooth velocity gradients. Our methods correctly account 

for ray curvature. We evaluate location errors using the 

event-station geometries provided by Loma Prieta aftershocks recorded at the PASSCAL instruments. An average of only six S-P observations were used in the location 

determinations. The results are summarized in table 1. 

Since earthquake locations are often computed from sparse 

data, an estimation of the location accuracy that can be 

obtained from so few observations is important. We found 

that in the absence of reading errors, S-P arrival-time data 

can locate earthquakes with comparable accuracy as S 

and P data separately. While the inclusion of realistic 

timing errors slightly degrades the accuracy of the S-P 

earthquake locations, the differential measure can still locate earthquakes with small uncertainties (< 2 km in epicentral location and c1 km in depth). 

The relocation of 50 Loma Prieta aftershocks using 

S-P arrival time data from the PASSCAL aftershock deployment are consistent with several other studies indicating complex faulting at depth within the Loma Prieta 

rupture zone. Although the locations suggest the existence of many fault segments, a definitive image of the 

faulting geometry at depth for this structurally complex 

region requires earthquakes to be located with greater accuracy than we can presently achieve. Such accuracy may 

be possible if present uncertainties in the velocity structure can be substantially diminished. 

Rapid response is essential for the success of aftershock 

deployments such as those initiated after the 1988 Armenian and 1989 Loma Prieta earthquakes. The installation, 

testing, and operation of sophisticated clock systems is 

both expensive and time consuming. Therefore, the assurance of accurate absolute timing can significantly decrease 

the amount of data recorded during a field operation. In 

the absence of reliable absolute timing, as was the case 

for the PASSCAL Loma Prieta aftershock data, where 

there were many clock problems, we have shown that 

accurate earthquake locations can be obtained using S-P 

differential travel times from a sparse, three-component 

network. The Loma Prieta earthquake occurred in a region of very dense pre-existing instrument coverage that 

was able to provide accurate earthquake locations for the 

aftershock sequence independent of the PASSCAL deployment. Although the locations we obtained in this study 

using S-P data from the sparse, three-component 

PASSCAL network may or may not have improved the 

previous determinations, their accuracy demonstrates the 

usefulness of S-P times to locate earthquakes in more 

remote regions. In remote locations, such as Armenia, 

field deployments consisting of sparse instrument coverage may be the only means available to obtain accurate 

earthquake locations. In these situations the location 

method discussed here can be very useful. 

ACKNOWLEDGMENTS 

We are grateful to John Vidale for his excellent advice 

throughout the course of this project and for his thorough 

review of this manuscript. We also appreciate and benefited from discussions with Thorne Lay. We thank Cliff 

Thurber, Lynn Dietz and Donna Eberhart-Phillips for reviewing this manuscript and Zhi Zhang for her help with 

data processing. This work was supported by USGS Grant 

no. 14-08-0001-G1836, IRIS Grant SC9Oll5, and a grant 

from the W.M. Keck Foundation. Contribution 180 of the 

W.M. Keck Seismological Laboratory and the Institute of 

Tectonics. 

REFERENCES CITED 

Buland, R., 1976, The mechanics of locating earthquakes: Bull. 

Seismol. Soc. Am., v. 66, p. 73-187. 

Dietz, L.D. and Ellsworth, W.L., 1990, The October 17, 1989, Loma 

Prieta, California, earthquake and its aftershocks: Geometry of the 

sequence from high-resolution locations: Geophys. Res. Lett., v. 

17, p. 1417-1420. 

Eberhart-Phillips, D.M., Labson, V.F., Stanley, W.D., Michael, A.J., 

and Rodriguez, B.D., 1990a, Preliminary velocity and resistivity 

models of the Loma Prieta earthquake region: Geophys. Res. Lett, 

V. 17, p. 11235-1238. 

Eberhart-Phillips, D.M., Michael, A.J., Fuis, G., and Luzitano, R., 

1990b, Three-dimensional crustal velocity structure in the region 

of the Loma Prieta, California, earthquake sequence from inversion of local earthquake and shot arrival times: Seismol. Res. 

Lett., v. 61, p. 48. 

Gomberg, J.S., Shedlock, K.M., and Roecker, S.W., 1990, The effect 

of 5-wave arrival times on the accuracy of hypocenter estimation: 

Bull. Seismol. Soc. Am.,v. 80, p. 1605-1628. 

James, D.E., Sacks, I.S., Lazo, E., and Araracio, G., 1969, On locating 

local earthquakes using small networks: Bull. Seismol. Soc. Am., 

v. 59, p. 1201-1212. 

Lees, J.M., 1990, Tomographic P-wave velocity images of the Loma 

Prieta earthquake asperity: Geophys. Res. Lett, v. 17, p. 1433- 

1436. 

Lemer-Lam, A., Simpson, D., Menke, W., Kim, W.Y., Hough, S., 

Pacheco, J., and Estabrook, C., 1990, Initial results of waveform 

analyses of IRISIPASSCAL Loma Prieta aftershock data: Trans. 

Am Geophys. Union, v. 71, p. 289. 

Nelson, G.D. and Vidale, J.E., 1990, Earthquake locations by 3-D finite-difference travel times: Bull. Seismol. Soc. Am., v. 80, p. 

395-410. 
LOMA PRIETA AFTERSHOCK RELOCATION WITH S-P TRAVEL TIMES FROM A PORTABLE ARRAY Dl03 

Schwartz, S.Y., Velasco, A., Protti, M., Bonarnassa, O., Nelson, G.D., Thurber, C.H., and Atre, S., 1993, Three-dimensional vplvs variations 

McNally, K.C., Vidale, J.E., Lay , T., and Flatt6, S., 1990, A along the Lorna Prieta rupture zone: Bull. Seismol. Soc. Am., v. 

PASSCAL instrument aftershock deployment along the rupture 83, p. 717-736. 

zone of the 18 October 1989 Loma Prieta Earthquake: Trans. Am. Vidale, J.E., 1990, Finite-difference calculation of travel time in 3D: 

Geophys. Union, v. 71, p. 290. Geophysics, v. 55, p. 521-526. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKE OCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

EMPIRICAL GREEN'S FUNCTION STUDY OF LOMA PRIETA 

AFTERSHOCKS: DETERMINATION OF STRESS DROP 

By H. Guo, A. Lerner-Lam, and W. Menke, 

Lamont-Doherty Earth Observatory of Columbia University, and 

S.E. Hough, 

U. S. Geological Survey 

CONTENTS 

Page 

Dl05 

105 

105 

108 

lo8 

115 

15 

115 

119 

119 

ABSTRACT 

This paper presents the estimation of stress drops of 

aftershocks of the Loma Prieta earthquake recorded by an 

array of 22 PASSCAL instruments. Using an empirical 

Green's function analysis, corner frequencies of 98 of the 

aftershocks are first determined. Brune-model stress drops, 

ACT, are calculated from the corner frequencies and moments. A dependence of stress drop on moment is observed. 

INTRODUCTION 

Estimates of earthquake source spectral parameters can 

be contaminated by effective attenuation along the sourcereceiver path, frequency-dependent propagation, and nearreceiver site response. The shaping of the spectrum of an 

arrival by attenuation is known to bias estimates of corner 

frequency and thus its interpretation in terms of stress 

drop (Hough and Anderson, 1988). To avoid this bias, a 

reference measurement technique must be employed when 

parameterizing the source spectrum, especially in a local 

aftershock survey. In this paper, we apply an empirical 

Green's function (eGf) analysis (Frankel and others, 1986; 

Geller and Mueller, 1980; Hartzell, 1978; Hough and others, 1991 ; Hutchings and Wu, 1990; Li and Thurber, 1988; 

Mori and Frankel, 1990; Mueller, 1985; Xie and others, 

1991) to aftershock doublets of the Loma Prieta sequence 

recorded on the IRISIPASSCAL digital seismographs. The 

smaller event in each doublet provides the reference signal for correcting the spectral characteristics of the larger 

source for spectral shaping by attenuation. 

In principle, eGf methods provide an efficient technique 

for quickly parameterizing large amounts of digital data. 

In practice, however, one has to be careful about the application of the method in the case of inadequate signalto-noise ratio for the smaller reference signal. We develop 

a methodology for identifying and mitigating such lowsignal problems and show how to assess the possibility of 

bias in the corner-frequency estimate. In addition, we perform sensitivity tests to demonstrate that the estimate of 

these spectral parameters is robust with respect to variations in fitting range and noise level. 

The combination of extensive coverage of the Loma 

Prieta aftershock zone with IRISIPASSCAL instrumentation, coupled with the locating capabilities of CALNET, 

provides a unique data set for waveform work and eGf 

analysis of spectral characteristics. 

In this paper, a frequency-domain deconvolution using 

the empirical Green's function method is used to determine the corner frequencies and the stress drops of 98 

aftershocks of the Loma Prieta earthquake. 

DATA 

The Loma Prieta earthquake and its aftershocks are an 

exceptionally well-recorded earthquake sequence. The aftershocks were recorded not only by the permanent 

CALNET stations but also by a temporary network. Following the Loma Prieta mainshock, an array of 22 IRIS1 

PASSCAL (Incorporated Research Institutions for SeismologylProgram for Array Seismological Studies of the 
Dl06 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Continental Lithosphere) stations was deployed in the epicentral area in the Santa Cruz mountains (see fig. 1). More 

than 2,000 events were recorded by three-component L22D 2-Hz velocity sensors paired with IRISIPASSCAL 

Reffek model 72A recorders. Two separate geophones, 

recorded at different gain levels and sampled at 200 

samples per second, were deployed at each site. These 

geophones were calibrated after the Loma Prieta deployment by Menke and others (1991). 

A total of 763 of these aftershocks, with local magnitude, Mb ranging from 1.5 to 4.5 and recorded from October 20 to November 25, 1989, have been time-associated 

with CALNET data. These associated aftershocks were 

relocated relative to preliminary CALNET hypocenters 

by inverting CALNET phase data with a one-dimensional 

velocity model and a modified HYPOINVERSE program 

with station corrections dependent on hypocentral location (Seeber and Armbruster, 1990). This method has been 

shown to give better relative locations (Seeber and 

Armbruster, 1990; Pujol, 1992). 

Candidate pairs for eGf analysis were selected from the 

relocated hypocenters. The fundamental assumptions of 

eGf analysis require that (1) the paired events must be 

closely spaced so that source-receiver paths are nearly 

coincident, (2) the event pair should have similar focal 

mechanisms and have similar waveform complexities, and 

(3) the rupture duration of the smaller or reference event 

should be short relative to the larger event so that the 

reference corner frequency is resolvably different from 

the corner frequency of the large event. This last criterion 

has implications in terms of signal-to-noise ratio of the 

quotient spectrum. It is most convenient to interpret this 

requirement in terms of relative magnitude difference, AMy 

The Mivalues must be large enough to resolve the difference in corner frequencies but not so large as to vitiate 

the assumption that attenuation and near-receiver effects 

are linear for two events. 

Our criteria for the relocated Loma Prieta aftershocks 

were (1) the hypocenters must be within 1 km of each 

other, (2) time- and frequency-domain signals must be 

122.0 

LONGITUDE, DEGREE 

Figure 1.-Portion of the aftershock zone of the Loma Prieta earthquake. Main shock (diamond), 22 PASSCAL stations (circles), 53 closest 

CALNET stations (triangles), San Andreas Fault (SAP), and Sargent fault (SF) are shown. A total of 124 events recorded by the PASSCAL array 

with magnitudes larger than 1.5 are used for empirical Green's function analysis and denoted by plus symbols. 
EMPIRICAL GREEN'S FUNCTION STUDY OF LOMA PRIETA AFTERSHOCKS: DETERMINATION OF STRESS DROP Dl07 

visually similar, and (3) AML>0.7. Comparable values of 

this lower bound for AMLhave been used in other studies 

(Li and Thurber, 1988: 0.5; Mueller, 1986: 0.7; Xie and 

others, 1991: 0.5). The largest AML for out data set is 2.9, 

which is small enough to maintain linearity. Of these criteria, the visual check is the most subjective. In the time 

domain, the comparison emphasizes the first motion direction, the detailed shape of the first several oscillations, 

and the overall shape of the envelope of the whole seis? 

mogram (for example, fig. 2). In the frequency domain, 

the comparison emphasizes spectral content below 30 Hz 

because of spurious geophone resonances detected by 

Menke and others (1991). Only those event pairs that satisfy the above criteria on 80 percent of the seismograms 

are suitable for further analysis. Seismograms with recording errors or signal-to-noise ratios less than approximately 20 dB were discarded. Low-gain data were used 

only if the signals were clipped on the high-gain data 

streams. 

Of the 763 relocated aftershocks with AMy>1.5, 347 

pairs were found to meet the hypocentral and AML restriction, and 226 event pairs were found to meet the subjec0 2 4 6 8 10 

TIME, SECOND 

Figure 2.-Three-component seismograms of a typical event pair (events 107 and 106). Amplitude scales differ between seismograms. 
Dl08 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

tive waveform similarity criteria. The larger events in each 

pair have magnitudes between 2.2 and 4.4, and the smaller 

events have magnitudes between 1.5 and 2.9. where A is the amplitude scaling and f is the corner 

frequency. The spectral domain deconvolution yields the 

spectral ratio RO: 

ANALYSIS 

DETERMINATION OF CORNER FREQUENCY 

The P-wave spectra were computed for a 2.5-second 

window starting just before the arrival. The length of the 

window was derived after experimentation to determine 

the trade-off between the successful characterization of 

the long-period signal and the inclusion of too much scattered energy. The spectra were computed using 47t-prolate multitapers, which yield a smooth spectral estimate 

with good resistance to leakage for relatively short time 

windows (Park and others, 1987). 

The paired spectra were divided to obtain spectral ratios, which should depend only on the relative source characteristics. Spectral ratios from different components and 

different stations were averaged to reduce the noise. This 

averaging procedure neglects the effects of directivity, 

which is not thought to be a large effect. 

The parameterization of the spectral ratio in terms of 

source characteristics is achieved by assuming a Brune 

spectrum, A(/), for each event (see fig. 3): 

0 l-- I I I I I I Ill I I I I IIIIJ 

0 1 2 

LOG FREQUENCY, HZ 

Figure 3. Comer frquencies of a hypothetical event pair, showing spectra as a function of frequency for large event (top), small event (middle), 

and ratio (bottom). fC=5 Hz and fC=20 Hz are the comer frequencies of 

the larger and smaller events, respectively. 

where subscripts 1 and 2 (oJ, 02) denote the larger and 

smaller members of the event doublet, respectively. Unless f is high enough relative to f, or A, is below the 

ambient noise, all four parameters must be included in the 

representation of the spectral ratio. We do this by finding 

the amplitude ratio Aol/Ao2 that minimizes the 2-norm of 

the residual R'i-RCf), where R'w is the observed spectral ratio, while fc and fc2 are varied on a grid. The values 

of fC and fC that yield the smallest residual are then taken 

to be estimates of the event corner frequencies. The corner frequencies are first gridded at 1-Hz increments between 1 Hz and a maximum of 100 Hz and 0.1-Hz 

increments for lower corner frequencies. R(f) is fitted between the instrument corner at 2 Hz and an upper noisefree frequency, which is determined independently for each 

pair. Poor choice of the fitting range yields obvious instabilities. The morphology of the residual manifold is assessed visually for each event pair studied. 

Two examples of this procedure, displayed in figure 4, 

show that this methodology can be problematic under certain circumstances. Figure 4A shows the residuals for the 

event pair (20, 24). In this case, the corner frequencies 

(f, fc)=(3, 16 Hz) are well-resolved. This figure also 

shows the bias that is possible if we assume the smaller 

event to be a delta function and the search is gridded only 

over the corner frequency fcl. Figure 4B shows the residuals for the event pair (58, 60). In this case no closed 

residual low is observed, and the corner frequency of the 

smaller event is not resolved at all. However, f has a 

small variation relative to the poor resolution of fC. This 

is generally true for event pairs where the smaller event 

magnitude is less than 2 or where the ambient noise is 

high. In order to get the best estimate off, we fix fC to a 

value consistent with the average value found in previous 

studies (Molnar and others, 1973; Boatwright and others, 

1991; Fletcher and Boatwright, 1991). Because fC in this 

case is assumed, it is indicated with a blank arrow in 

figure 5 and excluded from the averaging. 

Of the 226 pairs analyzed, 102, which include 124 individual events, are fitted well by the Brune ratio model, 

shown in figure 5. Averaging of common events and further exclusion of artifacts yielded 98 individual measurements of P-wave corner frequencies, with 53 from the 

larger events (see table 1). These corner frequencies fall 

in the range 3.8 to 37.5 Hz, with corresponding events 

having local magnitudes between 1.5 and 4.4. 
EMPIRICAL GREEN'S FUNCTION STUDY OF LOMA PRIETA AFTERSHOCKS: DETERMINATION OF STRESS DROP Dl09 

0 10 20 30 40 50 

CORNER FREQUENCY Fc1 

o~'ll~lllllll,ll ,1,1~~ 

0 10 20 30 40 50 60 70 80 90 100 

CORNER FREQUENCY Fc1 

Figure 4.-Typical least-squares residuals as functions 

of fC and fC. White represents the minimum residual, 

and black indicates five times the minimum residual. 

Variation is linear in logarithmic residual, giving 256 

grey scales. (A) Fitting for pair 20124. Low residuals are 

centered in a very narrow frequency range and give us 

well-resolved corner frequencies of fC=3 Hz and f=16 

Hz. (B) Fitting for pair 58/60. The residual contour is 

not closed. The residual increases with increasing f9, 

but fC changes less with increasing fc. An artifact limit, 

30 Hz, is assumed for fc2, and then fC=8 Hz is resolved 

in this case. 
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EMPIRICAL GREEN'S FUNCTION STUDY OF LOMA PR IETA AFTERSHOCKS: DETERMINATION OF STRESS DROP Dl 15 

CALCULATION OF STRESS DROP 

The determination of stress drop for a small earthquake 

is generally based on a source determination derived from 

either the corner frequency of the displacement spectrum 

or the pulse width of the waveform. We model a small 

earthquake as a circular rupture and estimate the stress 

drop from the corner frequency, fcl, using the formula 

(Brune, 1970; Brune, 197 1 ; Kanamori and Anderson, 

1975): 

where the rupture velocity v is taken to be 3.5 kmls 

(Dahlen, 1974; Xie and others, 1991; Hough and others, 

1991). 

The seismic moment, Mo, is calculated for each event 

(table 1) from the long-period displacement spectral levels using 

where a is the P-wave velocity, p is the density, r is the 

hypocentral distance, Sl is the long period spectral level, 

F is the free surface correction, and R is the average 

correction for radiation pattern. We use 0=6 kds, p=2.7 

gdcm3, F=2, and R-0.7. We also calculated moment 

taking into account exact radiation terms as determined 

from the focal mechanisms, but found these estimates to 

be less consistent (among moment estimates for the same 

event at different stations) than those obtained with a constant radiation term. The corner frequency is shown as a 

function of moment in figure 6. 

STRESS-DROP RESULTS 

The Brune-model stress drops are found to be highly 

variable, ranging from approximately 1 to nearly 800 bars 

(see table 1 and figure 7). The dashed lines on figure 7 

indicate the resolution limits corresponding to the frequency range over which we can generally resolve corner 

frequencies, 1 to 40 Hz. It appears that the apparent upper 

trend in figure 7 may in fact be an artifact of the high 

frequency resolution limit. However, the methodology and 

data should be sufficient to resolve low-stress-drop larger 

events, if they existed. 

A dependence of stress drop on moment can be interpreted in terms of a critical slip distance resulting from 

fundamental characteristics of established rate and state 

variable friction laws (for example, Dieterich, 1986). The 

critical slip distance to achieve dynamic rupture corresponds to a minimum earthquake rupture dimension (for 

example, Dieterich, 1986; Scholz, 1990), thus leading to a 

breakdown in the similarity of earthquake rupture processes. 

It is also possible to explain the inferred stress drop 

scaling in terms of tectonic factors. Global compilations 

(Abercrombie and Leary, 1993) show no resolved systematic scaling of stress drop with moment, implying that 

globally, there is no preferred earthquake rupture dimension (in other words, breakdown in similarity). However, 

within a limited tectonic region, it is plausible that a characteristic length scale will exist. If, for example, aftershocks in the Loma Prieta segment are interpreted as 

activation of cross faults within the fault zone (L. Seeber, 

1994, unpub. data), then their rupture dimensions may 

preferentially correspond to the width of the fault zone. 

CONCLUSIONS AND DISCUSSION 

It has been demonstrated that eGf analysis is a quick 

and effective method for estimating the source characteristics of a large number of small- to medium-sized earthquakes with good resolution. The frequency-domain 

deconvolution removes the effects of travel path, site, and 

the instrument and isolates the source characteristics of 

the two events. The Brune spectral ratio model appears 

sufficient to parameterize the spectral ratio, provided that 

the differences in magnitude are not too small and the 

ambient noise at high frequencies is not too large. It is 

important, moreover, to specifically address the properties of the fitting residual as a function of the gridding 

variables. The estimate of the large-event corner frequency 

can be biased by improper selection of the small-event 

corner frequency, and so it is important to grid additionally over small-event corner frequency. The grid searching yields a well-resolved estimate of the corner frequency 

of the large event and sometimes of the smaller event as 

well. The resolution of the corner frequency of the larger 

event will be much better than that of a single spectrum 

parameterization of source and path effects if we properly 
Dl 16 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 1 .-The 124 events used for empirical Green's function analysis 

[The 124 events involved in the 102 event pairs shown in figure 5. Of these, 98 event pairs have resolved corner frequencies, with 53 from the larger events (indicated in 

column L). The other 26 events are generally smaller events with magnitudes between 1.5 and 2.0. Their comer frequencies are not well-resolved from the ratio fitting 

and are excluded from the corner frequency averaging and the calculation of stress drops.] 

Event Date Time Long. Lat. Depth ML Moment fc Stress L 

(yymdd (Gmt (W) (N) (Km) (dyne*cm) (Hz) (bar) 
EMPIRICAL GREEN'S FUNCTION STUDY OF LOMA PRIETA AFTERSHOCKS: DETERMINATION OF STRESS DROP Dl 17 

Table 1 .-Continued 

Event Date Time Long. Lat. Depth ML Moment fc Stress L 

( ~~rnmdd (Gmt ) (W) (N) (Km) (dyne*cm) (Hz) (bar) 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 1 .-Continued 

Event Date Time Long. Lat. Depth ML Moment fc Stress L 

( yymdd (Gmt ) (W) (N) (Km) (dyne*cm) (Hz) (bar) 

100 

101 

102 

103 

104 

105 

106 

107 

108 

109 

110 

Ill 

112 

113 

114 

115 

116 

117 

118 

119 

120 

121 

122 

123 

124 

LOG MOMENT, DYNE*CM Figure 7.-Brune-model stress drop as a function of seismic moment for 

Figure 6.-Corner frequency as a function of seismic moment for 98 98 Loma Prieta aftershocks. Circles indicate the larger events and triLorna Prieta aftershocks. Circles indicate the larger events and triangles angles indicate the smaller events. The dashed lines are the resolution 

indicate the smaller events. limits corresponding to fixed corner frequencies of 1 Hz and 40 Hz. 
EMPIRICAL GREEN'S FUNCTION STUDY OF LOMA PRIETA AFTERSHOCKS: DETERMINATION OF STRESS DROP Dl 19 

choose the event pair. Although the empirical Green's 

function method is limited to available event pairs, we 

have shown that, within a dense aftershock sequence, it 

can be used for a large-scale investigation of source properties. In this study, we obtain a total of 98 well-resolved 

corner frequencies for events with magnitudes between 

1.5 and 4.4. 

ACKNOWLEDGMENTS 

We thank the many people who contributed to the successful deployment of the IRISIPASSCAL network, including David Simpson, Larry Shengold, Robert Busby, 

and James Fowler. We thank Professor Lynn Sykes for 

providing helpful comments and suggestions. We thank 

Leonardo Seeber and John Armbruster for providing relocation data. We also thank David Oppenheimer for supplying CALNET phase data used for relocation. The 

research was supported by U.S. Geological Survey under 

Grant 14-08-001 -G- 183 1, and the aftershock data collection was supported by the Incorporate Research Institutions for Seismology and by the National Science 

Foundation. The PASSCAL aftershock data are available 

from the IRIS Data Management Center or from the authors at Lamont-Doherty. This is Lamont-Doherty contribution number 5479. 

REFERENCES CITED 

Abercrombie, L. and P.C. Leary, 1993, Source parameters of small 

earthquakes recorded at 2.5 km depth, Cajon Pass, California: Implications for earthquake scaling: Geophysical Research Letters 

20, p. 1511-1514. 

Brune, J.N., 1970, Tectonic stress and the seismic shear waves from 

earthquakes: Journal of Geophysical Research, v. 75, no. 26, p. 

4997-5009. 

Brune, J.N., 1971, Correction: Journal of Geophysical Research, v. 76, 

no. 20, p. 5002. 

Boatwright, J., Fletcher, J.B., and Fumal, T.E., 1991, A general inversion scheme for source, site, and propagation characteristics using 

multiply recorded sets of moderate-sized earthquakes: Seismological Society of America Bulletin, v. 81, no. 5, p. 1754-1782. 

Dahlen, A., 1974, On the ratio of P-wave to S-wave corner frequencies 

for shallow earthquake source: Seismological Society of America 

Bulletin, v. 64, no. 4, p. 1 159- 1 180. 

Dieterich, J., 1986, State variable fault constitutive relations for dynamic slip, in Earthquake Source Mechanics: American Geophysical Union Geophysical Monograph 37, p. 3 1 1-3 18. 

Frankel, A., Fletcher, J., Vernon, F., Haar, L., Berger, J., Hanks, T., 

and Brune, J., 1986, Rupture characteristics and tomographic 

source imaging of ML-3 earthquakes near Anza, Southern California: Journal of Geophysical Research, v. 91, no B12, p. 12633- 

12650. 

Fletcher, J.B., and Boatwright, J., 1991, Source parameters of Loma 

Prieta aftershocks and wave propagation characteristics along the 

San Francisco peninsula from a joint inversion of digital seismograms: Seismological Society of America Bulletin, v. 81, no. 5, p. 

1783-1812. 

Geller, R.J., and Mueller, C.S., 1980, Four similar earthquakes in central California: Geophysical Research Letters, v. 7, no. 10, p. 821- 

824. 

Hartzell, S.H., 1978, Earthquake aftershocks as Green's functions: 

Geophysical Research Letters, v. 5, no. 1, p. 1-4. 

Hough, S.E., and Anderson, J.G., 1988, High frequency spectra observed at Anza, California: Implications for Q structure: Seismological Society of America Bulletin, v. 78, no. 2, p. 692-707. 

Hough, S.E., Seeber, L., Lerner-Lam, A., Armbruster, J.G., and Guo, 

H., 1991, Empirical Green's function analysis of Loma Prieta aftershocks: Seismological Society of America Bulletin, v. 81, no. 5, 

p. 1737-1753. 

Hutchings, L., and Wu, F., 1990, Empirical Green's functions from 

small earthquakes: a waveform study of locally recorded aftershocks of the 1971 San Fernando earthquake: Journal of Geophysical Research, v. 95, no. B2, p. 1 187-1214. 

Kanamori, H., and Anderson, D.L., 1975, Theoretical basis of some 

empirical relations in seismology: Seismological Society of 

America Bulletin, v. 65, no. 5, p. 1073-1095. 

Li, Y., and Thurber, C.H., 1988, Source properties of two microearthquakes in Kilauea volcano, Hawaii: Seismological Society 

of America Bulletin, v. 78, no. 3, p. 1123-1 132. 

Menke, W., Shengold, S.L., Guo, H., Hu, G., and Lerner-Lam, A., 

1991, Performance of the short-period geophones of the IRIS/ 

PASSCAL array: Seismological Society of America Bulletin, v. 

81, no. 1, p. 232-242. 

Molnar, P., Tucker, B.E., and Brune, J.N., 1973, Corner frequencies of 

P and S waves and models of earthquake sources: Seismological 

Society of America Bulletin, v. 63, no. 6, p. 2091-2104. 

Mori, J., and Frankel, A., 1990, Source parameters for small events 

associated with the 1986 North Palm Springs, California, earthquake determined using empirical Green's functions: Seismological Society of America Bulletin, v. 80, no. 2, p. 278-295. 

Mueller, C.S., 1985, Source pulse enhancement by deconvolution of 

empirical Green's function: Geophysical Research Lettters, v. 12, 

no. 1, p. 33-36. 

Park, J., Lindberg, C.R., and Vernon, F.L., 1987, Multitaper spectral 

analysis of high-frequency seismograms: Journal of Geophysical 

Research, v. 92, no. B12, p. 12,675-12,684. 

Pujol, J., 1992, Joint hypocentral location in media with lateral velocity variations and interpretation of station correction: Physics of 

the Earth and Planetary Interiors, v. 75, no. (1-3), p. 7-24. 

Scholz, C.H., 1990, The mechanics of earthquakes and faulting: Cambridge, Massachusetts, Cambridge University Press, 441 p. 

Seeber, L. and Armbruster, J.G., 1990, Fault kinematics in the 1989 

Loma Prieta rupture area during 20 years before that event: Geophysical Research Letters, v. 17, no. 9, p. 1425-1428. 

Xie, J., Liu, Z., Herrmann, R.B., and Cranswick, E.D., 1991, Source 

processes of three aftershocks of the 1983 Goodnow, New York, 

earthquake: high-resolution images of small, symmetric ruptures: 

Seismological Society of America Bulletin, v. 81, no. 3, p. 818-843. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

CONTENTS 

Abstract ............................................................... 

Introduction ............................................................... 

GEOS data ............................................................... 

Large-length-scale seismic-wave-propagation studies ------ 

Influence of local geology on ground shaking --------------- 

Collocations with strong-motion instruments ---------------- DR-200 data .............................................................. 

Santa Cruz .......................................................... 

Moss Landing power plant ...................................... 

LO~ Gates ........................................................... 

Liquefaction sites .................................................. 

Topographic effects studies ..................................... 

Collocations with strong-motion instruments ---------------- USGS seismic cassette recorders ..................................... 

Acknowledgments ....................................................... 

References ............................................................... 

By Leif Wennerberg, 

U.S. Geological Survey 

INTRODUCTION 

ABSTRACT 

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Portable seismographs were deployed to record aftershocks of the Loma Prieta earthquake for studies of earthquake source processes, effects of local geology on ground 

motions, and crustal and basin wave-propagation effects. 

This report provides maps and descriptions of specific 

experimental objectives of the aftershock deployments. 

Thirty-eight digital seismographs were deployed at 195 

sites during the period from the day after the October 17, 

1989, main shock until mid-January 1990. Aftershock recordings were obtained at or near 26 sites where the main 

shock was recorded by permanent strong-motion instruments. Fifteen aftershock recording instruments were at, 

or near, locations where previous USGS studies have been 

done; nuclear-explosion recordings and/or borehole geologic and velocity data are available for these sites. There 

was also a brief deployment of 60 USGS refraction-survey analog instruments in the epicentral area on the night 

after the main shock. 

Aftershock sequences provide a reliable source of earthquake ground motions which can be used to address issues relating to the physics of the seismogenic process 

and to address practical problems of strong-ground-motion prediction, such as the relevence of weak-motion 

records to strong-ground motions or the effects of local 

geologic conditions on ground motion. Following the Loma 

Prieta earthquake, several groups from the USGS recorded 

aftershocks in an effort to further our understanding of 

earthquake hazards in the San Francisco Bay area. The 

digital data sets include a total of more than 5,000 threecomponent records from 195 sites in the greater San Francisco Bay area, and have been collected on two CD-ROMs 

(see Mueller and Glassmoyer, 1990, and Carver and others, 1990, for data access). 

In this paper I summarize the USGS data-collection 

efforts and describe the experimental objectives. Certain 

recording sites could be used for more than one purpose, 

some of which may not have been considered, so the objectives listed here cannot be exhaustive. As I discuss the 

experiments, I refer to any published results I know of 

that have used these data. I do not discuss the results of 

these studies. In the "References" section I also list a few 

analyses not referred to explicitly in the text. There is still 

a significant amount of work that could be done using 

these data. 

This discussion provides an overview and introduction 

to the data sets, and supplements Mueller and Glassmoyer's 

(1990) and Carver and others' (1990) presentations of the 

digital data. No report has been written specifically devoted to the analog data set. Mueller and Glassmoyer's 

report (1990), while providing a comprehensive description of a significant portion of the digital data set, provides a much briefer discussion of experimental objectives 

than presented here. It contains no discussion of previous 

work or published analyses. My discussion of the data 

presented in Carver and others (1990) briefly summarizes 

their report and includes references to relevant work that 

has appeared since their report was published. 
Dl22 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

For the purposes of this report, I divide the USGS aftershock investigations into three parts corresponding to 

three instrumentation groups, one of which collected only 

a small and quite specialized data set, but which is included for the sake of completeness. The two main datacollection efforts were undertaken by a group using GEOS 

recorders (Borcherdt and others, 1985; Mueller and 

Glassmoyer, 1990) and a group using Sprengnether DR200 recorders (Carver and others, 1986; Carver and others, 1990; Cranswick and others, 1990). The data collected 

by the GEOS group are mostly from Santa Clara County 

and the San Francisco peninsula, with some data recorded 

in Monterey, San Benito, Santa Cruz, and southern 

Alameda Counties. These data are described in detail in 

Mueller and Glassmoyer (1990). The DR-200 data are 

mostly from Santa Cruz and Santa Clara Counties, with 

some records obtained in Monterey County and are described in Carver and others (1990) and King and others 

(1990). The GEOS records digital data with 16-bit resolution. The DR-200 has a 12-bit digitizer, but is gain-ranged 

to give an effective 16-bit recording range. Both types of 

instruments sampled at 200 samples per second. Threecomponent velocity transducers with natural periods between 1 and 2 Hz were used. Most GEOS recorders were 

also deployed with three-component force-balance accelerometers. See Mueller and Glassmoyer (1990) and Carver 

and others (1990) for details on sensors. 

In this paper I discuss the GEOS experiments, presenting station maps and a table of sites and experiments, and 

note related work done by the USGS. Then I discuss the 

DR-200 data, presenting maps and discussing experiments 

and selected sites. Finally, I briefly describe the refraction-survey-instrument data set. 

GEOS DATA 

Twenty-seven GEOS recorders were deployed at 94 sites 

during the aftershock studies. 2,615 three-component velocity records were collected, and 2,592 accelerograms 

were collected from sensors collocated with velocity transducers. Figures 1 through 8, taken from Mueller and 

Glassmoyer (1990), show the sites occupied by GEOS 

recorders. Table 1 lists the sites where data were obtained 

and indicates the reason(s) for each site's occupation. The 

main shock was recorded by permanent strong-motion instruments at twenty-three GEOS sites. Mueller and 

Glassmoyer (1990) give more detailed information on site 

locations and characteristics, instrument parameters, and 

aftershocks recorded. The following is a somewhat idealized description of the aftershock studies undertaken. 

Mueller and Glassmoyer (1990) give a clearer indication 

of how much data relevant to each study were actually 

obtained. 

The topics of interest during the GEOS deployments 

can be placed in three categories: large-length-scale wavepropagation effects, local site effects, and collocations with 

strong-motion instruments. The large-scale studies were 

intended to clarify crustal propagation effects, or the possibility of a distinct seismic response of the sedimentary 

basin of the south San Francisco Bay. Studies of local site 

effects were primarily motivated either by damage or by 

an interest in the shaking properties of common geologic 

units. These studies typically consisted either of instruments deployed on bay mud and/or alluvium sites with 

simultaneously recording instrument(s) "nearby" on one 

or more rock sites, or of dense arrays designed to study 

wave propagation and (or) the small-scale variability of 

ground motions. The collocations with strong-motion instruments were intended to collect small-earthquake data 

that could be used for an empirical-Green's-function analysis of the main shock (for example, Hartzell, 19781, or for 

information on the relationship between small-earthquake 

records and large-earthquake ground motions. 

LARGE-LENGTH-SCALE SEISMIC-WAVEPROPAGATION STUDIES 

The main large-scale deployment was a line of stations 

along the San Francisco peninsula from the aftershock 

zone to San Francisco. The stations are listed in table 1 

under the experiment PP. The main interest of this deployment was crustal wave propagation, with a secondary 

interest in local site response. Results of analysis of the 

data are presented by Fletcher and Boatwright (1991) and 

Boatwright and Fletcher (1991a). McGarr and others 

(1991) also used some of these stations to analyze wave 

propagation up the peninsula. Two shorter transverse lines, 

oriented roughly southwest-northeast, were deployed in 

the south bay: a line from the southern peninsula to the 

southern east bay flatlands (from SRL on the west to FLB 

and POR on the east, fig. 4) and a line crossing the northern Santa Clara Valley (from BMT on the west to TUL on 

the east, fig. 4). The stations involved are listed in table 1 

as PE and SV, respectively. These lines were deployed to 

observe possible basin response. 

INFLUENCEOFLOCALGEOLOGYONGROUND 

SHAKING 

Most of the aftershock studies focused on the effects of 

local geology on ground motion. Two particularly significant efforts were studies of the Marina District of San 

Francisco and of the San Francisco Airport area (Mb and 

Ms, and SF0 in table 1; figs. 7 and 5, respectively). The 

Marina suffered dramatic damage, which was surprising 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA D 123 

given its distance from the main shock, and was the object of quite thorough instrumentation. The USGS was 

requested by the airport authorities to conduct a study to 

interpret damage patterns at the facility. Analyses of the 

Marina data have appeared in Boatwright and others 

(1991b, 1992). McGarr and others (1991) analyzed the 

airport data. 

A study of San Francisco (SF in table 1) focused on the 

correlation of site amplification and damage with geology 

in the city. An analysis of these data appears in Seekins 

and Boatwright (1994). Instruments were also deployed 

near Foster City on the bay margin (FC in table 1, fig. 4). 

These data have been discussed in Celebi and McGarr 

(1991a, b). The trans-south-bay line (PE in table 1, fig. 4) 

was also designed to sample a variety of alluvial units to 

determine if significant variations in ground shaking occurred. These data were analyzed by Margheriti and others (1994). Six instruments were deployed on or near the 

Figure 1 .-GEOS station locations (dots) in greater San Francisco Bay area. Main faults are indicated by solid and broken lines. 

Approximate Lorna Prieta aftershock zone is indicated by shaded area. Boxes outline detail maps in figures 2, 3, 4, and 5. Map 

from Mueller and Glassmoyer (1990). 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Oakland Bay Bridge in an effort to gather data relevant to 

the bridge closure (BB in table 1, fig. 6). Many instruments were located at sites of particular interest due to 

damage or geologic conditions. These site features are 

described under the heading U in table 1. 

Fifteen instrument locations were the same as, or less 

than a kilometer away from, sites that have been studied 

previously by the USGS; one location was the site of a 

small array (STl+, figs. 4, 8). These are referred to as Old 

sites in table 1. The earlier studies began with comparisons of the response of different sites around the bay to 

nuclear explosions at the Nevada Test Site (Borcherdt, 

1970; Gibbs and Borcherdt, 1974; Borcherdt and Gibbs, 

1976) and included geologic and velocity logs to 30 meters 

in depth at selected sites (Gibbs and others, 1975, 1976a, 

and 1976b). A data analysis indicating that the velocitylog data were useful for interpreting the amplitudes of the 

explosion data was given by Borcherdt and others (1978). 

A deeper (180 m) borehole was drilled at station RAV, 

which has been the subject of quite detailed analysis 

(Warrick, 1974, and Joyner and others, 1976). Table 2 

lists the aftershock-instrument sites that are near to (significantly less than a kilometer), or the same as, sites 

where nuclear explosions were recorded. It also lists specifically the correspondence between Borcherdt and 

Gibbs's recordings and the aftershock-instrument sites. The 

GEOS sites BMT and COY are noteworthy because they 

were also reference sites for the earlier work. 

Also listed in table 2 are sites for which near-surface 

velocity data are available. Most of these sites were used 

in the nuclear explosion work. GEOS stations AP7, MAL, 

and SF1 were not included in the nuclear studies but are 

the sites of strong-motion accelerometers that recorded 

the main shock. Recently, a 150-m velocity log was obtained at station SF1 (Gibbs and others, 1992). In addition 

to a borehole log at SWS (Kayen and others, 1990), simultaneous downhole/uphole small-earthquake data have 

been recorded there (Liu and others, 1992a, b). 

Four sub-kilometer-scale arrays recorded data relevant 

to spatial variations in ground motions. Three instruments 

were deployed in the Santa Cruz mountains as part of an 

attempt to observe topographic effects (SC in table 1 and 

KKO+ in figs. 3, 8, 14). This deployment was concurrent 

with DR-200 instruments, as discussed below. Two arrays 

Figure 2.-GEOS stations in Monterey and San Benito Counties (southernmost region in figure 1). Main shock was recorded at MON, SAG, and 

TOM. Map from Mueller and Glassmoyer (1990). 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA Dl25 

were intended to study the propagation of surface waves 

in deeper sediments: a three-component array deployed in 

Palo Alto (SP in table 1, DUV+ in figs. 4, 8) and a fourcomponent array was deployed in the north central Santa 

Clara Valley (SA in table 1, AG1+ in figs. 4, 8). A fourcomponent array was deployed near Stanford University, 

mainly as a test of array instrumentation techniques (ST 

in table 1, ST1+ in figs. 4, 8. See also table 2). 

COLLOCATIONS WITH STRONG-MOTION 

INSTRUMENTS 

Twenty-three GEOS recorders were located where the 

main shock was also recorded. These sites are designated 

as CS in table 1. All stations but one were located with 

strong-motion accelerometers operated by the state of California (Shakal and others, 1989) or by the USGS (Maley 

and others, 1989). For further information on these strongmotion collocations, see Mueller and Glassmoyer (1990). 

AMF was deployed across the street from the Hyatt Hotel 

in Burlingame, where accelerometers operated by the building owners recorded the main shock (Celebi and McGarr, 

199la, b). The high-rise building in which AMF was located has been razed, apparently because of earthquake 

damage (Celebi, oral commun., 1993). 

DR-200 DATA 

During the aftershock deployment, 11 DR-200 recorders were deployed at 101 sites. A total of 2,926 threecomponent velocity records were recorded. The main focus 

of these studies was on local site effects on ground motion. In special studies, data were recorded to investigate 

topographic effects and sites where liquefaction occurred. 

Five instruments were collocated with, or very near, sites 

where the main shock was recorded. Two of these latter 

sites were also occupied by GEOS recorders. Information 

analogous to table 1, describing the DR-200 deployments, 

is available in the detailed reports given by Carver and 

others (1990) and King and others (1990). Figures 9 

Figure 3.-GEOS stations in Santa Cruz and southern Santa Clara Coun- shock was recorded at CAP, CRGIKOI (collocated GEOSs), DMD, GAOI 

ties (south-central region in figure 1). KKO+ marks a three-component, GA2lGAV (GEOS's in and around building where main shock was resub-kilometer-scale array that was part of a topographic experiment that corded), and SAR. Map from Mueller and Glassmoyer (1990). 

also included eight DR-200 recorders (see figures 8 and 14). The main 
Dl26 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

through 14, taken from Carver and others (1990), show 

the locations of the sites they studied. 

SANTA CRUZ 

Thirty-seven sites were occupied in the Santa Cruz area 

(fig. 10; Carver and others, 1990). The main objective 

was to characterize the response of different geologic units 

in the city. Part of the deployment was a systematic study 

of site response of the flood plain of the San Lorenzo 

River (fig. 11). This study was presented by Cranswick 

and others (1990). Brief comments on some of the sites 

occupied during this deployment are listed in table 3. In 

addition to aftershock recordings, King and others (1990) 

analyzed small-scale refraction studies of 10 sites in the 

Santa Cruz area. Their results are relevant to velocities at 

depths of up to approximately 20 to75 meters. See also 

Williams and others (1994). 

MOSS LANDING POWER PLANT 

Three DR-200's were deployed at and in the power 

plant (fig. 9) to gather information regarding this significant site. One instrument was in the free field and two 

were deployed in the control room in the upper floors of 

the power plant. These stations are discussed in Carver 

and others (1990). The Electric Power Research Institute 

also deployed instruments at the same time as the DR200's (Schneider, 1990). 

LOS GATOS 

Twenty-eight sites were occupied to study the effects 

of local site geology on ground motion in the city of Los 

Gatos (fig. 12). Stations LG3 and LG6 were located in the 

area of heaviest damage in the old business district along 

Santa Cruz avenue. Stations L12 and LEX, which was 

collocated with the CSMIP strong-motion instrument on 

Figure 4.-GEOS stations in San Mateo, southern Alameda, and north- BMT, COY, PKT, RAV, ST1+, and USG. 30-meter velocity logs are 

ern Santa Clara Counties (north-central region in figure 1). Main shock available from boreholes near AP7, BMT, MAL, and ST1+. A 180- 

was recorded at AMF, AP2, AP7, ASH, BBP, HALmA2, MAL, POR, meter velocity log is available for RAV. See table 2 for references to 

and SAR. AG1+, DUV+, and ST1+ mark sub-kilometer-scale arrays (see these data. Map from Mueller and Glassmoyer (1990). 

figure 8). Nuclear explosions have been recorded at or near stations AP2, 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA Dl27 

~ the left abutment of Lexington Dam, served as hard-rock 

reference sites. 

LIQUEFACTION SITES 

Seven instruments were installed at sites in Monterey 

and south Santa Cruz Counties where liquefaction occurred 

during the main shock. They are sites AIR, ART, BEL, 

MIL, MOL, MOS, and ORD and are shown in figure 9. 

TOPOGRAPHIC EFFECTS STUDIES 

Two deployments were made in the Santa Cruz mountains to study variations of ground motion with topography over distances of less than a kilometer. These studies 

were motivated by dramatic damage to structures along 

ridge tops. The site locations and topography are indi- I 

I cated in figures 9, 13 and 14. 

The study of the ridge topped by Rebecca Drive, east 

of Boulder Creek (fig. 9), included 15 stations installed 

along the top and face of the ridge and in the canyons on 

either side of the ridge (fig. 13). Stations MID and REB 

were located near minor damage to residences. TOP was 

placed in an area of significant structural damage, which 

seemed largely due to its position at the head-scarp of a 

landslide (King, written commun., 1990). In addition to 

variations in topography, there were significant variations 

in site geology among the stations in this study. 

The second study was of a ridge topped by Robinwood 

Lane above Hester Creek off of Old Santa Cruz highway 

(fig. 9). The ridge appeared to be geologically homogeneous. Eight DR-200's and three GEOS recorders were 

deployed along the ridge top and descending approximately 

45 meters down slope from the site of severely damaged 

houses near stations LP5 and LP6 (fig. 14). The GEOS 

sites are denoted by the letters KK in the station name, 

and the DR-200 sites are denoted by LP. 

Figure 5.-GEOS stations in San Mateo, San Francisco, and Alameda Counties (northernmost region in figure 1). Area in box is detailed in figure 

6. Main shock was recorded at DIA, FPO, MAL, and SF1. A 150-meter velocity log is available for station SF1 (see table 2). No data were obtained 

from CRA. Map from Mueller and Glassmoyer (1990). 
Dl28 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

COLLOCATIONS WITH STRONG-MOTION 

INSTRUMENTS 

Instruments were deployed at five sites where, or close 

to where, the main shock was recorded. Three of the sites 

are shown in figure 9: HAL, SAL and SAR. Instruments 

were also collocated with a strong-motion accelerometer 

on the left abutment of Lexington Dam (LEX in fig. 12) 

and with the strong-motion accelerometer at UC Santa 

Cruz (Station LOE in fig. 10). With the exception of SAR 

(see Maley and others, 1989), the strong-motion instruments were operated by CSMIP (Shakal and others, 1989). 

deployed with 2-Hz vertical geophones. See Healy and 

others (1983) for a description of the instruments. The 

recorders do not operate in a triggered mode, so they 

were programmed to turn on for six time intervals lasting 

12 to 13 minutes each. The time intervals began at 0600, 

0630, 1000, 1030, 2300, and 2300 universal time on October 19, 1989. These data have been digitized and analyzed by Eberhart-Phillips and others (1990). They are 

also discussed in this professonal paper (Eberhart-Phillips 

and others, this chapter). No report has been written specifically describing these data. 

ACKNOWLEDGMENTS 

USGS SEISMIC CASSETTE RECORDERS 

On the night following the main shock, 60 cassette recorders were deployed in the epicentral area to study 

crustal structure (fig. 15). These analog recorders were 

Paul Spudich suggested that this report would be of use 

to future workers. Jack Boatwright provided particularly 

thorough comments on the GEOS data set. Dave Carver 

and Charles Mueller were helpful by providing copies of 

LEA Ã AUD 

CAL Ã 

CCS ÃˆSN Ã RIN,R12 

GROÃ 

k 

CH@~^~ 

Ã KIS 

9 CWH 

Figure 6.-GEOS stations in San Francisco and Alameda Counties. Marina district is outlined and detailed in figure 7. Main shock was 

recorded at CAL, EMT, and RINtRI2. Nuclear explosions have been recorded at or near stations AUD, BLX, EMT, RIN, R12, and SND 

(see table 2). No data were obtained at CCS or YBI. Map from Mueller and Glassmoyer (1990). 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

SAN FRANCISCO BAY 

- - - 

MARINA GREEN 

0 

lllll 

meters LOMBARD ST 

Figure 7.-GEOS stations in the Marina District of San Francisco. Borehole geologic, velocity, and downhole seismic data are available at 

(see table 2). Map from Mueller and Glassmoyer (1990). 

DAW 

DUV 0 

DAS KK1 0 

KKO 

sws 

Figure 8.-Details of small GEOS arrays indicating 

dimensions and relative station postions. North is up. 

See figures 3 and 4 for array locations. Figure 14 shows 

topographic detail for, and DR-200 stations recording 

simultaneously with, the array KKO, KK1, and KK2. 

Map from Mueller and Glassmoyer (1990). 
Dl30 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

their figures as well as review comments on an earlier 

draft. An unpublished report of Ken King's was helpful. 

Andy Michael provided timing and location information 

for the cassette recorder data. I would also like to thank 

Mehmet Celebi, Chris Dietel, Art Frankel, Jim Gibbs, Gary 

Glassmoyer, Tom Hanks, Tom Holzer, Art McGarr, and 

Linda Seekins for their helpful input. 

REFERENCES 

Boatwright, J., Seekins, L.C., and Mueller, C., 1990, Ground motion 

amplification in the Marina District, in Effects of the Loma Prieta 

earthquake on the Marina District, San Francisco, California: U. 

S. Geological Survey Open-File Report 90-253, P. F1-F24. 

Boatwright, J., Fletcher, J.B., and Fumal, T., 199la, A general inversion scheme for source, site and propagation characteristics using 

multiply recorded sets of moderate-sized earthquakes: Bulletin of 

the Seismological Society of America, v. 81, no. 5, p. 1754-1782. 

Boatwright, J., Seekins, L.C., Fumal, T.E., Liu, H-P., and Mueller, C., 

199 lb, Ground motion amplification in the Marina District: Bulletin of the Seismological Society of America, v. 81, no. 5, p. 1980- 

1997. 

Boatwright, J., Seekins, L.C., Fumal, T.E., Liu, H-P., and Mueller, C., 

1992, Ground motion amplification, in The Loma Prieta, California, earthquake of October 17, 1989- Marina District: U.S. Geological Survey Professional Paper 1551-F, p. 35-50. 

Borcherdt, R.D., 1970, Effects of local geology on ground motion near 

San Francisco Bay: Bulletin of the Seismological Society of 

America, v. 60, no. 1, p. 29-60. 

Borcherdt, R.D., and Gibbs, J.F., 1976, Effects of local geological conditions in the San Francisco Bay region on ground motions and 

the intensities of the 1906 earthquake: Bulletin of the Seismological Society of America, v. 66, no. 1, p. 467-500. 

Borcherdt, R.D., Gibbs, J.F., and Fumal, T.E., 1978, Progress on 

ground motion predictions for the San Francisco Bay region, California, in Proceedings of the Second International Conference on 

Microzonation for Safer Construction-Research and Application, 

San Francisco, California: National Science Foundation, 

UNESCO, American Society Civil Engineers, EERI, SSA, Universities Council for Earthquake Engineering Research, p. 241. 

Borcherdt, R.D., Gibbs, J.F., and Lajoie, K.R., 1975, Prediction of 

maximum earthquake intensity in the San Francisco Bay region, 

California, for large earthquakes on the San Andreas and Hayward 

faults: US Geological Survey Miscellaneous Field Studies Map 

MF-709, scale 1: 125,000, 11 p. 

Borcherdt, R., Fletcher, J., Jensen, E., Maxwell, G., VanShaak, J., 

Warrick, R., Cranswick, E., Johnston, M., and McClearn, R., 

1985, A general earthquake observation system (GEOS): Bulletin 

of the Seismological Society of America, v. 75, no. 6, p. 1783- 

1825. 

Carver, D.L., Cunningham, D.R., and King, K.W., 1986, Calibration 

and acceptance testing of the DR-200 digital seismograph: U.S. 

Geological Survey Open-File Report 86-340, 28 p. 

Carver, D.L., King, K.W. , Cranswick, E., Worley, D.M., Spudich, P., 

and Mueller, C., 1990, Digital recordings of aftershocks of the 

October 17, 1989, Loma Prieta, California, earthquake: Santa 

Cruz, Los Gatos, and surrounding areas: U.S. Geological Survey 

Open-File Report 90-683, 203 p. 

Celebi, M., and McGarr, A., 1991a, Discussion of site-response at Foster City and San Francisco Airport, in Proceedings of the Fourth 

International Conference on Seismic Zonation, August 25-29, 

199 1, Stanford, California: Earthquake Engineering Research 

Institute, v. 111, p. 367-373. 

Celebi, M., and McGarr, A., 1991b, Site-response at Foster City and 

San Francisco Airport as Loma Prieta studies, in Soil Dynamics 

and Earthquake Engineering V, Computational Mechanics Publications: Southampton and London, England, Elsevier Applied Science, p. 35-46. 

Clark, J.C., 1981, Geologic map and sections of the Felton-Santa Cruz 

area, Santa Cruz County, California: U.S. Geological Survey Professional Paper 1168, plate 2. 

Cranswick, E., King, K., Carver, D., Worley, D.,Williams, R., 

Spudich, P., and Banfill, R., 1990, Site response across downtown 

Santa Cruz, California: Geophysical Research Letters, v. 17, no. 

10, p. 1793-1796. 

Eberhart-Phillips, D.M., Michael, A. ., Fuis, G., and Luzitano, R., 

1990, Three-dimensional crustal velocity structure in the region of 

the Loma Prieta, California, earthquake sequence from inversion 

of local earthquake and shot arrival times: Seismological Research 

Letters, v. 61, no. 1, p. 48. 

Fletcher, J.B., and Boatwright, J., 1991, Source parameters of Loma 

Prieta aftershocks and wave propagation characteristics along the 

San Francisco peninsula from a joint inversion of digital seismograms: Bulletin of the Seismological Society of America, v. 81, 

no. 5, p. 1783 -1812. 

Gibbs, J.F., and Borcherdt, R.D., 1974, Effects of local geology on 

ground motion in the San Francisco Bay region, California-a 

continued study: U.S. Geological Survey Open-File Report 74- 

222, 146 p. 

Gibbs, J.F., Fumal, T.E., and Borcherdt, R.D., 1975, In-situ measurements of seismic velocities at twelve locations in the San Francisco Bay region: U.S. Geological Survey Open-File Report 75- 

564, 87 p. 

1976 a, In-situ measurements of seismic velocities in the San 

Francisco Bay region; part 11: U.S. Geological Survey Open-File 

Report 76-73 1, 145 p. 

Gibbs, J.F., Fumal, T.E., Borcherdt, RD., and Roth, E.F., 1976 b, Insitu measurements of seismic velocities in the San Francisco Bay 

region; part 111: U.S. Geological Survey Open-File Report 77-850, 

142 p. 

Gibbs, J.F., Fumal, T.E., Boore, D.M., and Joyner, W.B., 1992, Seismic velocities and geologic logs from borehole measurements at 

seven strong-motion stations that recorded the Loma Prieta earthquake: U.S. Geological Survey Open-File Report 92-287, 139 p. 

Hartzell, S., 1978, Earthquake aftershocks as Green's functions: Geophysical Research Letters, v. 5, no. 1 p. 1 -4. 

Healy, J.H., Mooney, W.D., Blank, H.R., Gettings, M.E., Kohler, 

W.M., Lamson, and R.J., Leone, L.E., 1983, Saudi Arabian seismic deep-refraction profile: final project report: U.S. Geological 

Survey Open-File Report 83-390, 360 p., 9 plates. 

Helley, E.J. and LaJoie, K.R., 1979, Flatlands deposits of the San 

Francisco Bay Region, California: their geology and engineering 

properties and their importance to compreshensive planning: U.S. 

Geological Survey Professional Paper 943, 88 pp. 

Joyner, W.B., Warrick, R.E., and Oliver, A.A. Ill, 1976, Analysis of 

seismograms from a downhole array in sediments near San Francisco Bay: Bulletin of the Seismological Society of America, v. 

66, no. 3, p. 937-958. 

Kayen, R.E., Liu, H-P., Fumal, T.E.,Westerlund, R.E., Warrick, R.E., 

Gibbs, J.F., and Lee, H.J., 1990, Engineering and seismic properties of the soil column at Winfield Scott School, San Francisco, in 

Effects of the Loma Prieta earthquake on the Marina District, San 

Francisco, California: US Geological Survey Open-File Report 90- 

253, p. Gl-G18. 

King, K., Carver, D., Williams, R., Worley, D., Cranswick, E., and 

Meremonte, M., 1990, Santa Cruz seismic investigations following 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

Table 1 .-GEOS data sites described in Mueller and Glassmoyer (1 990) and main experimental objectives 

Abbreviations: 

BB 

cs 

FC 

rr 

Mb 

Ms 

NS 

Old 

PE 

PP 

SM 

S A 

sv 

SF 

SF0 

SP 

u 

Oakland Bay Bridge study 

Colocated with a USGS, CSMIP or private strong-motion accelerograph that recorded the mainshock 

Foster City study of bay-margin response 

Instrumentation test 

Surface instrument near subsequent hole drilled in Marina District of San Francisco 

Site study in the Marina District of San Francisco 

Colocated with a strong-motion accelerograph without a mainshock record 

In previously published site response study. See table 2. 

Line from southern peninsula to the southern East Bay 

Peninsula wave-propagation study 

Sub-km-scale joint deployment with DR-200 recorders for Santa Cruz mountain topographic study 

Sub-lun-scale four-station array on alluvium in the central Santa Clara Valley 

Line crossing northern Santa Clara Valley 

San Francisco site-response study 

Few-km-scale array at San Francisco Airport 

Sub-krn-scale three-instrument array on the alluvium in the southern peninsula 

Geologic unit of specific interest. Alluvial unit descriptions from Helley et a1 (1979) 

BB CS FC IT Mb Ms NS Old PE PP SA SM SV SF SF0 SP U 

ACQ ........................... ** .......................... gravel quarry 

AG1 .................................. **.. ... ** 

AG2 .... ** ....... (same location as ASH) ....... ** ..... * * 

AG3 .................................. ** ..... * * 

AG4 .................................. **.. ... * * 

AMF .... ** ................................................. bay mud, damaged high-rise 

AP2 ....** . ** ............... ** ............................. baymud 

AP7 .... ** . ** ............... **. ... * * 

ASH .... ** ....... (same location as AG2) .............. * * 

AUD ........................ ** .................. ** ......... near damaged freeway 

BB4 ** 

BB5 ** 

BBP .... ** ................................................. bay mud 

BBT ** 

BCC .............................. ** 

BEA ................. **.. .................................. hydraulicfill 

the October 17, 1989 Loma Prieta earthquake: U.S. Geological 

Survey Open-File Report 90-307, 59 p. 

Liu, H-P., Warrick, R.E., Westerlund, R.E., Sembera, E. D., and 

Wennerberg, L., 1992a, Observation of local site effects at a 

downhole-and-surface station, in The Loma Prieta, California, 

earthquake of October 17, 1989-Marina District: U.S. Geological 

Survey Professional Paper 155 1 -F, p. 5 1-74. 

Liu, H-P., Warrick, R.E., Westerlund, R.E., Sembera, E.D., and 

Wennerberg, L., 1992b, Observation of local site effects at a 

downhole-and-surface station in the Marina District of San Francisco: Bulletin of the Seismological Society of America, v. 82, no. 

4, p 1563 -1591. 

Maley, R., Acosta, A., Ellis, F., Etheredge, E., Foote, L., Johnson, D., 

Porcella, R., Salsman, M., and Switzer, J., 1989, U.S. Geological 

Survey strong-motion records from the northern California (Loma 

Prieta) earthquake of October 17, 1989: U.S. Geological Survey 

Open-File Report 89-568, 85 p. 

Margheriti, L., Wennerberg, L., and Boatwright, J., 1994, A comparison of coda and S-wave spectral ratios: estimation of site response 

in the south San Francisco Bay, Bulletin of the Seismological Society of America, v. 84, no. 6, p. 1815-1830. 

McGarr, A., Qslebi, M., Sembera, E., Noce, T., and Mueller, C., 1991, 

Ground motion at the San Francisco International Airport from the 

Loma Prieta earthquake sequence, 1989: Bulletin of the Seismological Society of America, v. 81, no. 5, p. 1923-1944. 

Mueller, C. and Glassmoyer, G.,1990, Digital recordings of aftershocks of the 17 October 1989 Loma Prieta, California, earthquake: U.S. Geological Survey Open-File Report 90-503, 147 p. 

Schneider, J., 1990, Aftershock recordings of the Loma Prieta earthquake: project RP 3 18 1 - 1, EPRI Technical Report. 

Seekins, L.C., and Boatwright, J., 1994, Ground motion amplification, 

geology, and damage from the 1989 Loma Prieta earthquake in 

the City of San Francisco: Bulletin of the Seismological Society 

of America, v. 84, no. 1, p. 16-30. 

Shakal, A., Huang, M., Reichle, M., Ventura, C., Cao, T., Sherburne, 

R., Savage, M., Darragh, R., and Peterson, C., 1989, CSMIP 

strong-motion records from the Santa Cruz mountains (Loma 

Prieta), California earthquake of 17 October 1989: California Department of Conservation, Division of Mines and Geology, Office 

of Strong Motion Studies Report OSMS 89-06, 195 p. 

Warrick, R.E., 1974, Seismic investigation of a San Francisco Bay 

mud site: Bulletin of the Seismological Society of America, v. 64, 

no. 2, p. 375-385. 

Williams, R.A., Cranswick, E., King, K.W. , Carver, D.L., and 

Worley, D.M., 1994, Site-response models from high-resolution 

seismic reflection and refraction data recorded in Santa Cruz, 

California, in Borcherdt, R.D., ed., Strong ground motion: U.S. 

Geological Survey Professional Paper 155 1 -A, p. 2 17-242. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 1 .-Continued 

BB CS FC IT Mb Ms NS Old PE PP SA SM SV SF SF0 SP U 

BLX ........................ ** .................. ** ......... near damage 

BMT ............. B VF ............. BV2 ............. CAL .... ** ....... CAP .... * * 

CGC ............. CHB ............. CNT ............. COY ............. CRG .... ** .... ** . CWH ............. DAS ............. DAW ............. DEM ............. DIA .... ** ....... 

........... **. ... * * 

. ** ............... (BVF inside shed with 

........ ** ............... bad resonance. BV2 outside) * * .... 

* * ................. 

............................... ** ......... bay mud 

.............. ** .......................... Holocene coarse-grain alluvium 

........... ** . * * .......................... hard rock 

(Guralp accelerometer collocated with KOI) 

............................... * * ......... near liquefaction 

.............. ** ...................... ** .. Holocene fine-med. alluvium 

.............. * * ...................... * * .. Holocene fine-med. alluvium 

. ** .................................... on (near?) hydraulic fill * * ............................... DMD ....... * * 

DUV ........................... ** ...................... ** .. Holocene fine-med. alluvium 

EDS * * ......................................... EMT**..** * * .................. ESC ........................... ** .......................... Holocene med. grain alluvium 

FER ............................................ ** ......... building on pilings 

FLB ........................... * * .......................... near-fault, Pleistocene alluvium 

FOX ....... ** .............................................. bay margin, near damage? 

FPO ....** ...................................... * * ......... near damage 

GAO ... ** .... (GAV and GA2 were 

GAV .... * * ... sequentially deployed 

GA2 .... ** ... in and near one structure) 

GRO ............................................ ** ......... dune sand near damage 

HAL .... ** .... (HA2 same location as HAL, 

HA2 ... ** ... but with different sensors) 

IVE ........................... ** .......................... Pleistocene alluvium 

JSQ ............................................ ** ......... bay mud near undamaged 

unreinforced masonry 

KIS ............................................ ** ......... dune sand 

KKO * * ...................................... KK1 ...................................... * * 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

Table 1 .-Continued 

BB CS FC IT Mb Ms NS Old PE PP SA SM SV SF SF0 SP U 

KK2 ...................................... ik* 

KO1 .... ** . (Collocated with Guralp, station CRG) .. ** 

LEA ................. ....**. .................... * * 

LMS ................. ** ................................... near (but off?) hydraulic fill 

MAL .... ** . ** ............... * * 

MAS ................. ** ............... (MAS and MA2 were sequentially 

MA2 ................ ** ............. near each other on hard rock) 

MDC ........................... ** .......................... Holcene fine-med. alluvium 

MON ....... * * 

NPT ................. ** .................................... close to greatest damage; fire 

OAB ** 

PKT .... ** .................. ** ............... * * 

POR ... * * 

PUC ................. ** .................................... off fill, building renovated 1906 

RAV .......... ** ............ ** ............................. bay mud 

RIN ....** .................. ** . (RIN relocated to ...... * * 

R12. .... ** ................. ** . R12. Same building) ... * * 

RUS ............................................ ** ......... near damage 

SAG .... * * 

SAR ....**. ....................... * * 

SBM ........................ **. * * ... SF1 .... ** .................. ** ..................... * * 

SF2 ............................................... **.. .... hardrock 

SF3 .............................. ** ............... ** ...... near damage 

SF4 .............................. **.. ............. * * 

SND ............................................ * * ......... near damage 

SRL ........................... ** . ** 

ST 1 .......... ** ............ * * 

ST2 .......... ** ............ * * 

ST3 .......... ** ............ * * 

ST4 .......... ** ............ * * 

STQ ........................... ** .......................... damage, Pleist. alluvium 

SUN ......................................... * * 

sws ............. ** . * ** 

TOM .... ** 

TUL ......................................... * * 

USG .......... ** ............ * * 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 2.-GEOS aftershock recording locations at or near previous USGS studies 

[The left column indicates analyses of seismic data from sites collocated with or nearby the listed GEOS aftershock sites. Letter-number 

combinations denote particular site and nuclear explosion combinations as discussed in Borcherdt (1970), Gibbs and Borcherdt (1974), 

Borcherdt and others (1975). and Borcherdt and Gibbs (1976). Two more site-specific seismic studies, which analyze borehole data, are 

explicitly listed. The right column lists velocity and geologic log data sources for the GEOS sites.] 

Nuclear explosion, and Velocity and geologic log data 

borehole seismic studies 

AP2 

AP7 

AUD 

BLX 

BMT 

COY 

EMT 

MAL 

I5c (less than 1 km away) 

"Pulgas Tunnel" in Gibbs and others (1976) 

05, P7 

T5 (less than 1 km away) 

S3, S4, S5, S16, J17, J18, S19 "Black Mountain" in Gibbs and others (1975) 

K1, K12, S13, S14, S15 

H 18 (less than 1 km away) 

"Audobon School" (- 1 km away) in 

Gibbs and others (1977) 

PKT PI8 

RAV 11 1, Joyner and others (1976) Warrick (1974) 

RIN J5, J7 

R12 J5, J7 

SF 1 "SF Airport" in Gibbs and others (1992) 

SND R7 (less than 1 km away) 

ST1 L3 "Page Mill" in Gibbs and others (1977) 

ST2 L3 â€ ̃'Pa Mill" in Gibbs and others (1977) 

ST3 L3 "Page Mill" in Gibbs and others (1977) 

ST4 L3 "Page Mill" in Gibbs and others ( 1976b) 

S W S Liu and others (1992a, b) Kayen and others (1990) 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

Central California Study Region 

Figure 9.-Locations of DR-200 recorders in Santa Cruz, Monterey and southern Santa Clara Counties. Boxes outline areas 

detailed in figures 10 to 14. Liquefaction occurred during the main shock at stations AIR, ART, BEL, MIL, MOL, MOS, and 

ORD. Main shock was recorded at HAL, SAL, and SAR. Map from Carver and others (1990). 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Santa Cruz Study Area 1220 I 

0 

- 

1 Kilometer 

Tertiary Sedimentary 

..... Marine Terrace . . . . . Deposits \ Crystal line Rocks 

Figure 10.-Locations of DR-200 recorders (triangles) in the city of Santa Cruz. Surficial geology adapted from Clark (1981). 

Main shock was recorded at LOE. Map from Carver and others (1990). 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

Santa Cruz Flood Plain Array 

1 ALLUVIUM TERTIARY SEDIMENTARY 

ROCKS 

1 KM ' 2 1 KILOMETERS 0 

Figure 11.-Locations of DR-200 recorders (triangles) for the study of seismic response of the flood 

plain of the San Lorenzo River (Cranswick and others, 1990). These sites were occupied simultaneously, Bottom of figure shows idealized geologic cross section, a portion of which is offset and 

drawn to a different scale than the rest of the figure. Map from Carver and others (1990). 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 3.-Noteworthy selected sites in Santa Cruz study 

[See Carver and others (1990) and King and others (1990).] 

ALM Marine terrace, minor chimney damage 

BAR Hard rock 

BLA Heavy damage 

CE2 Heavy damage 

CED Heavy damage 

KAL Hard rock 

LOE Mainshock recorded at site 

MA1 Marine terrace 

MON Heavy damage 

PAC Heavy damage 

WAL Heavy damage 

122 1 Los Gatos Study Area 

Figure 12.-Locations of DR-200 recorders (triangles) in the city of Los Gatos (roads shown). The shaded areas 

denote lakes. Highway 17 is indicated. Highway 9 intersects highway 17 at cloverleaf. Main shock was recorded at 

LEX. Map from Carver and others (1990). 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

Rebecca Ridge Study Area 

Contour Interval 40 feet 

0 0.5 Mile 

0 0.5 Kilometer 

Figure 13.-Locations of DR-200 recorders (triangles) used in the Rebecca Ridge topographic study. See figure 9 for array location. Map from Carver 

and others (1990). 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Robinwood Lane Study kea 

Contour Interval 40 feet 

0 0.5 Mile 

0.5 Kilometer 

Figure 14.-Locations of DR-200 and GEOS recorders used in the Robinwood Lane topographic study. GEOS 

instruments are denoted by KK. See figure 9 for array location. Map from Carver and others (1990). 
U.S. GEOLOGICAL SURVEY AFTERSHOCK GROUND-MOTION DATA 

Lorna Prieta 

Cassette recorder deployment 

Figure 15.-Locations of USGS seismic cassette recorders deployed the night of October 19, 1989 (triangles). Recordings were made continuously 

for 12- to 13-minute intervals starting at the following times: 0600, 0630, 1000, 1030, 2300, and 2330 universal time. For comparison, note that the 

main shock source zone is shown in figure 1 along with the same geologic features as in this figure. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKE OCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

RESPONSE OF US. GEOLOGICAL SURVEY CREEPMETERS TO THE 

LOMA PRIETA EARTHQUAKE 

CONTENTS 

By K.S. Breckenridge and R.W. Simpson, 

U.S. Geological Survey 

Abstract ............................................................... 

Introduction ............................................................... 

Data from eight creepmeters closest to the earthquake ---------- 

Preseismic slip deficit: precursory retardation or drought? ---- Signal versus noise in the creep data ................................ 

Coseismic steps and static stress changes .......................... Creep rate changes and static stress changes ...................... Coefficient of apparent friction ....................................... 

Creep rate versus stress law ........................................... 

Models exploring depth-origin of Loma Prieta afterslip --------- 

Retardation and anomalous behavior before the earthquake? --- Conclusions ............................................................... 

Acknowledgments ....................................................... 

Appendix A: Site notes ................................................. 

References cited ......................................................... 

Page 

Dl43 

143 

147 

147 

152 

152 

155 

159 

160 

164 

166 

171 

175 

176 

177 

ABSTRACT 

At the time of the Loma Prieta earthquake, 18 U.S. 

Geological Survey creepmeters in central California recorded coseismic steps ranging in size from -0.35 to +6.8 

mm. Five of the closest instruments on the San Andreas 

fault and three on the Calaveras fault also displayed longterm rate changes in the months after the earthquake. The 

coseismic steps seem to bear little relation in magnitude 

or sense to static stress changes calculated using dislocation models of the earthquake rupture, but 1-year average 

creep-rate changes (faster on the San Andreas and slower 

on the Calaveras) do correlate well with static stress 

changes. This correlation favors low values of apparent 

coefficient of friction. Observed advances and deficits in 

cumulative slip at the closest sites caused by the positive 

and negative rate changes are in fair agreement with deformation predicted by a three-dimensional dislocation 

model that requires anomalous slip to extend to depths in 

excess of 10 km. Rate changes observed at several 

creepmeters in the years before Loma Prieta may be precursory but are difficult to interpret unambiguously because they fall in a period of extended drought. A 2-year 

period of retarded creep at the Cienega Winery site before 

the Lorna Prieta earthquake contains an interesting interval with normal right-lateral creep events superposed on 

unusual left-lateral background drift. Assuming that this 

behavior was tectonic in origin and not drought- orinstrument-related, it can be explained by stresses imposed from 

two different sources. One source could be slip on the 

San Andreas fault below the creepmeter, causing the rightlateral creep events. The second could be slip on the 

subparallel Calaveras fault or other local structures which 

were caught up in regional tectonic adjustments following 

the 26 January 1986 Tres Pinos earthquake, causing the 

left-lateral movement and retarding the cumulative progress 

of the instrument. In this scenario, the retardation at 

Cienega Winery could have been a precursor to Loma 

Prieta in the sense that it too was a manifestation of regional tectonic changes that would eventually trigger the 

Loma Prieta earthquake. 

INTRODUCTION 

At the time of the Loma Prieta earthquake, 27 U.S. 

Geological Survey (USGS) creepmeters were operating 

within 205 km of the epicenter (fig. 1A). Eighteen of these 

instruments recorded coseismic steps ranging in size from 

-0.35 to +6.8 mm, and eight instruments located closest 

to the epicenter recorded significant changes in slip rate 

in the months after the earthquake (table 1). For several 

creepmeters, the Loma Prieta signal is the largest anomaly 

recorded in up to 25 years of operation. Possible precursory rate changes were observed on several instruments 

(Breckenridge and Burford, 1990). 

The USGS creepmeters are of several different designs, 

and data are collected at various intervals ranging from 

minutes to months, depending upon the configuration of a 

particular instrument (table 1). Descriptions of the instruments and sites have been published in Schulz and others 

(1982, 1983) and more recently in Schulz (1989). Eighteen sites are equipped with satellite telemetry that samples 
Dl44 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

the creepmeter every 10 minutes (Silverman and others, discussed by Behr and others (1990) and by Behr and 

1989), while micrometer dial readings at the remaining 9 others (this chapter). Alignment arrays on the Calaveras, 

sites are taken quarterly. Hayward, and other Bay Area faults maintained by reAdditional creepmeters on the San Andreas fault in- searchers at the University of California at San Francisco 

stalled by researchers from the University of Colorado are are described by Galehouse (1992, and this chapter). 

Figure 1.-A, Locations of USGS creepmeters in central California op- creepmeters which experienced long-term rate changes (table 1). B, Loerating at the time of the Lorna Prieta earthquake. The epicenter of the cations of eight creepmeters lying within 60 km of the Loma Prieta 

earthquake is shown by a star. Filled triangles indicate creepmeters rupture. Epicenters of six earthquakes discussed in the text are also 

which recorded definite coseismic steps; larger triangles indicate shown as stars. 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl45 

Numerous instances of triggered slip on fault segments 

at considerable distances from an earthquake rupture are 

reported in the literature. For example, such slip occurred 

after the 1968 Borrego Mountain earthquake (Allen and 

others, 1972), the 1979 Imperial Valley earthquake (Fuis, 

1982; Sieh, 1982; Bilham and Williams, 1985), the 1981 

Westmorland earthquake (Sharp and others, 1986a), the 

1983 Coalinga earthquake (Mavko and others, 1985; 

Schulz and others, 1987), the 1984 Morgan Hill earthquake (Schulz, 1984), the 1986 Tres Pinos earthquake 

(Simpson and others, 1988), the 1986 North Palm Springs 

earthquake (Fagerson and others, 1986; Sharp and others, 

1986b, Wesson and others, 1986; Williams and others, 

1986, 1988), the 1987 Superstition Hills earthquake (Sharp, 

1989; Hudnut and Clark, 1989; McGill and others, 1989), 

and the 1989 Loma Prieta earthquake (Behr and others, 

1990; Galehouse, 1990; McClellan and Hay, 1990). 

If the triggered slip persists, it may be difficult to distinguish from afterslip, although traditionally "afterslip" 

has been used to describe postseismic slip occurring on 

the same fault that the earthquake occurred on and localized in the rupture region or in its immediate vicinity (for 

example, Smith and Wyss, 1968; Scholz and others, 1969; 

Burford, 1972; Burford and others, 1973; Cohn and others, 1982; Wesson, 1987; Bilham, 1989; Marone and others, 1991). 

Lorna Prieta 

(840424, M6.2) 

Coyote Lake (79 

Figure 1 .-Continued 
Dl46 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 1.-Summary table of all USGS creepmeters in operation at the time of the Lorna Prieta earthquake 

[Asterisk (*) in first column indicates a site on the Southwest Fracture near Parkfield. Asterisk (*) in the coseismic change column 

indicates initial RL movement which was offset by subsequent LL movement within hours to days of the earthquake. Question marks in 

the coseismic change column indicate lack of data for determining coseismic movement. Negative values indicate left-lateral coseisrnic 

steps; positive values, right-lateral.] 

XS J 

XHR 

cwc 

XFL 

XMR 

MRW 

BIT 

XMP 

xsc 

XMM 

XMD 

XVA 

XRS * 

XPK 

XTA 

WKR 

XHS * 

CRR 

XGH 

X46 

HWR 

HWE 

HWW 

HWP 

XSH 

HLC 

HLD 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

San Andreas 

Hay ward 

Hay ward 

Hay ward 

Hayward 

Calaveras 

Calaveras 

Calaveras 

Site Fault Origin Sample Rate Distance from Coseismic Post-LP 

(moly r) Loma Prieta (km) Change (mm) Rate Change(?) 

1 1/74 10 minute 39 5.2 * Y 

10 minute 

10 minute 

10 minute 

10 minute 

quarterly 

quarterly 

quarterly 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

10 minute 

quarterly 

quarterly 

quarterly 

quarterly 

10 minute 

quarterly 

4/70 quarterly 48 ? Y 

Some of these triggered responses, especially coseismic 

steps recorded on creepmeters, are most likely produced 

by the shaking of fault and instrument during passage of 

the seismic waves (King and others, 1977; Allen and others, 1972; Fuis, 1982; Williams and others, 1988; McGill 

and others, 1989). Such shaking probably triggers the release of a backlog of slip that had been held in the Earth's 

near-surface or in the creepmeter itself by friction. Other 

longer-term responses, including alterations in creep-rate 

and occasionally in creep direction (Mavko, 1982; Mavko 

and others, 1985; Simpson and others, 1988), may be 

caused by static changes in the stress field produced by 

the distant earthquake's fault offset. The reality of such 

static stress changes in the earth is not in question, because sensitive strainmeters have detected them at great 

distances from earthquakes (for example, Johnston and 

others, 1987; Shimada and others, 1987; Johnston and 

others, 1990). Microseismicity rate changes on central 

California faults also appear to correlate with calculated 

static stress changes after the Loma Prieta earthquake 

(Reasenberg and Simpson, 1992, and this chapter), suggesting that these faults can and do react to small stress 

perturbations at seismogenic depths. 

In the following sections, we present data for the eight 

USGS creepmeters that were situated nearest to the Loma 

Prieta epicenter. One of the largest uncertainties in the 

interpretation of creepmeter data is introduced by seasonal and rainfall-induced fluctuations! Therefore, rainfall 

records are also presented, and we explore the possible 

effects of weather on creepmeter behavior. Because tectonic accelerations or retardations in creep rates may occur before certain earthquakes (for example, Nason, 1973; 

Burford, 1988), establishing the tectonic as opposed to 

the rainfall-related significance of creep-rate changes is 

important, although difficult to accomplish (Goulty and 

Oilman, 1978, Langbein and others, 1993). 

We attempt to relate the observed coseismic steps and 

longer-term creep-rate changes that occurred at the time 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl47 

of the Loma Prieta earthquake to static stress changes 

calculated from simple dislocation models of the rupture 

in an elastic half-space. To put the Loma Prieta observations into perspective and to improve the statistics, we 

have compiled creepmeter data for four other moderateto-large Bay Area earthquakes that occurred before 1989, 

in addition to the 1990 Chittenden earthquake, which was 

a large aftershock to the Loma Prieta earthquake. The 

results described here suggest that although the sense 

(right-lateral, RL, or left-lateral, LL) of coseismic steps at 

creepmeter sites seems to bear little relation to the imposed static shear-stress changes, there is a statistically 

significant relation between the sense (RL or LL) of longterm average creep-rate changes and the sign of coseismic 

shear-stress changes. For the Loma Prieta earthquake, the 

magnitudes of these quantities are also well correlated. 

We also describe three-dimensional boundary element 

models used to estimate the total anomalous slip advance 

or deficit that can be expected at creepmeter sites as a 

result of the Loma Prieta creep-rate changes. The models 

suggest that, under some instruments, complex spatial distributions of stress can exist, and the effects of RL stress 

changes imposed at the surface might be superseded over 

time by larger LL stresses imposed at depth. These models suggest that for the five creepmeters on the San Andreas fault, cumulative advances ranging from 12 to 60 

mm can be expected as a result of the Loma Prieta earthquake, whereas at the three sites on the Calaveras fault, 

cumulative delays of -6 to -29 mm might be expected. 

Although post-Loma Prieta adjustments at the creepmeters 

are not yet complete, these estimates appear to agree with 

extrapolations of observed advances or deficits in longterm pre-earthquake trends. 

DATA FROM EIGHT CREEPMETERS 

CLOSEST TO THE EARTHQUAKE 

In this paper we concentrate our discussion on the responses of the eight closest USGS instruments on the San 

Andreas and Calaveras faults (fig. 1B). Data for all of the 

operating instruments is available to interested investigators. 

Long-term data from the five closest instruments on the 

San Andreas fault are shown in figure 2A. Data from a 3- 

year time window centered about the Loma Prieta earthquake date are displayed in figure 2B. These plots were 

made using daily data (1 point per day) selected (for the 

most part) from 10-minute interval data by an automatic 

algorithm. These daily data have been adjusted to agree 

with quarterly micrometer readings where data gaps exist 

or calibrations are in question. Such adjustments are usually unnecessary or small, and none of the significant variations in creep rate presented here are likely to result from 

instrument calibration problems. 

The most obvious features in the long-term data in figure 2A are the post-Loma Prieta rate increases at XSJ, 

XHR, and CWC. Post-seismic changes at XFL and XMR 

are smaller compared with the other three sites and with 

other variations recorded at those sites over the years. At 

the bottoms of figure 2A and 2B, plots of daily rainfall at 

Paicines, California, located near the fault between 

creepmeters CWC and XFL, show that the Loma Prieta 

earthquake occurred during a drought lasting from 1986 

to about 1991 (U.S. Dept. of Commerce). 

Figure 3A shows cumulative data for the three 

creepmeters along the Calaveras fault near Hollister. Two 

traces are plotted for the Shore Road creepmeter; XSH for 

daily data selected from 10-minute telemetered data and 

XSHM for approximately quarterly micrometer measurements to facilitate comparison with dial readings from 

Central Avenue (HLC) and D Street (HLD) in Hollister. 

The most obvious feature is the slip-rate decrease since 

the Loma Prieta earthquake at all three sites. There are 

indications that this decrease began before the earthquake 

and may be related to the M=5.6 Tres Pinos earthquake 

on 26 January 1986. These declining slip-rates occur during a period of drought, as shown by the plot of daily 

rainfall from Paicines, which casts some uncertainty on 

the tectonic significance of these changes. Galehouse (this 

chapter) has reported slower creep rates at a number of 

sites on the Hayward and southern Calaveras faults after 

the Loma Prieta earthquake. 

PRESEISMIC SLIP DEFICITS: 

PRECURSORY RETARDATION OR 

DROUGHT? 

Small, pre-Loma Prieta rate changes occurred at XHR, 

CWC, and XMR in the 2-3 years before the earthquake 

(figs. 2A, B). It is of some interest to know whether these 

changes are of tectonic origin, representing precursors to 

the Loma Prieta earthquake. Burford (1988), for example, 

discussed the possible significance of retardations in creep 

rate as precursory signals for magnitude 4-6 earthquakes. 

Burford also speculated that a larger, more distant shock 

might be associated with slip deficits across a broader 

region. The rate reductions seen at several creepmeters 

before the Loma Prieta earthquake are tantalizing in that 

respect. However, creepmeters are also known to respond 

to rainfall, and the prolonged drought that occurred from 

1986 to 1991 might also play a role in preseismic creep 

rates. 

Schulz and others (1983) examined the impact of rainfall at XSH, XSJ, and XMR during the 1970's, which 

includes the drought period from 1976 to 1978. Reviewing 8-11 years of creep data, they found no significant 

correlation between rainfall and either the total amount of 

annual creep, the number of annual episodic events, or 
x SJ 

XHR 

cwc 

PAIR 5 

- z 

^ 

1 I-' 

XSJ 

XHR 

cwc 

XFL 

XMR 

1. 

oI._lltn,JllJ.. "hLh&., I&;. 

PAIR 

0 . 

1987 1988 1989 1990 1991 1992 

Figure 2.-A, Cumulative creep data for five creepmeters on the San Andreas fault five creepmeters on the San Andreas fault from three years before to three years after 

from 1970 to 1993. Rainfall occurrences are also shown at the bottom of the plot. the Loma Prieta earthquake. Rainfall occurrences are also shown at the bottom of the 

Vertical lines A through F mark the times of six moderate to large central California plot. Vertical lines E and F mark the times of the Loma Prieta earthquake and the 

earthquakes discussed in the text and shown in figure 1B. B, Cumulative creep data for Chittenden aftershock. 
HLD 

Figure 3.-A, Cumulative creep data for three creepmeters on the Calaveras fault Calaveras fault from three years before to three years after the Loma Prieta earthfrom 1970 to 1993. XSHM is micrometer readings at site XSH, presented for easier quake. Trace XSHM represents quarterly micrometer readings at site XSH, presented 

comparison with the micrometer readings used for HLC and HLD. Rainfall occur- for easier comparison with the dial gauge readings from HLC and HLD. Rainfall 

rences are also shown at the bottom of the plot. Vertical lines A through F mark the occurrences are also shown at the bottom of the plot. Vertical lines E and F mark the 

times of six moderate to large central California earthquakes discussed in the text times of the Loma Prieta earthquake and the Chittenden earthquake. 

and shown in figure IB. 5, Cumulative creep data for three creepmeters on the 
D 150 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

the total creep accumulated from annual episodic events. 

The study identified slight RL or LL increases at XSJ and 

XMR associated with rainfall, and at XSH, oscillatory 

events of about 1-day duration accompanying rain. No 

obvious seasonal trends are present at XHR, CWC, or 

XFL. Figures 4A and 4B depict residual creep data with 

residual rainfall from Paicines, located just south of the 

CWC creepmeter. An annual average rate of 15.4 inches 

per year has been removed from the rain record to visually enhance seasonal changes. Drought periods, from 1976 

to 1978 and from 1986 to 1991, appear as downward 

trends, and are marked with horizontal bars above the 

traces. 

For San Andreas fault creepmeters in this study, there 

is no apparent seasonal trend associated with rainfall except at XMR, where strong LL movement occurred durFigure 4.-A, Detrended creep data for five creepmeters on the San the period 1970-1993. Trend was determined by calculating best-fitting 

Andreas fault for the period 1970-1993. Trend was determined by cal- least-squares line. XSHM is micrometer readings at site XSH, presented 

culating best-fitting least-squares line. Horizontal bars denote periods of for easier comparison with the micrometer readings used for HLC and 

creep retardation. Bottom plot shows detrended cumulative rainfall record. HLD. Horizontal bars denote creep retardation. Bottom plot shows 

B, Detrended creep data for three creepmeters on the Calaveras fault for detrended cumulative rainfall record. 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl51 

ing the 1991 and 1992 rainy seasons. The average creep 

rate at XMR for those years remains consistent with the 

long-term average of about 18 mmlyr despite these seasonal fluctuations. On the Calaveras fault, rainfall signals 

are more pronounced at XSH after 1986, when the instrument was rebuilt. The oscillatory character of the signal 

remains. Appendix A discusses the reconstruction at XSH 

in more detail. During the 1976 drought, XSH recorded a 

slip deficit before the Coyote Lake earthquake in 1979. 

When normal rainfall resumed in 1978, there was no 

change in the character of the retardation signal in response to increased precipitation. Schulz and others (1983) 

note that while lack of rain may delay onset of a creep 

event, stress build-up from slip at depth will eventually 

override conditions in the shallow soil to produce a creep 

event. We find no strong relationship between changes in 

creep rate and periods of drought or deluge for the 

creepmeters in this study. 

Turning to creep retardation, horizontal bars above creep 

data in figures 4A and 45 mark slip deficits identified by 

Burford (1988), updated through December 1992. Although 

the definition is subjective, as a general rule we look for 

rate decreases of at least 30 percent compared to the longterm average rate, using a duration of at least 12 to 18 

months to avoid seasonal effects. By this rule, the deficit 

at XHR in 1988-89 does not qualify as retardation due to 

its brevity, but we include it nevertheless because of its 

uniqueness in the long-term data set. On the San Andreas 

fault, four out of five creepmeters recorded slip retardations of some duration before the Loma Prieta earthquake, 

though there is little consistency in the onset. However, 

even in hindsight, it is not clear that these retardations are 

precursors to the Loma Prieta earthquake or that they could 

have been used to somehow predict the earthquake. Table 

2 lists retardation parameters for these episodes. At two 

of these sites, XHR and CWC, rate reductions end just 

Figure 4.-Continued 
D 152 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 2.-Possible creep retardations before the earthquake 

[The retarded rate and the background rate were determined by fitting least square lines to the intervals of interest. The background interval 

extends from the origin of the respective meter to 17 October 1989.1 

Site From To Duration (months) Rate (mmlyr) Background Rate (mmlyr) 

XSJ August 1982 April 1990 92 3.8 6.4 

XHR September 1988 September 1989 12 6.1 7.5 

CWC August 1987 October 1989 26 6.9 10.1 

XMR August 1988 May 1991 3 6 12.3 18.0 

XSH May 1989 continuing 30 -0.3 11.8 

before or with the quake. At XSJ and XMR however, slip 

rates continued to be lower than the long-term average 

until 6 months and 19 months after the quake. 

Of the Calaveras creepmeters, the rate at XSH began to 

drop in May 1989, while at HLC and HLD slip deficits 

are completely postseismic effects. These retardation episodes amount to only a shadowy precursor, given the variability of response across the network. However, since 

the Loma Prieta is the first M=7 quake in proximity to 

USGS creepmeters in almost 25 years of monitoring, it is 

still premature to discount these deficits as local, transitory effects without further rigorous examination. 

SIGNAL VERSUS NOISE IN THE CREEP 

DATA 

As is indicated in the above discussion, separating tectonic signals from rainfall-induced signals and noise is an 

important problem. Langbein and others (1993) have estimated the 1/f2 noise of the creepmeters near Parkfield, 

California, to have a standard deviation of 6 mmlyr. We 

have attempted to estimate the noise levels for the eight 

instruments discussed in this report by calculating average creep rates using 6-month and 1-year averaging windows moved across four intervals of interest (table 3). 

The standard deviations in table 3 give some indication 

of noise levels and how large rate changes need to be at 

particular sites before they can be regarded as significant. 

Although our analysis is simpler and not directly comparable to that of Langbein and others (1993), the results for 

XHR, CWC, and XFL show standard deviations consistently below the 6-mdyr level, suggesting that these instruments may have higher signal-to-noise ratios than 

typical Parkfield instruments. On the contrary, XSJ, XMR, 

XSH, and HLC have standard deviations that are typically 

worse than 6 mdyr, at least for some intervals. 

Another factor that helps convince us that rate changes 

at the time of Loma Prieta are real is the occurrence of 

similar changes on a number of instruments. Even though 

the changes on individual instruments might be called into 

question because of noise problems, the agreement of a 

suite of instruments in the sense and magnitude of rate 

change can be convincing. 

COSEISMIC STEPS AND STATIC STRESS 

CHANGES 

Coseismic steps were recorded at the six closest USGS 

creepmeters on the San Andreas and Calaveras faults with 

sampling intervals short enough (10 minutes) to detect 

such steps. These "coseismic" responses occurred within 

10-20 minutes of the earthquake origin time, the uncertainty in timing being caused by the coarseness of the 

sampling interval. Table 1 lists coseismic steps at all 

creepmeters in the network, and figure 5 shows raw data 

for the six closest instruments. 

Figure 6 compares the coseismic steps at the six 

creepmeters with the calculated Loma Prieta stress changes, 

based on a dislocation model of Lisowski and others (1990) 

used to match geodetic measurements of coseismic surface displacements. Static stress changes at the creepmeter 

sites were calculated using equations derived by Okada 

(1992) for dislocations in an elastic half-space. We show 

only horizontal shear stress changes and normal stress 

changes because the vertical shear changes are 5-10 times 

smaller on these vertical faults (Simpson and Reasenberg, 

1994). Coseismic steps could not be inferred for HLC or 

HLD because these instruments are read manually at quarterly intervals. Of six observed coseismic steps, one (at 

XFL) is LL and all of the others are RL. No strong relation exists between either sign or magnitude of the steps 

and sign or magnitude of the calculated stress changes, 

especially in view of the tendency of coseismic steps in 

RL fault systems to occur in a RL sense. 

Five earthquakes other than the Loma Prieta with magnitudes greater than 5 have occurred in the region since 

1974 (fig. 15); namely, Thanksgiving Day (M=5.1), Coyote Lake (M=5.9), Morgan Hill (M=6.2), Tres Pinos 

(M=5.3), and Chittenden (M=5.4). For all six earthquakes, 

4 LL and 17 RL coseismic steps were observed (table 4, 

figure 7). Of the 4 LL steps, two occurred at sites where 

model calculations predict LL shear-stress changes, and 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl53 

Table 3.-Average creep rates for five creepmeters on the Sun Andreas Fault and three creepmeters on the Calaveras 

fault 

[Average rate is determined by least square fitting of a line to data in either a 6-month window (A) or in a 1-year window (B); the 

windows are advanced by 3-month and 6-month increments, respectively, over the period of interest, and the resulting rates are 

averaged. The first interval is a background period which ends more than 3 years before the earthquake so as to avoid possible 

precursory Loma Prieta effects. The second interval includes possible precursory Lorna Prieta effects. The third interval is for 1 year 

after the earthquake. The fourth interval is for 2 years after the earthquake. Negative values indicate left-lateral rates; positive values, 

right-lateral.] 

A. 6-Month Average Creep Rates, in mmlyr 

[The 6-month sampling interval advances in 3-month steps. First number is average creep rate in mmlyr; 

second number after 5 is one standard deviation; value in parentheses is the number of six-month intervals 

within the overall interval that was used to calculate average and standard deviation.] 

Interval Dates: 01/01/77-01/01/86 01/01/86-10/01/89 11/01/89-11/01/90 11/01/89-11/01/91 

XS J 

XHR 

cwc 

XFL 

XMR 

XSH 

HLC 

HLD 

B. 1-Year Average Creep Rates, in mdyr 

[The 1-year sampling interval advances in half-year steps. First number is average creep rate in mmlyr; 

second number after k is one standard deviation; value in parentheses is the number of six-month intervals 

within the overall interval that was used to calculate average and standard deviation.] 

Interval Dates: 01101/77-01/01/86 01/01/86-10/01/89 11/01/89-11/01/90 11/01/89-11/01/91 

XSJ 

XHR 

cwc 

XFL 

XMR 

XSH 

HLC 

HLD 

two occurred at sites where calculations predict RL shear- These results imply that the sense (RL or LL) of 

stress changes. Conversely, at the six sites with coseismic coseismic steps at the six creepmeter sites bears little resteps where LL stress changes are predicted, only two lation to the imposed static shear-stress changes. If we 

showed LL coseismic steps. adopt the null hypothesis that the direction of a coseismic 
Dl54 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

step is independent of the sign of imposed static stress 

change, then a two-sided chi-square test (for example, 

Sachs, 1982) applied to the data displayed in figures 6 

and 7 does not allow us to reject the null hypothesis at 

even the SO-percent confidence level. We conclude that 

the sense of observed coseismic steps bears no obvious 

strong relation to the calculated static stress changes. Although all but one of the large coseismic steps did occur 

for RL static stress changes, there was only one instance 

of a large imposed LL static change, so the geometry of 

faults and earthquakes in the region has not lent itself to 

testing the hypothesis in question. 

It does appear from visual inspection of figure 7 that 

the 26 occurrences of increased RL shear were kore likely 

to have a sizable coseismic step than the 19 occurrences 

of added LL shear, many of which had a zero step and are 

not plotted in figure 7, The larger amplitude and the larger 

number of RL, steps presumably reflects the fact that all 

of the creepmeters are installed across faults that normally creep in a RL sense, 

More often than not, the coseismic step probably reflects the triggered release of a backlog of slip either in the 

near-surface of the earth or in the instrument itself by the 

shaking that accompanies the passage of seismic waves. 

The magnitude of the dynamic stress changes associated 

with the passage of seismic waves is, in general, much 

larger than the magnitude of the static stress changes 

(Spudich and others, 1995). On a right-lateral fault, such a 

backlog would normally be right-lateral, although a few 

sites might have a left-lateral backlog resulting from seasonal variations in slip direction of the instrument caused 

by rainfall, local soil conditions, and installation configuration. We cannot rule out the possibility that, because of 

radiation patterns and ground conditions, those sites with 

RL shear added also might have been the sites that experienced the greater amount of dynamic shaking during the 

earthquake. Nevertheless, the results shown in figure 7 

suggest that added RL static shear encourages the triggered 

release of a backlog of slip in a RL system, while added 

LL static shear can discourage such a triggered release, 

XSH 

XS J 

XHR 

cwc 

XFL 

XMR 

OCTOBER 

Figure 5.-Coseismic response to the Loma Prieta earthquake for six creepmeters. XSH is on the Calaveras fault, while the others are 

on the San Andreas fault. Time extends from one day before to one day after the Loma Prieta earthquake. 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAW D 155 

Williams and others (1988) suggest that the amount of 

slip triggered on part of the southern San Andreas fault 

by the 1986 North Palm Springs earthquake agreed well 

with the size of the slip deficit at the site. McGill and 

others (1989) suggest that slip deficit is probably only 

part of the story, because some sites showed triggered 

events associated with both the 1989 Elmore Ranch and 

Superstition Hills earthquakes, separated by only 11 hours, 

whereas one might expect the deficit to have been shaken 

out by the first shock. We looked for a correlation between the magnitude of the coseismic step and the slip 

deficit at our eight San Andreas and Calaveras sites, where 

slip deficit is defined as the difference between the longterm rate at a site and the average rate in the 1-year period preceding the earthquake (fig. 8). (The two points in 

figure 8 with large negative deficits come from longterm rates before the Chittenden earthquake-an aftershock to the Lorna Prieta earthquake. These negative deficits are caused by faster-than-nomal rates at the two sites 

caused by Lorna Prieta afterslip.) Although there is a suggestion that the largest steps correlate with the largest 

deficits as defined here, the strength of the correlation is 

not great. 

We could find no significant agreement between sign 

of coseismic steps and sign of 1-year average rate changes 

(fig. 9). The two-sided chi-square test did not allow us to 

reject the null hypothesis that no relation existed with a 

confidence level of any greater than 40 percent. The absence of any such a relation is consistent with the idea 

that, in most cases, coseismic steps largely reflect the 

release of a backlog of slip rather than a clean response to 

newly imposed stress changes, although it is not possible 

to rule out a combination of the two effects. 

CREEP RATE CHANGES AND STATIC 

STRESS CHANGES 

Average 1-year slip rates for the five creepmeters on 

the San Andreas fault increased at the time of the Loma 

HORIZONTAL SHEAR STRESS CHANGE, IN MBAR 

2 

z 

gzci 

LLl lz 

cmz 

5!? 

LLI 

rn 

0 

0 0. 

NORMAL STRESS CHANGE, IN MBAR 

-600 -400 -200 0 200 400 600 

I I I 

e 

XSH 

I I I 

2 

=om z - aLU 

I-- 

rn 

Gmz 

2 

LU 

m 

8 

Figure 6.4oseismic steps for Lorna Prieta earthquake at six creepmeters compared with calculated static stress changes. Coseismic steps are + for 

RL, - for LL. Shear stress is horizontal component, + for RL, - for LL. Normal stress is perpendicular component, + for tension, - for compression. 

I I I 

- 

0 

cwc - 0 0 

XHR XSJ 

XMR 

0 

XFL 

I I 

-600 -400 -200 0 200 400 600 

I I I 

. 

cwc 

. 0 

XSH XSJxHR . 

XM 

X FL 

I I I 

I I I 

- 

- 

F3 

I I I 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 4.-Observed coseismic steps and I-year average rates fur eight creepmeters and six earthquakes 

[The earthquakes in column 1 are ch = Chittenden, lp = Loma Prieta, tp = Tres Pinos, mh = Morgan Hill, cl = Coyote Lake, td = 

Thanksgiving Day. For coseismic steps in column 3, '-' indicates no step, and '?' indicates that the data were not available. I-year average 

rates were calculated by fitting a least-squares line to the data in the one year interval and ignoring any coseismic offset. Fractional rate 

change in column 6 is calculated from columns 4 and 5 as (rate before - rate after)/(rate before). Stress changes in columns 7 and 8 were 

calculated from dislocation models described in table 3 and in the text. Negative values indicate left-lateral changes or rates; positive 

values, right-lateral.] 

Earthquake Creep- Coseisrnic Step I-Year Rate Before 1-Year Rate After Fractional Shear Stress Normal Stress 

ch 

ch 

ch 

ch 

ch 

ch 

ch 

ch 

IP 

1P 

IP 

1 P 

1 P 

1 P 

IP 

[P 

t P 

t P 

tP 

tP 

t P 

t P 

tP 

t P 

mh 

mh 

mh 

mh 

rnh 

rnh 

rnh 

rnh 

cl 

cl 

cl 

cl 

cl 

cl 

cl 

cl 

td 

td 

td 

td 

td 

td 

td 

td 

XSH 

HLC 

HLD 

XSJ 

XHR 

cwc 

XFL 

XMR 

XSH 

HLC 

HLD 

XSJ 

XHR 

cwc 

XFL 

XMR 

XSH 

HLC 

HLD 

XSJ 

XHR 

cwc 

XFL 

XMR 

XSH 

H LC 

HLD 

XSJ 

XHR 

cwc 

X FL 

XMR 

XSH 

H LC 

HLD 

XSJ 

XHR 

cwc 

XFL 

XMR 

XSH 

H LC 

HLD 

XSJ 

XHR 

cwc 

X FL 

XMR 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE D 157 

Prieta earthquake, whereas rates for the three sites on the changes, three different dislocation models of the Lorna 

Calaveras fault decreased (figs. 2, 3; tables 3,4). To corn- Prieta rupture were used, based on geometries and slip 

pare creep-rate changes with calculated static stress distributions proposed by Lisowski and others (1990), 

HORIZONTAL SHEAR STRESS CHANGE, IN MBAR 

-600 -400 -200 0 200 400 600 

NORMAL STRESS CHANGE, IN MBAR 

Figure 7.-Coseismic steps for six earthquakes, including Loma Prieta earthquake, compared with calculated static stress changes. Sign conventions as in figure 6. 

-20 -1 5 -1 0 - 5 0 5 10 15 20 

SLIP DEFICIT FOR ONE YEAR, IN MM 

Figure 8.-Coseismic steps for six earthquakes compared with one-year slip deficit (defined as difference between the long term rate and the oneyear average rate prior to the earthquake). 
Dl58 AFTERSHOCKS AND POSTSEISMIC ElTECTS 

Marshall and others (19911, and Beroza (1991). Simpson 

and Reasenberg (table 1, 1994) tabulate the details of these 

three Loma Prieta models. For one site, changes in the 

rupture orientation from model to model changes the sign 

of calculated shear-stress because this site is near a node 

in the stress field. In general, the stresses calculated at 

distances of several rupture lengths from the epicenter are 

quite comparable model to model. 

Figure 10 compares the fractional change in 1-year average creep rates before and after the Loma Prieta earthquake with calculated static stress changes from the three 

Loma Prieta models, The fractional change in average rate 

is defined by AV I Vo where AV is the change in average 

rate over the given time window and % is the average 

rate before Loma Prieta over the given time window. 

To see if there was any significant relationship between 

the signs of the rate changes and the signs of the stress 

changes, we tested the null hypothesis that these quantities were independent by again applying a two-sided chisquare test to the respective fourfold tables. 

For shear stresses, the null hypothesis that the quantities are independent can be rejected at the 99-percent, 96- 

percent, and 83-percent confidence levels (x2 values of 8, 

4.4, 1.9) for the Lisowski, Beroza, and Marshall models, 

respectively, when the stresses are calculated at the surface. The confidence limits are even better if the stresses 

are calculated at greater depths (for example, 10 km) because the horizontal shear-stresses change sign under HLC 

and HLD in the Marshall model (see fig. 15 and table 6). 

For normal stresses, the null hypothesis that the quantities are independent can be rejected at the 0-percent, 52- 

percent, and 96-percent confidence levels (x2 values of 0, 

0.5, 4.4) for the Lisowski, Beroza, and Marshall models, 

respectively. The sense of relation for normal stresses suggested by the Marshall model would yield an increase in 

creep rate in response to negative values of normal stress 

which, in the convention used here, would imply more 

compression. If a relation of this sort exists, it would be 

counter-intuitive given that Coulomb's law predicts greater 

friction on the fault as a result of greater compression. 

We interpret these results to mean that for the three 

models tested, there is a good relation between slip-rate 

change and shear stress change, but that a relation between slip-rate change and normal stress change is unlikely. 

In an effort to put this relationship on a firmer footing, 

we repeated the statistical tests using data for four earlier 

earthquakes and the Chittenden aftershock, as well as for 

the Loma Prieta earthquake (table 4). Model ruptures for 

each earthquake were made using dislocation rectangles 

positioned and oriented from main shock and aftershock 

locations and from focal-mechanism information (table 

5). Amounts of slip were assigned to these dislocation 

surfaces to yield the observed seismic moments. By examining the results of six earthquakes together, we hoped 

that model-dependent effects would be either made apparent or minimized in the statistics. The data for all six 

earthquakes are shown in figure 11. For shear stress, there 

is a relation at the 95-percent confidence level (x2= 3.88). 

For normal stress, the null hypothesis of independence 

cannot be rejected, implying that no strong relation exists. 

A slightly sharper result is obtained if the lower calculated values of stress change are discarded. For example, 

Reasenberg and Simpson (1992, and this chapter) reported 

that significant correlation exists between shear-stress 

changes and microseismicity rate changes down to stress 

levels of about 0.1 bar. If we repeat the statistical tests, 

eliminating those data points for which the shear-stress 

changes were less than 0.05 bar (fig. 12), then for shear 

stresses, there is a relation at the 99-percent confidence 

level (x2= 6.58). These results are summarized in table 6. 

Similar statistical tests applied to each of the earthquakes individually show that, not unexpectedly, the largest earthquakes (Loma Prieta and Morgan Hill) give the 

FRACTIONAL CHANGE IN 1 -YEAR AVERAGE RATE 

Figure 9.-Coseismic steps for six earthquakes compared with fractional change in 1-year average creep rates. Fractional change is defined as the 1- 

year average rate before the earthquake minus the 1-year average rate after thcearthquake, divided by the 1-year average rate before the earthquake. 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl59 

most definitive results. The results of such tests are also 

dependent upon the choice of averaging interval. We used 

a 1-year interval for the tests described above. For a 3- 

year window, XMR yields a slower average rate after 

Loma Prieta compared to the rate before, which worsens 

the statistical results because the calculated static-stress 

changes require an increase in rate at XMR. It would be 

desirable to discover some objective criterion for choosing the averaging interval, but in the absence of such a 

criterion perhaps the best that can be done is to try a 

range of reasonable intervals to demonstrate that any results obtained do not depend strongly on the choice of 

interval length. 

The best overall relationship between fractional 1-year 

average rate changes and applied static stress changes from 

table 6 can be expressed as AV/V0=(8 bars-l)~~, where AT 

is in bars, although multipliers obtained range from 4 to 9 

bars-'. This is not a very general relationship because it is 

tied to a 1-year averaging window, but the data on the 

whole do not seem adequate to support a more complex 

relation. The data from XHR and CWC might be good 

enough to warrant a fit to some of the empirical relations 

described below, but we have not done this. 

COEFFICIENT OF APPARENT 

FRICTION 

In correlating microseismicity rate changes with Coulomb failure function changes, Reasenberg and Simpson 

(1992, and this chapter) found that the best correlation 

was obtained for low assumed values of the apparent coefficient of friction p'. (In the terminology used by 

Reasenberg and Simpson (this chapter), the apparent coefficient of friction is the value inferred by neglecting 

pore fluid pressure changes.) The absence of any significant relationship between creep-rate changes and calculated static stress changes reported in the previous section 

is consistent with a low value of p'. 

As another way to examine this, we correlated the Coulomb stress for different values of p' with the fractional 

rate change. Values of correlation coefficient p are weakly 

1000 -500 0 500 1000 1500 

HORIZONTAL SHEAR STRESS CHANGE, IN MBAR 

XSJ 

NORMAL STRESS CHANGE, IN MBAR 

Figure 10.-Comparison of the fractional change in 1-year average creep rate at the time of Lorna Prieta earthquake with calculated 

static stress changes for the three Lorna Prieta slip distributions: A, Lisowski and others (1990), B, Marshall and others (1991), C, 

Beroza (1991). Best fit line passing through the origin is also shown (see table 4 for slopes). 
Dl60 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

but systematically peaked at low values of the apparent 

coefficients of friction p' (fig. 13). The correlation coefficient p improves slightly if we delete all values of Coulomb stress less than 0.05 bars. In both cases the correlation 

is significant at the 95-percent confidence level, although 

the slight differences for neighboring values of p' in figure 13 are not likely to be very significant. 

CREEP RATE VERSUS STRESS LAW 

It would be very desirable to have a rheological law 

that would allow us to predict how a creeping fault would 

respond to applied stresses. The calculated stress distributions discussed in the next section suggest that inferring 

information about such a law from the response of a single 

creepmeter to a single earthquake will not be easy. 

Nason and Weertman (1973) pointed out that although 

the geometry of fault creep events requires an upper yield 

point behavior on the part of the fault gouge, a unique 

constitutive law cannot be inferred from the shape of the 

creep events alone. Nonetheless, a number of authors have 

attempted to infer parameters for various plausible types 

of creep laws by looking at creep events or afterslip decay. 

Crough and Burford (1977) used a power law faultzone rheology to relate the shapes of individual creep 

events with stress. The power law is 

and the resulting displacement for creep events is given 

by 

LL 

-1 500 -1 000 -500 0 500 1000 1500 

HORIZONTAL SHEAR STRESS CHANGE, IN MBAR 

6 -1500 -1 000 -500 0 500 1000 1500 

NORMAL STRESS CHANGE, IN MBAR 

Figure 10.-Continued 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl61 

where U(t) is the displacement at time t after the onset of 

the creep event, t is the driving stress, Uf is the final 

displacement, C is a constant of proportionality, and n is 

the power law exponent. Wesson (1987, 1988) used this 

law for his discussions of fault dynamics and Bilham 

(1989) used it to fit afterslip data to the 1987 Superstition 

Hills, California, earthquake. Although the shapes of the 

Loma Prieta perturbation at XHR and CWC (fig. 14) approximate the shapes of the classic creep event, the shapes 

of the perturbation at the remaining six sites are far from 

classic. Wesson (198 8) has simulated composite events 

with somewhat similar shapes to those seen at XSJ and 

XMR in a model using multiple interacting slip patches, 

each obeying a power-law rheology. 

We note that the power law rheology as written, if applied to the fault at all depths, does not take into account 

the possibility of a sign change in T at some depths that 

Ill 

< 

a 

0. 

UJ 

Ill 

u 

0 

z 

Ill 

a 

I 

0 

_J 

a 

2 

0 

I0 

ec. 

might temporarily reverse the creep direction. A better 

version might be written 

For small changes in stress, the power-law yields 

where AV is the change in creep rate and AT is the change 

in driving stress, which motivated our use of the fractional creep rate change in the previous section. Comparison with observations (table 6) yields the result that n/~=8 

bars-l. 

u -1 500 -1 000 -500 0 500 1000 1500 

HORIZONTAL SHEAR STRESS CHANGE, IN MBAR 

c -1500 -1 000 -500 0 500 1000 1500 

NORMAL STRESS CHANGE, IN MBAR 

UJ 

Figure 10.Ã‘Continue 

I- 

<au 

0. 

am-- 

u 

0 

2"*- 

UJ 

ONa 

I 

0 0- _1 

a 

59- 

I0 

a 7,- 0- 

XSJ 

1 I I I I 
Dl62 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 5.-Values used in dislocation models of the six earthquakes 

[Negative horizontal-slip value indicates left-lateral offset; positive values, right-lateral. Negative dip-slip value indicates reverse faulting. Earthquake information from Oppenheimer 

and others (1990).] 

date name M no. Ion lat depth len. ht. strike dip horiz dip 

(deg) (deg) (km) (km) (km) (deg) (deg) slip (m) slip (m) 

74 1 128 Thanksgiving Day 5.1 1 121.46 36.92 4.5 1 1 33 90 -1.5 0 

790806 Coyote Lake 5.9 1 121.53 37.10 6.0 13 8 335 90 0.2 0 

840424 Morgan Hill 5 6.2 324 121.61 37.22 - 27 12 328 90 0.245 0 

860126 Tres Pinos 5.3 1 121.28 36.80 7.5 1 1 353 83 3.0 0 

89101 8 Loma Prieta f 7.0 1 121.91 37.06 11.2 37 13.3 136 70 1.66 -1.19 

90041 8 Chittenden 5.4 1 121.64 36.94 4.8 1 1 312 90 4.2 0 

M = local magnitude 

no. = number of elements in dislocation model 

depth = depth to center of model 

len. = horizontal length of rectangular dislocation surface 

ht. = down-dip dimension of rectangular dislocation surface 

strike, dip = orientation of rectangular dislocation surface 

horiz. slip = component of slip in horizontal direction 

dip slip = component of slip in down-dip direction 

5 Model based on Harzell and Heaton (1986) used by Oppenheimer and others (1988). Slip value is average. 

f This is the Lisowski model described in Simpson and Reasenberg (1994) 

We can infer a value for the "average" ambient stress 

on the fault plane under the creepmeters by substituting 

likely values for n. Crough and Burford (1977) report 

values of n between 1.0 and 2.5 with an average of 1.6 

inferred from typical creep events. For such values, T 

would lie between 0.1 and 0.3 bars, which seems quite 

small but which may represent average stress levels at the 

shallow depths most likely to be reflected in the first year 

of creep rate change. 

Sharp and Saxton (1989) proposed an empirical law to 

describe the time-evolution of afterslip observed on the 

Superstition Hills fault after two earthquakes in November 1987 

where U(t) is displacement at time t after the earthquake, 

and U', B, and c are constants. (L. Wennerberg, oral 

commun., 1993, has proposed a refined and better-fitting 

version of this empirical law.) Boatwright and others 

(1989) discuss an inversion method to infer the three parameters in the law from field data. 

Scholz (1990) and Marone and others (1991) assumed 

a relation based on constitutive laws developed by 

Dieterich (1979, 1981), Ruina (1983), and Rice and Gu 

(1983) to explore the nature of afterslip curves. Their starting constitutive relation is 

where T~~ is steady state frictional strength, V is sliding 

speed, A-B is the friction rate parameter, and T* is the 

strength for steady state sliding at speed V*. The afterslip 

displacement Up(t) at time t after the earthquake is given 

by 

A-B up~t~=-J[~)t+i]+~, k A-B 

where k is the thickness averaged stiffness for the velocity strengthening region, V& is the thickness-averaged 

coseismic slip velocity within the velocity strengthening 

region, and V. is the long-term slip rate. 

There is a difficulty in applying the constitutive relation when V becomes zero or negative. A more general 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl63 

formulation of the constitutive relation suggested by 

Dieterich (for example, 1992) can allow for these possibilities. 

At depths of several kilometers, we would expect that 

the imposed shear-stress changes under our creepmeter 

sites (typically tenths of bars or less) are small compared 

with the usual shear-stress levels driving these faults. If 

so, then a linearized form of the above laws relating creeprate changes and stress changes may be appropriate, and 

the data provided by the creepmeters may not be adequate 

to distinguish among the various laws or, in fact, to distinguish them from a linear viscous response. The fact 

that the three sites that were close to regions where the 

fault surface had LL shear-stress applied actually went 

LL for varying periods after Loma Prieta suggests that the 

usual shear-stress levels in the upper meters or kilometers 

are normally quite small, and that the rheological laws 

governing this depth range on the fault can probably not 

be linearized. The behavior of the Middle Mountain 

creepmeter near Parkfield after the Coalinga earthquake 

(Mavko and others, 1985; Simpson and others, 1988) provides another example of such behavior. 

Clearly, more work needs to be done to propose and 

test creep-rate laws with the appropriate depth-dependent 

behaviors, in order to explain the perturbations at these 

creepmeters. One benefit of such a law is the ability to 

calculate synthetic creep records at various sites based on 

the influences of nearby earthquakes. When such a law is 

calibrated, it can also be used to infer the existence of 

stressing not obviously associated with seismicity, but perhaps caused by the passage of aseismic tectonic waves. 

Gwyther and others (1992) describe a post-Loma Prieta 

shear-strain anomaly near San Juan Bautista recorded on 

a tensor strainmeter that could perhaps be used in conjunction with creep data to constrain a fault rheology. 

If indeed the creep rates are responding to static stress 

changes caused by earthquakes, including earthquakes on 

other faults, this raises the question as to whether the 

effect is shallow or extends to depth. The c eepmeters, if 

they are sitting over low-friction shallow racks, could 

just be acting as sensitive strainmeters. Th 1 observation 

that microseismicity rates on Bay Area faults also responded to static stress changes after the Loma Prieta 

earthquake (Reasenberg and Simpson, 1992), suggests that 

HORIZONTAL SHEAR STRESS CHANGE, IN MBAR 

NORMAL STRESS CHANGE. IN MBAR 

Figure 11 .-Comparison of the fractional change in 1-year average creep rate at the times of six earthquakes with calculated static stress changes for 

simple dislocation models of the earthquakes. 
Dl64 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

some part of the observed changes in surface creep-rates 

reflect changes in slip rate at seismogenic depths. 

We have not attempted to fit any of these laws to the 

Loma Prieta perturbations displayed in figure 14, partly 

because these perturbations are not very cleanly defined, 

and partly because we believe that the perturbations represent a composite response to sometimes complex stress 

distributions. We will instead attempt to use some simple 

dislocation models to put bounds on the total amount of 

perturbed slip that might occur at the various sites, and to 

estimate the depths from which the slip might be coming 

as a function of time. 

MODELS EXPLORING DEPTH-ORIGIN 

OF LOMA PRIETA AFTERSLIP 

Although it is difficult to infer fault-zone rheology from 

the available "afterslip" observations, it is possible to estimate the depth to which the earthquake-induced anomalous slip extends. It might be the case, for example, that 

the afterslip recorded at the creepmeters was a superficial 

phenomenon, confined to the upper kilometer or two of 

the crust. 

To bound the possible depth to which anomalous slip 

might extend, we used Okada's (1992) dislocation subroutines and the Loma Prieta slip distribution of Marshall 

and others (1991) to make a model of the San Andreas 

and Calaveras faults (fig. 15). Stress changes were calculated at the centers of 2 km by 2 km square dislocation 

patches extending down to 20 km. We permitted the dislocation squares between the Earth's surface and some 

chosen depth to slip freely in response to the earthquakeinduced static stress changes, so that the stress at their 

centers was canceled. All squares below the chosen depth 

were not allowed to slip. No squares farther than 9 km 

north of creepmeter XSJ were allowed to slip in response 

to the stress chahges. We assumed that the total slip in 

response to the stress changes would occur instantaneously, 

although, in fact, the slip must occur in viscous fashion 

over the space of several years. 

The model contains an interruption in the Calaveras 

fault between HLC and HLD. This discontinuity, although 

1000 -500 0 500 1000 

HORIZONTAL SHEAR STRESS CHANGE, IN MBAR 

-500 0 500 

NORMAL STRESS CHANGE, IN MBAR 

Figure 12.-Comparison of the fractional change in 1-year average creep rate at the times of six earthquakes with calculated static stress changes 

greater than 0.05 bars in amplitude for simple dislocation models of the earthquakes. 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE D 165 

Table 6.-Statistical tests to determine significance of relation between changes in one-year average creep rates and calculated 

changes in static stress 

- - 

Earthquake Stress Depth Two-sided Confidence Rho N Conf Slope0 Slope Intercept 

Model(s) Component (km) Chi-sq (%I (96) (bar") (bar") 

L horiz. shear 0 8.00 99 0.96 8 99 8.38 8.16M.97 0.19k0.29 

B 4.44 96 0.90 8 99 4.16 4.14M.83 0.03k0.48 

M ,I ,I 1.90 83 0.87 8 99 4.75 4.74k1.08 0.01k0.53 -- 

L normal 0 NA 0 0.18 8 <SO -1.32 2.1W4.80 1.39k1.29 

B 0.53 5 3 -0.71 8 95 -6.55 -7.04k2.88 1.19k0.70 

M ,I 9, 4.44 96 -0.50 8 <80 -5.50 -4.77k3.35 0.53k0.92 

L horiz. shear 10 8.00 99 088 8 99 6.50 6.34k1.36 088k0.46 

B ,, 8.00 99 0.86 8 99 4.38 4.17k1.02 0.71k0.51 

M 9, t, 8.00 99 0.85 8 99 5.05 4.79k1.21 0.66k0.52 

L normal 10 NA 0 0.30 8 '50 -0.03 2.48k3.22 1.54k1.16 

B I, I1 0.53 53 0.13 8 <SO 0.28 0.78k2.50 1.14k1.06 

M ,, 1, 1.90 83 008 8 <SO 0.74 0.57k2.79 1.12k1.10 

ALL horiz. shear 0 3.88 95 0.48 48 99 9.35 8.81k2.36 0.6W0.34 

ALL>0.05 I, I, 6.58 9 8 0.48 25 98 9.39 8.38k3.23 1.06k0.64 

ALL normal 0 0.30 41 0.07 48 <50 -0.82 1.58k3.50 0.83k0.41 

ALL>0.05 ,, t, 0.88 65 0.00 14 0 -0.88 0.02k3.16 0.35k0.68 

ALL horiz. shear 10 6.50 99 0.45 48 99 8.51 8.09k2.36 0.66k0.34 

ALL>0.05 ,, If 9.38 99 0.46 20 96 8.48 7.72k3.52 1.45k0.79 

ALL normal 10 0.23 36 0.08 48 <SO 0.24 1.48k2.57 0.82k0.39 

ALL>0.05 ,, It 0.20 34 0.13 11 <SO 0.28 1.06k2.64 0.52k0.84 

Models: L=Lisowski, B=Beroza, M=Marshall, ALL=Lisowski plus 5 other earthquakes, ALL>0.05 = same as ALL except with stress 

changes less than 0.05 bar omitted. 

Depth = depth at which static stress changes were calculated. 

Two-sided Chi-sq = Confidence that the fractional creep rate and change in stress component are not independent based on the twosided Chi sqaured test applied to the respective four-fold table (Sachs, 1982). 

Rho = Correlation coefficient for fractional creep rates and static stress change values. 

N = number of samples. 

Conf = Confidence in best fit line. 

Slope0 = Slope of best fit line forced to pass through the origin. 

Slope, Intercept = Parameters for best fit line not forced to pass through the origin. 

small, will reduce the total slip estimated for nearby sites 

because of the fault connectivity effect of Bilham and 

Bodin (1992). (Sites on model faults that are distant from 

fault ends can slip farther than points close to ends of 

discontinuous segments.) 

The results (table 7, fig. 16) show, for example, that at 

CWC, a maximum of 5 mm of anomalous surface slip 

will occur if only the upper 2 km of the fault are able to 

respond to the stress changes. Figure 14 shows that slip at 

CWC had exceeded the amount expected at the pre-earthquake creep rate by this amount within several months of 

the earthquake, even if the coseismic step is not included. 

Compared to the creep rate in the year before the earthquake, at the beginning of 1993, anomalous excess slip of 

30-35 mm had occurred at CWC (fig. 14), requiring 

anomalous slip to have occurred to depths in excess of 10 

km in our model. Note that anomalous slip at some sites 

is highly dependent on the choice of background interval 

(compare last two columns in table 7). 

Figure 16 suggests that the total anomalous slip at the 

surface would not get much larger if even deeper levels 

were allowed to slip. We anticipate that the results obtained from a model where slip could extend to 50 km or 

100 km would not be greatly different from the results 
Dl66 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

observed at 20 km. As of 1 July 1992, observed anomalous slip advances and deficits were estimated to range 

from -25 to +45 mm for the 8 creepmeters on the San 

Andreas and Calaveras faults (table 7). These observations fall within the calculated extremes of -28 to +60 

mm in the 20-krn column of table 7. Because the anomalous slip is continuing, further comparisons will be needed. 

The three-dimensional distribution of earthquake-induced shear-stress in our model (fig. 15) suggests that 

efforts to infer fault-zone rheology from afterslip behavior recorded at the Earth's surface need to take into account the geometry of the stress field both laterally and 

with depth on the fault surface. For example., using the 

slip distribution of Marshall and others (1991), the stresses 

imposed by the earthquake close to HLC and HLD are RL 

in the upper 6 km and LL from 6 to at least 20 km (fig. 

15). The expected signal at these sites would initially be 

an increase in creep rate as the upper levels of the fault 

move faster in response to the added RL shear, followed 

by a slowing creep rate as the LL shear imposed at greater 

depths and to the north retards creep in these regionswhich eventually propagates to the creepmeter sites. One 

can see suggestions of such behavior in figure 14 at sites 

HLC and HLD, although the noise level in these records 

is large enough to cast some doubt on this interpretation. 

RETARDATION AND ANOMALOUS 

BEHAVIOR BEFORE THE 

EARTHQUAKE? 

A number of possible precursors to the Loma Prieta 

earthquake have been suggested. For example, FraserSmith and others (1990) reported anomalous electromagnetic radiation in the days and hours before the earthquake. 

Gladwin and others (1991) described a strain anomaly 

near creepmeter XSJ that began mid-way through 1988. 

Galehouse (this chapter) suggested that some alignment 

arrays along the Hayward fault might have slowed down 

in the months before Loma Prieta, and Reasenberg and 

Simpson (this chapter) reported a possible slowdown in 

Hayward fault seismicity beginning in 1988. 

We have listed five possible creep retardations (table 2, 

fig. 4) that might have foreshadowed the Loma Prieta 

earthquake. We examined the retardation at CWC in detail because the behavior of CWC (fig. 17) had been so 

consistent from about 1975 to 1987 that the retardation 

appears more convincing than do some of the others. We 

cannot rule out the possibility that drought conditions caused 

the slowdown at this site, but in the following discussion 

we assume a tectonic origin and see where that leads. 

I I I 1 # I 

0.0 0.2 0.4 0.6 0.8 1 .O 

APPARENT COEFFICIENT OF FRICTION 

Figure 13.-Correlation coefficient as a function of apparent coefficient of friction for correlation of fractional changes in 1-year creep rates with calculated shear stress change. Curve 1 represents results for all 

data. Curve 2 is for data with Coulomb failure function values greater than or equal to 0.05 bars. 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE Dl67 

An important observation is that in spite of possible 4), suggesting that whatever is happening at CWC is mostly 

retardations identified at XHR and XSJ (4 and 15 km local in scope. 

northwest of CWC, respectively), the anomaly at CWC The Tres Pinos earthquake of 26 January 1986 ought to 

does not correspond very well in duration or character have increased the static shear-stress at CWC by about 

with the behaviors of neighboring creepmeters (figs. 2, 3, one-third as much as the Loma Prieta event (table 4), and 

Detrended Using 1 Year Before 

XSJ 

XHR 

cwc 

XFL 

XMR 

XSH 

HLC 

HLD 

Figure 14.-Plots of the Lorna Prieta perturbation for eight creepmeters with 1-year average rates for 

the year before the Lorna Prieta earthquake subtracted to enhance the perturbations. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

our model suggests that these stress changes at CWC 

should have increased the creep-rate just as the Loma 

Prieta earthquake did, rather than retard it. (A small increase in rate is apparent immediately after the Tres Pinos 

earthquake, but it was soon followed by the more obvious 

decrease.) 

Another clue is offered in the detailed behavior of CWC 

during part of the retardation period (fig. 18). Although 

discrete RL events continued at about the same rate and 

size as before, a trend change in background movement 

beginning soon after a M=3 earthquake on 871 102 carries 

the instrument in a LL sense between the RL events. Histograms of creep step sizes for part of this period and for 

a more normal period are shown in figure 19. Left-lateral 

movement between RL events at CWC also took place 

between 1978 and 1981, but such behavior has not been 

the norm. 

Bilham and Behr (1992) proposed a two-layer model 

for creep observed on the Superstition Hills fault. They 

described episodic creep events superimposed on a more 

constant background slip rate, and they proposed that the 

episodic slip is coming from deeper layers on the fault, 

while the background slip represents steady sliding in a 

near-surface layer. One plausible explanation for the CWC 

retardation would be to have the RL events produced by 

slip propagating to the surface from some region on the 

San Andreas fault under or adjacent to CWC. The LL slip 

between RL events would be produced by stress transferred to the San Andreas fault under CWC by slip on 

some other fault. The most obvious source is the Tres 

Pinos earthquake, which was followed by a diffuse set of 

aftershocks (fig. 20), suggesting regional stress changes 

in the complex region between the San Andreas and Calaveras faults (Burford and Savage, 1972). Afterslip on the 

Tres Pinos fault plane itself does not produce LL stresses 

near CWC with our model geometry. But if the Tres Pinos 

earthquake triggered RL slip on parts of the subparallel 

Paicines fault, the resultant stresses at CWC could be LL. 

Figure 21 shows regions on nearby vertical faults where 

RL slip (red) would produce LL stress at CWC. 

A second scenario would be to have the Tres Pinos 

earthquake induce slip on some other structure even closer 

to CWC that could again induce LL stress changes on the 

San Andreas fault in the near surface under CWC. Small 

earthquakes define several such structures just to the north 

and east of CWC (fig. 20). Activation of a small structure 

close to CWC could explain why the effect is not seen 

with the same character on other nearby creepmeters. The 

next creepmeter 4 lun to the north (XHR) might in fact 

see stresses of opposite polarity depending on the geometry and location of the slipping region. Future models 

will be constructed to test the magnitudes of slip needed 

on various nearby faults to produce the observed rates of 

LL stepping. 

A third scenario would have regional, mostly aseismic 

slip on a sub-horizontal detachment structure following 

XSH 

fc~~ HLC 

HLD 

XSJ 

Sm Andmas Fault 

AS -0.4 -0.3 -0.2 -0.1 0.0 0.1 _ .J 0.4 1.- 

HORIZONTAL SHEAR SThkSS, IN BARS 

Figure 15.Ã‘Geometr of one model used to estimate total size of post- plotted in color. Red colors indicate right-lateral stress changes, blue 

Loma Prieta slip perturbation for various depths to the bottom of the colors indicate left-lateral changes. Each small rectangle is approximately 

freely slipping layer. This is an oblique view from the southwest, with 2-km by 2-km in size. The fault planes in this model extend to a depth of 

the San Andreas fault south of the Loma Prieta rupture in the foreground, 20 km. Locations of eight creepmeters discussed in the text are shown by 

and the Calaveras fault in the background. Horizontal shear stresses cal- triangles. 

culated using the Marshall and others (1991) slip distribution are shown 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE 

Table 7.-Total post-Loma Prieta anomalous slip, in millimeters, at eight creepmeters predicted 

for various depths to the bottom of the freely slipping layer 

[Calculations based on a dislocation model using the Lorna Prieta slip distribution of Marshall and others 

(1991). Negative values indicate left-lateral slip; positive values, right-lateral. The values shown in the last two 

columns are observed anomalous slip as of 1 July 1992 relative to a 1-year average rate before Lorna Prieta 

(figure 14) and relative to a 3-year average rate before Lorna Prieta.] 

-- - 

depth-> 2km 4km 6km 10km 20km Observed(1) Observed(3) 

XSJ 18.8 30.8 39.1 49.4 60.3 30 25 

XHR 6.2 12.1 17.7 27.5 41.8 45 3 3 

CWC 4.7 9.0 13.5 21.9 35.4 3 0 3 3 

XFL 1.9 3.5 5.4 9.7 19.4 7 7 

XMR 1.2 2.1 3.1 5.6 12.2 20 -10 

XSH -7.0 -11.4 -14.6 -19.7 -28.6 -7 -25 

HLC 0.1 -1.2 -2.8 -5.9 -10.5 5 -15 

HLD 0.2 -0.7 -1.6 -3.3 -5.7 - 5 - 5 

- 

0 5 10 15 20 

DEPTH OF SLIP. IN KM 

Figure 16.-Plot of total post-Loma Prieta anomalous slip predicted at eight creepmeter sites as a function of the depth to the bottom of the freely 

slipping layer in model. 
D 170 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

the Tres Pinos earthquake. Again, slip in the proper re- scale stress changes that could also have triggered the 

gions of such a structure could probably induce LL stress Loma Prieta earthquake. 

at CWC. Perhaps it is significant that the region south of San 

In all of these scenarios we suggest that the Tres Pinos Juan Bautista in which most creep retardations have been 

earthquake caused or accompanied regional stress changes &.scribed (Burford, 1988) contains subparallel fault strands 

in the triangular wedge between the Sari Andreas and Ca- (San Andreas and Paicines/Calaveras). If RL slip were to 

laveras faults which could have been recorded at CWC as alternate on these strands, then creepmeters on the movLL drift. Thus CWC retardation might have reflected large- ing strand would speed up, while those on the other would 

Figure 17.-Plot of creep record for CWC showing times of earthquakes that might have had an 

effect at this site. Earthquakes were selected from the CALNET catalog by scaling their moment by 

distance-3, which is the fall-off of maximum stress from a point dislocation, disregarding orientation. 

Numbers at tops of lines compare potential impact of the earthquake's static stresses (without regard 

to orientation information, so the value is an upper bound) to the impact expected from an optimally 

oriented magnitude 4.0 earthquake at a distance of 10 km. 
RESPONSE OF U.S. GEOLOG ICAL SURVEY CREEPMETERS TO THE 

slow down. Perhaps when the slowed strand begins to 

move again the fault might have become more brittle because it has had a chance to heal, so that if sufficient 

stress has accumulated to produce an earthquake, then 

one would be more likely to occur at the end of a retardation period than at other times. 

Burford (1988) suggested fault interactions of this sort 

as a possible explanation for the observed retardations. 

He also offers other explanations for the phenomenon, 

including growth of asperities, strain hardening, stress 

waves, and fluctuation in driving stress. There appears to 

be some hope of evaluating these possibilities using simple 

dislocation models to explore plausible sources of imposed stresses. 

LOMA PRIETA EARTHQUAKE 

CONCLUSIONS 

Four mechanisms have been proposed to explain triggered slip on faults (Allen and others, 1972; Fuis, 1982; 

Williams and others, 1988; McGill and others, 1989): (1) 

static stress changes produced by the earthquake rupture, 

(2) dynamic stresses from the passage of seismic waves, 

(3) creep migrating from the earthquake source region, 

and (4) a regional strain event that produces both aseismic 

slip on some faults as well as earthquakes on others. 

The Loma Prieta earthquake produced coseismic steps 

on many of the central California creepmeters. We think 

that these steps were caused by shaking of the sites and 

Figure 18.-CWC creep record from 1986-1989 showing pre-Lorna Prieta retardation beginning 

in mid-1987. Times of the Tres Pinos and Loma Prieta earthquakes are indicated by vertical 

lines. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

-- 1 

-0.2 -0.1 0.0 0.1 

CREEP STEP SIZE, IN MM 

CREEP STEP SIZE, IN MM 

Figure 19.-Histogram of creep step sizes smaller than 0.25 mm from daily data between 1 October 1987 and 1 October 

1988 (bottom) compared with a similar histogram for the period from 1 January 1981 to 1 January 1982 (top). 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE D 173 

the instruments because they do not seem to correlate this chapter). These changes are generally consistent in 

very well in either size or direction with calculated static magnitude and sign with the static shear-stress changes, 

stress changes, favoring explanation (2) above. and statistically significant correlation exists for three difThe Loma Prieta earthquake produced significant ferent models of the Loma Prieta rupture, although the 

changes in average creep rate at a number of sites on the quality of the correlation varies to some degree from model 

San Andreas, Calaveras, and Hayward faults (Galehouse, to model. The change in horizontal shear-stress appears to 

Figure 20.-A, Earthquakes in the area of creepmeter CWC during the period from 1983-1988 from the CALNET catalog. B, 

Earthquakes for the same area from 1 January 1987 to 1 September 1988 during the period of retardation observed on 

creepmeter CWC when left-lateral steps were most obvious. Earthquakes do not lie exactly under the fault traces, probably in 

pan because of deficiencies in the velocity models and in part because the faults may not be vertical. A star marks the epicenter 

of the 26 January 1986 Tres Pinos earthquake. 
D 174 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

be the significant variable. Changes in calculated normal 

stress do not seem to correlate at significant confidence 

levels, which suggests that coefficients of apparent friction are low for creep on these faults. A comparison of 

correlation coefficients for various assumed values of apparent coefficient of friction finds the best correlation for 

low values in the range 0.0-0.3. These observations seem 

consistent with explanation (1) above. Explanation (3) 

might explain rate changes for creepmeters on the San 

Andreas fault, but hardly seems to explain the rate changes 

on the Calaveras or Hayward faults. 

A three-dimensional boundary element model was used 

to examine the depths that were being sampled by surface 

slip. This model can only put approximate limits on the 

depths that slip is "coming from" but suggests at CWC, 

for example, that by 1 year after the earthquake at the 

latest, slip from 10 km depth was being sampled at the 

surface. Total expected anomalous slip at the creep sites 

... . . ...... 

. . *. - *. .o 

, . 

XHR . rn ' . :. 

PACIFIC \ 

OCEAN \, 

Figure 20.-Continued 
RESPONSE OF U.S. GEOLOGICAL SURVEY CREEPMETERS TO THE LOMA PRIETA EARTHQUAKE 

can also be estimated from this three-dimensional model. 

The estimates range in value from -6 mm to +60 mm, 

which are in fair agreement with observed slip deficits 

and advances observed at the sites. 

Rainfall-induced and seasonal variations in creepmeter 

behavior are considerable and raise the possibility that 

some rate variations that might be interpreted as precursors to earthquakes are weather related. Because the 

drought conditions starting in 1987 were especially severe, creep retardations observed at four sites are suspect 

to some degree. We attempted to use creep behavior during the earlier drought years of 1976-1977 to calibrate the 

more recent drought but found no strong link between 

climate and slip fluctuations. 

Burford (1988) has proposed that creep retardations can 

occur at creep sites before local earthquakes and has tabulated 25 instances of possible retardations on the San Andreas and Calaveras faults between 1957 and 1983. 

Assuming that a possible retardation at CWC that began 

in 1986-1987 and ended with the Loma Prieta earthquake 

might be of tectonic origin, we considered some possible 

tectonic causes of this retardation. The instrument continued to record large RL events during this interval but 

small LL slip events in the intervals between the large RL 

events slowed the total creep-rate. This behavior is similar to that described by Bilham and Behr (1992), who 

ascribe the large creep steps and the background creep at 

creepmeter sites on the Superstition Hills fault to different sources. If such an explanation holds here, it would 

seem likely that the large RL steps that represent fairly 

normal behavior at CWC come from slip on nearby parts 

of the San Andreas fault. We propose three scenarios for 

the origin of the LL drift: slip triggered on the Paicines 

fault, slip on a nearby structure to CWC revealed in the 

seismicity, or slip on a regional subhorizontal detachment. 

In all three scenarios, this movement would be triggered 

by regional tectonic adjustments following the 26 January 

1986 Tres Pinos earthquake in the complex triangle 

(Burford and Savage, 1972) between the San Andreas and 

Calaveras faults. If this is correct, then the retardation at 

CWC could be regarded as a precursor to the Loma Prieta 

earthquake if these same adjustments ultimately brought 

that earthquake closer to fruition. This scenario would 

also seem to give some credence to explanation (4) above 

as a viable mechanism. 

The ability of the creepmeters to respond to Loma Prieta 

stress changes in predictable ways suggests to us that other 

regional stress-change information is contained in the 

creepmeter signals. It becomes increasingly important to 

understand the effects of weather and seasons on the instruments so that the true signals of tectonic origin can be 

extracted and interpreted. 

ACKNOWLEDGMENTS 

This paper was greatly enhanced by discussion with 

Paul Reasenberg, Evelyn Roeloffs, John Langbein, Roger 

Bilham, Bob Burford, and Jeff Behr. We are also grateful 

to Rich Liechti, whose expertise in creepmeter maintenance ensures a foundation for this and other studies, and 

to Jon Galehouse and his students at San Francisco State 

Lorna Prieta Rupture 

INFLUENCE ON CWC, IN BARSM 

Figure 21.Ã‘Obliqu view from the northeast with the Calaveras fault in shown in the figure, LL slip would be required to produce LL stress at 

the foreground and the San Andreas fault in the background, showing CWC, whereas on the Calaveras fault, there are regions (red) where RL 

influence coefficients relative to the rectangular patch under creepmeter slip can produce LL stress at CWC. Each small rectangle is approxiCWC. Red indicates fault patches where RL slip would induce LL stress mately 2-km by 2-l^m in size. The fault planes in this model extend 

changes at CWC; blue indicates patches where RL slip would induce RL vertically to a depth of 10 km. 

stress changes at CWC. Note that on the parts of the San Andreas fault 
D 176 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

University for reading the creepmeters manually during 

their surveying expeditions. 

APPENDIX A: SITE NOTES 

The following section provides further detail on conditions influencing the data from different creepmeters. These 

aspects are typically site-specific and modulate the way 

we interpret the long-term record. As such they form the 

basis of several assumptions in this paper. 

XHR: COMBINED DATASETS 

The record for XHR used in the study is a composite of 

data from XHRl and XHR2. Zero was lost on creepmeter 

XHRl when it was destroyed in 1984 (Schulz 1989). To 

estimate a projected starting point for XHR2 data we calculated the amount of movement in a 245 day window, 

which is the duration of the gap between instruments, 

sliding the window every 10 days. The greatest change in 

a window with continuous data was 9.4 mm for the period 

ending 81 1102. We chose instead a correction value of 

5.4 mm, which occurred most often in the series of windows. Comparison with creep data from Cienega winery, 

3 km southeast of the site supports this as a reasonable 

adjustment for the level of activity at the time. 

CWC: INSTRUMENT CHARACTERISTICS AND 

RAINFALL 

At the winery, creepmeters actually measure offset of 

adjacent concrete floor slabs separated by the fault trace. 

Individual rainstorms produce only nominal changes in 

creep at this site. It is possible that a sustained drought 

might be expressed as a decrease in the background slip 

pattern such that episodic events, an indication of slip at 

depth, would be of normal amplitudes, while inter-event 

slip could be more sensitive to changing conditions in the 

shallow soil. De-coupling between the instrument and the 

fault at this site makes the association between drought 

and slip deficit tentative at best. In fact, we could argue 

that the retardation seen at the winery from 1987 to 1989 

is caused by the same phenomenon that produced the slip 

deficit at XHR from late 1988 to September 1989. Perhaps both sites were responding to drought. Alternatively, 

these instruments may have been sensing a local perturbation in fault activity that was overridden by the Loma 

Prieta earthquake. 

Another instrument at the winery, CWN, is not included 

in this study. At this creepmeter, located about 30 meters 

northwest of CWC, an obstruction gradually developed 

inside the instrument enclosure which inhibited the amount 

of movement recorded. Approximate onset of this condition is difficult to determine, and the slip released when 

the problem was corrected in February 1990 was insufficient to resolve the discrepancy between what was recorded on CWN and what was observed on CWC. Prior 

to 1987 the two instruments tracked each other very well, 

both in rate of slip and in creep event characteristicsonset, duration, and amplitude. 

XSH: SITE MODIFICATIONS AND RAINFALL 

In spring 1986 a new creepmeter was installed at Shore 

Road, slightly south of the original site and spanning an 

additional five meters of fault zone (Schulz 1989). The 

anchor pier of the new instrument is now within a few 

meters of an adjacent slough embankment. The stronger 

RL movement occurring during winter and spring may be 

related to increased movement of the anchor pier during 

wet weather. The fact that this pattern of activity was not 

seen at the site before 1987 adds credence to this account. 

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Spudich, Paul, Steck, L.K., Hellweg, Margaret, Fletcher, J.B., and 

Baker, L.M., 1995, Transient stresses at Parkfield, California, produced by the M7.4 Landers earthquake of June 28, 1992: Observations from the UPSAR dense seismograph array: Journal of Geophysical Research, v. 100, p. 675-690. 

U.S. Department of Commerce, National Oceanic and Atmospheric 

Administration. Climatological Data, California. Ashville, NC: 

National Climatic Data Center, 1970-1 992. 

Wesson, R.L., 1987, Modelling aftershock migration and afterslip of 

the San Juan Bautista, California, earthquake of October 3, 1972: 

Tectonophysics, v. 144, p. 2 15-229. 

1988, Dynamics of fault creep: Journal of Geophysical Research, v. 93, p. 8929-8951. 

Wesson, R.L., Jibson, R.W., Morton, D., Campbell, R.H., and 

Nicholson, C., 1986, Interpretation of surface cracks and other 

fault-line surface deformation associated with the North Palm 

Springs earthquake of July 8, 1986: Eos, Transactions, American 

Geophysical Union, v. 67, p. 1090. 

Williams, P., Fagerson, S., and Sieh, K., 1986, Triggered slip of the 

San Andreas fault after the July 8, 1986 North Palm Springs 

earthquake: Eos, Transactions, American Geophysical Union, v. 

67, p. 1090. 

Williams, P.L., McGill, S.F., Sieh, K.E., Allen, C.R., and Louie, J.N., 

1988, Triggered slip along the San Andreas fault after the 8 July 

1986 North Palm Springs earthquake: Bulletin of the Seismological Society of America, v. 78, p. 11 12-1 122. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

INCREASED SURFACE CREEP RATES ON THE SAN ANDREAS FAULT 

SOUTHEAST OF THE LOMA PRIETA MAIN SHOCK 

CONTENTS 

By Jeff Behr, Roger Bilham, and Paul Bodin, 

University of Colorado, Boulder; 

Kate Breckenridge, 

U.S. Geological Survey; and 

Arthur G. Sylvester 

University of California, Santa Barbara 

Abstract ............................................................... Introduction ............................................................... 

Postseismic surface slip above the southeastern Loma Prieta 

Rupture zone ....................................................... 

Aseismic surface slip on the San Andreas fault southeast of 

the Loma Prieta rupture .......................................... 

Elastic model estimates for aseismic slip in central 

California associated with the Loma Prieta earthquake ---- 

Creepmeter and borehole strainmeter data ......................... Conclusions ............................................................... 

Acknowledgments ....................................................... 

References cited ......................................................... 

Page 

Dl79 

179 

181 

181 

186 

187 

191 

191 

192 

ABSTRACT 

Surface creep rates increased substantially on the northwestern 50 km of the creeping section of the San Andreas 

fault following the Loma Prieta earthquake. Although 

postseismic creep southeast of the rupture zone initially 

resembled previously observed earthquake afterslip, creep 

rates between October 1990 and September 1992 have 

remained nearly constant at rates higher than pre-event 

values. Two-dimensional elastic boundary element models indicate that if the San Andreas and Calaveras faults 

were frictionless, vertical dislocations, then more than 20 

cm of surface slip would be induced on the 

northwesternmost 20 km of the creeping section in response to strain changes accompanying the Loma Prieta 

event. Induced additional postseismic surface slip nowhere 

exceeds 4 cm, indicating that much of the potential fault 

displacement in the northwestern creeping section has not 

occurred. Whereas postseismic displacements up to September 1992 have attained only 10 to 20 percent of modeled values, the spatial distribution of postseismic surface 

slip rate increases reflects the results of these simple models. In particular, potential slip and creep rates are predicted to be much reduced south of the San Andreas' 

intersection with the Calaveras fault. Changes in borehole 

strain and creep rates recorded along a 3.1-km-long section of the northwestern San Andreas creeping section 

provide strong indications of the depths of pre- and postevent shallow surface creep but satisfy a wide range of 

models in which the fault may be slipping at depth. The 

substantial difference in creep rate in that region suggests 

that observed strain changes may be locally enhanced by 

elastic inhomogeneity or local fault-slip processes. 

INTRODUCTION 

The Lorna Prieta earthquake was not associated with a 

throughgoing surface rupture even though numerous discontinuous surface cracks were manifest near the surface 

fault trace (U.S. Geological Survey Staff, 1990). It is possible that distributed shear could have absorbed surface 

displacements over a several kilometer wide zone, although 

geodetic data indicate that the Loma Prieta rupture did 

not reach the surface (Lisowski and others, 1990). In several recent strike-slip earthquakes in California, the amplitude of surface fault displacement has increased in the 

weeks to years following the main shock (Smith and Wyss, 

1968; Burford, 1972; Cohn and others, 1982; Bilham, 

1989; Sharp and others, 1989; Williams and Magistrale, 

1989). Anticipating the possible development of post-event 

surface slip, we installed three digital creepmeters (Bilham, 

1989) within 2 to 17 days after the Loma Prieta earthquake. Two of these, Madonna Road and Chittenden 

Bridge, were installed over the southeastern rupture zone, 

whereas the other, at Nyland Ranch, was installed near 

the northwestern end of the San Andreas fault creeping 

section (Behr and others, 1990) (fig. 1). 

Although the two instruments over the rupture zone 

have not recorded significant post-event tectonic slip, the 
Dl80 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Nyland Ranch creepmeter and several USGS creepmeters 

installed on the San Andreas fault creeping section 

southeast of the rupture zone (creepmeters XSJ2, XHR2, 

CWC3, XFL1 and XMR1) have recorded substantial 

changes in creep rate since the event. The form of this 

creep signal exhibits a decay in slip rate characteristic of 

postseismic slip observed following other earthquakes 

(Smith and Wyss, 1968; Sylvester, 1986; Marone and 

others, 1991). Although the term "afterslip" has hitherto 

been used to describe this characteristic surface slip behavior over the rupture zone, it conveniently describes the 

impulsive initiation of slip and subsequent decay in slip 

rate observed on adjacent fault segments. In the case of 

the Loma Prieta earthquake, however, the increased surface slip rates on the northwestern 50 km of the creeping 

section have not continued to decay toward pre-event 

levels, but have remained elevated and roughly constant 

during the "500 to 1,100 days following the main 

shock. 

The response of the creeping section of the San Andreas fault to previous earthquakes was examined by 

Mavko and others (1985), Burford (1988), and Simpson 

and others (1988). The northwestern creeping section, 

which terminates a few kilometers northwest of San Juan 

Bautista, defines an area of high surface creep rate on the 

San Andreas fault. The Loma Prieta earthquake epicenter 

was located approximately 50 km northwest of San Juan 

Bautista, and its aftershocks define a rupture zone that 

extends to within approximately 10 km of that town (fig. 

1). In this paper, we describe the nature of fault creep 

observed in the northwestern 50 km of the creeping section and estimate the additional fault slip induced by the 

Loma Prieta earthquake. Using two-dimensional boundary element methods, we examine the possible effect of 

Figure 1.-Location map showing University of Colorado (CU) digital creepmeters and U.S. Geological Survey (USGS) 

creepmeters used in this study in relation to the Loma Prieta main shock and aftershock zone (shaded), the San Andreas fault 

creeping section (hachured), and the principal San Francisco Bay area faults. The inset map shows the relationship between 

creepmeters at Nyland Ranch and XSJ2, and the borehole strainmeter, SJT, near San Juan Bautista; San Juan Bautista Mission 

is marked by a cross (t). 
INCREASED SURFACE CREEP RATES ON THE SAN ANDREAS FAULT SOUTHEAST OF THE MAIN SHOCK Dl81 

variable San Andreas and Calaveras fault geometries on 

the amplitude and spatial distribution of surface creep 

southeast of the Loma Prieta rupture. We further estimate 

the potential amplitude of slip induced by the Loma Prieta 

main shock on the San Andreas fault using simple frictionless models of the San Andreas and Calaveras faults. 

Finally, we attempt to reconcile changes between pre- and 

postseismic creep and strain rates near San Juan Bautista 

with changes in the depths of shallow and deep slip on 

the creeping fault . 

POSTSEISMIC SURFACE SLIP ABOVE 

THE SOUTHERN LOMA PRIETA 

RUPTURE ZONE 

The longest of the discontinuous surface fractures in 

the San Andreas fault zone associated with the Loma Prieta 

earthquake was at Madonna Road (Rymer, 1990). This 

crack, which displayed right-lateral offset of "2 cm, was 

instrumented 2 days after the event with a 10-m-long digital creepmeter. Even though right-lateral slip increased by 

8 mm on this fracture during the 10 months following the 

earthquake, slip since September 1990 has displayed significant apparent left-lateral displacement (fig. 2A). We 

interpret this apparent left-lateral signal as contraction 

during periods of heavy rain, followed by extension, or 

right-lateral motion, during dry spells. We suspect that 

much of the signal monitored by this instrument is related 

to hillslope instability. 

The Chittenden rail bridge crossing the Pajaro River at 

Pajaro Gap was instrumented in early November 1989 on 

the assumption that the structure crosses the 1906 fault 

break (Behr and others, 1990). Since mid-1990 the instrument has monitored no cumulative displacement other than 

an annual signal considered to be thermoelastic response 

of the bridge box-girder. If we consider that signal to be 

typical annual behavior, then during the first several 

months of instrument operation this creepmeter recorded 

a cumulative bridge contraction of approximately 3 cm 

(fig. 2B). Several mapped strands of the San Andreas fault 

do not pass beneath the bridge and, although it was pulled 

from its abutment during the 1906 earthquake, it is possible the Chittenden Bridge may not cross the main active 

fault trace (Prentice and Schwartz, 1991). 

ASEISMIC SURFACE SLIP ON THE SAN 

ANDREAS FAULT SOUTHEAST OF THE 

LOMA PRIETA RUPTURE 

A wooden fence at Nyland Ranch (1.4 km northwest of 

San Juan Bautista Mission), offset at a rate of 8 mmlyear 

between 1942 and 1978 (Burford and Harsh, 1980), represents the northwesternmost location of aseismic creep observed on the San Andreas fault. A 43-m-long alignment 

array was established adjacent to this wooden fence in 

1967 (Nason and Tocher, 1970). The periodic measurement of this array was supplemented by a biaxial rod 

extensometer with a dial gauge read-out installed 5 m 

southeast of the alignment array, and by a continuously 

operating creepmeter (USGS instrument SJN1) that was 

discontinued in 1985 due to repeated flooding of the instrument. Infrequent measurement data from the alignment array reveal an approximately linear slip rate of 7.0 

mdyear prior to 1989, although over short periods the 

sense of motion fluctuates in rate and sign (fig. 3A). Between 1968 and 1985 the creepmeter SJN1 recorded somewhat more rapid creep at 8.1k0.2 mmlyear, with rare creep 

events of amplitude approaching 1 mm (Schulz and others, 1982; Schulz, 1989). An ML=4.6 earthquake near the 

site in 1972 resulted in accelerated creep with a duration 

of several months and with a decaying rate typical of 

earthquake afterslip (Burford and others, 1973). 

Surface cracks were not present near the alignment array at Nyland Ranch a few hours after the Loma Prieta 

main shock, although they were evident on the fault 

through the town of San Juan Bautista to the southeast. 

Measurement of the alignment array a few days after the 

main shock revealed a small (=I mm) left-lateral signal. 

On October 22, 1989, we installed a digital creepmeter 

(Bilham, 1989) next to the east-west arm of the dial-gauge 

creepmeter to provide a continuous record of creep at this 

location. During the three years following the Loma Prieta 

earthquake, the creep rate at Nyland Ranch has averaged 

more than 20 mdyear (fig. 3). This increased slip rate 

did not begin until 44 days after the main shock; in the 

first 10 days of operation the creepmeter recorded less 

than 0.1 mm of left-lateral signal (possibly a response to 

instrument installation), subsequently replaced by slow 

right-lateral displacement of less than 10 microns per day. 

An abrupt increase in creep rate, which began 44 days 

after the main shock, exhibits the characteristic decay in 

rate described as earthquake afterslip. However, the creep 

rate has not continued to decay in the manner of previously described earthquake afterslip but instead has established an approximately linear 13.4-mdyear rate between 

480 and 1,120 days after the main shock (fig. 3B). If the 

mean pre-earthquake creep rate of 7 mdyear is subtracted 

from the mean post-earthquake rate of nearly 20 mm/ 

year, then an additional 3.9 cm of surface slip is inferred 

to have been induced by the Loma Prieta main shock in 

the three years since the event. 

Rainfall clearly modulates the short-term creep rate at 

Nyland Ranch. At times of rainfall, the single-component 

creepmeter displays significant shortening (1-5 mm). Following these periods, however, the instrument recovers 

and resumes recording dextral slip. These rainfall-induced 

transients have been removed from the record (fig. 3B) to 
Dl82 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

allow for more precise least-squares estimation of the over- creepmeter are significantly less contaminated by rainfall 

all creep rate. In an attempt to suppress signals related to signals. However, the continuous records from this instrusoil expansion induced by rainfall, the digital creepmeter ment and from the single creepmeter have been interwas supplemented by a differential creepmeter in March rupted by flooding of the instruments at times of high 

1990 (Bilham and Behr, 1992). Data from the differential water table, a problem responsible for discontinuing opCHITTENDEN BRIDGE 

NOV. 4, 1989 TO MARCH 23, 1992 

1/1/90 51 1/90 91 1/90 11119 1 51119 1 91119 1 1/1/92 

8 

DATE 

....... .... .... I I I:' I I ,I::.:' ' i..l.:z.l: ..::, I I ..... ......... 

......... ...... ....... - \ 

- A \ ....... \ ..... 

Figure 2.-A, Digital creepmeter record from Madonna Road, 20 Octo- positive. The high-amplitude diurnal and annual signal is due to thermal 

ber 1989 to 22 March 1992. Right-lateral displacement is positive. The expansion of the 110-m-long box girder to which the instrument rod is 

cumulative displacement is believed to result from ground failure of the attached. The vertical hachured lines in both A and B mark the April 

sediments in which the instrument is installed. B, Digital extensometer 1990 Loma Prieta aftershock sequence. Gaps in the data are caused by 

record from Chittenden Bridge, 4 November 1989 to 23 March 1992. instrument malfunction and are constrained by jaw separation readings 

Extension of the bridge and corresponding right-lateral displacement is of the instrument's digital caliper sensor. 

\ :: 

MADONNA ROAD 

6 - \ ,:' OCT. 20, 1989 TO MARCH 22,1992 - \ :!' \ .:" 

4 

2 

\ :? 

:t. 

.::: \ ........ :. ... ... .... .: \ ............ 

:. :-: ::::.-:: 

-:Ã£:: 

- - .... ... 

....... 

: :. r:- - -: . .... ::: \ 

. . ... 

.. :.: ... ' \ - ..... . ::,.::..:;- / 

4 

^. 

\ ........ .... ........ ....... + \ .......... . .......... .-. - ...... .. \ ......... . 8.28 ":. ... ? 

. . .- ... .... \ ... 

. . ... . . :: 

\ ...... 

..... ..... \ ..... ......... ..... ...... -...- ....... 

0- \ 

\ 

- .... .... .... \ .... - ..... ...... ....... 

\ 

\ 

..... .. ..... \ ... ... .... ... ........ ., " ..... -2+'Il1 ... llllltlt\llll It1l Ill1 'It1 11111111 lllllltlth....dtlllll* 

1/1/90 51 1/90 91 1/90 11119 1 51119 1 91119 1 1/1/92 

DATE 
INCREASED SURFACE CREEP RATES ON THE SAN ANDREAS FAULT SOUTHEAST OF THE MAIN SHOCK Dl83 

eration of USGS creepmeter SJN1 in 1985. The data seg- ered by measuring changes in the physical separation of 

ments in figure 3 represent a composite record constructed the jaws of the digital caliper-displacement sensors. 

from discontinuous data from both creepmeters. Surface We examine data from the five northwesternmost USGS 

displacements during gaps in the recorded data are recov- creepmeters on the San Andreas fault (fig. 1) from the 

200 NYLAND RANCH ALIGNMENT ARRAY 

1967- 1992 

150 

100 

50 

0 

1/1/68 1/1/72 1/1/76 1/1/80 1/1/84 1/1/88 1/1/92 

DATE 

- 

- NYLAND RANCH COMPOSITE 

50 - 

- 

- 

- 40 - 

- 

- 

- 

- 30 - 

- 

- 

- - 

100 200 300 400 500 600 700 800 900 1000 1100 

DAYS AFTER OCTOBER 18,1989 

Figure 3.-A, Record of alignment array offset (diamonds) at Nyland record. B, Nyland Ranch digital composite record, 4 November 1989 to 

Ranch from 1967 to 1992. The straight line represents the 7.0-mmlyr 5 November 1992. Transient rainfall-induced signals are omitted from 

average creep rate at Nyland Ranch from 1967 to 1990. A composite the record with all interruptions to the observed data constrained by 

record of post-Loma Prieta creep from the single- and double-rod digital manual caliper readings. The data after January 1991 are linearly fit by a 

creepmeters at Nyland Ranch has been appended to the alignment array creep rate of -13.4 mrntyr. 
D 184 AFTERSHOCKS AND POSTSEISMIC EmCTS 

time of the May 2, 1983, Coalinga earthquake until 2.8 mm to 6.8 mm of triggered dextral slip (Schulz, 1989; 

years after the Lorna Prieta earthquake (fig. 4). These table 1). During the 6 years prior to the Lorna Prieta main 

creepmeters recorded coseismic offsets ranging from -0.35 shock, creep rates at XSJ2, CWC3, XFLl and XMRl 

- A SAN ANDREAS FAULT CFEEP 

M 7.1 

1/1/84 1/1/85 1/1/86 1/1/87 1/1/88 1/1/89 1/1/90 1/1/91 1/1/92 

DATE 

240 - 

- 

- B DETRENDEDCMEP 

- 

- 

- 

- J 

- 

1/1/84 1/1/85 1/1/86 1/1/87 1/1/88 1/1/89 1/1/90 1/1/91 1/1/92 

DATE 

Figure 4.-A, USGS creepmeter records from the northwestern 45 km of (1983-1989) creep rate (see table 1). The detrended creep records show 

the creeping section of the San Andreas fault, 2 May 1983 to 1 Septem- significant changes in creep rate at about the time of the Loma Prieta 

ber 1992. The vertical line marks the Loma Prieta main shock. B, Records earthquake. 

from A, detrended by subtracting the least squares estimated pre-event 
INCREASED SURFACE CREEP RATES ON THE SAN ANDREAS FAULT SOUTHEAST OF THE MAIN SHOCK Dl85 

Table 1 .-Creep rates before and after the Loma Prieta earthquake determined from least squares analyses and end-point values 

[The pre-event data begin after the Coalinga earthquake (2 May 1983). Post-event data include a transient period marked by decreasing slip rate (cumulative afterslip 

column) and a period of approximately constant rate beginning in October 1990 (least squares and end-point fits, 1990-19921, The triggered slip estimates are from 

Schulz (1989). The last column shows estimates of the change in creep rate between the six years prior to Loma Prieta and the steady rate established after October 

1990. Uncertainties reflect the range of differences obtained through the least squares and end-point fit methods.] 

Creepmeter Distance Least End-point Triggered Cumulative Least End-point Rate change 

south of squares fit slip afterslip squares fit 1990-92 minus 

San Juan 1983-89 1983-89 18 Oct. '89 1989-90 1990-92 1990-92 1983-89 

(km)] (mrd~r)~ (md~r)~ (mm) (mm) (wyrl3 (m4yrI3 (md~ r) 

Nyland -1.4 7 .O 7 .O <2.0 3 5 13.4 12.9 +6.2&0.2 

XS J2 1.7 3.5 3.2 5.18 22 14.0 11.4 +9.5k1.7 

XHR 1 -2 12.8 8.1 7.9 4.25 40 15.7 17.2 +8.4k0.8 

CWC3 16.9 9.5 9.1 6.8 40 13.2 14.0 +4.3&0.6 

XFLl 30.3 6.6 6.4 -0.35 20 8.5 7.3 +1.4*0.7 

Xh4Rl 41.7 18.9 17.4 2.6 3 5 16.3 18.3 -0.8k1.8 

Distance along San Andreas fault trace south of Mission, San Juan Bautista. 

Least-squares and end-point fits to creep data from May 2, 1983 to October 17, 1989. Rates at Nyland Ranch are long-term 

averages established from alignment array data. 

Least-squares and end-point fits to creep data from October 1990 to September 1992. 

ranged from 4.0 to 0.75 mdyear lower than previously 

estimated long-term averages (Schulz and others, 1982). 

An apparent increase in creep rate of 0.9 mdyear at Harris Ranch (HR) about the time of the Coalinga earthquake 

may be a consequence of instrumental friction in 

creepmeter HRS3. The record from HRS3 was used by 

Schulz and others (1989) to estimate the creep rate from 

1970 to 1982, whereas the rate estimate between the 

Coalinga and Loma Prieta earthquakes was calculated from 

the records of XHRl, active from 1980 to 1984, and its 

replacement, XHR2, installed in 1985. Overall, Schulz 

(1989) noted that the long-term creep rate at Harris Ranch 

averages from 6 to 9 mdyr, depending on the time period considered. A further decrease in creep rate prior to 

the Lorna Prieta event, noted by Burford and Schulz (1988) 

and by Sylvester and others (1990), is evident in the creep 

records at XSJ2, CWC3 and XMRl (fig. 4, table 1). As at 

Nyland Ranch, surface slip at XHR2 and CWC3 appeared 

in a form similar to earthquake afterslip for approximately 

1 year after the earthquake. However, creep rates over the 

whole of the northwestern creeping section were significantly altered, with rates increased by a factor of 3.3 near 

the rupture zone (XSJ2), to a 10 percent reduction distant 

from the rupture (XMRl). Following the initial postseismic 

increases on the four northwesternmost USGS instruments, 

slip rates decreased until the fall of 1990. From then 

through September 1992, surface slip rates remained nearly 

constant along the northwestern 35 km of the creeping 

section at rates significantly increased above pre-event 

velocities (fig. 4, table 1). Because of seasonal variations 

in creep rate and the occurrence of episodic creep in some 

of the records, the mean rates in table 1 are sensitive to 

the methods by which they are estimated. As noted by 

Schulz and others (1982) the mean rate obtained by leastsquares analysis yields a different value from that obtained by taking the end-point values and dividing by the 

appropriate time interval. This is because of the "staircase" character of creepmeter records consisting of a cumulative signal formed by minor contributions from steady 

creep and substantial contributions from large episodic 

events. The resultant uncertainty in the actual change in 

creep rate initiated by the Loma Prieta earthquake is indicated by a range of values for each creepmeter (column 9, 

table 1). Because we have used different averaging windows, the mean, pre-Loma Prieta creep rates we present 

in table 1 differ from those reported by Schulz and others 

(1982) and Schulz (1989). 

We are unable to offer a satisfactory explanation for 

the 44-day delay between the Lorna Prieta main shock 

and the sudden emergence of rapid dextral creep at Nyland 

Ranch. If it were a propagating phenomenon, then 

creepmeters farther to the southeast should experience a 

sequential delay in their response to the main shock. Although there were similar 67- and 42-day delays in the 

onset of increased creep rates at XSJ2 and XFLl, respectively, a search for a ubiquitous southeastward (or 

northwestward) propagating process was unsuccessful. We 

do, however, exclude rainfall as the triggering process for 
Dl86 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

initiating creep, because heavy periods of rain occurred 

twice during the 44 days of creep quiescence with no 

significant perturbation to the creep record. 

ELASTIC MODEL ESTIMATES FOR 

ASEISMIC SLIP IN CENTRAL 

CALIFORNIA ASSOCIATED WITH THE 

LOMA PRIETA EARTHQUAKE 

Stress changes on faults in central California resulting 

from the Loma Prieta rupture (Reasenberg and Simpson, 

1992) cannot be used directly to estimate potential aseismic 

slip along the creeping section because slip locally modifies the calculated stress, which in turn leads to an incorrect estimate for potential slip. To overcome this problem, 

we use boundary element methods (Mavko, 1982) to estimate the slip at various locations along the northwestern 

San Andreas fault creeping section. Our approach (Bilham 

and King, 1989a) minimizes the stress in the medium surrounding the fault generated both by slip on the Loma 

Prieta rupture plane and by slip adjustment on segments 

of the Calaveras and San Andreas faults. That is, we ignore the steadily accumulating plate boundary loading rate 

responsible for the pre-event creep rates in central California. The creeping zone is simplified as a two-dimensional frictionless vertical cut in an elastic halfspace (fig. 

5) with the observed geometry of the San Andreas creeping section. We confine the model to these two faults, 

which further simplifies the analysis compared to that reported by Mavko (1982). We represent the Loma Prieta 

rupture as a vertical, rectangular dislocation with a length 

of 38 km on which is imposed a dextral slip of 1.66 m 

and a convergent displacement of 41 cm to simulate 1.2 

m of dip slip on a fault surface dipping 70' to the southwest (Lisowski and others, 1990). This simplified model 

for the Loma Prieta earthquake implies that rupture extends throughout a uniformly thick elastic crust which 

an Juan Bautista 

Figure 5.-Block diagram of the two-dimensional boundary element 

model used to estimate slip on the creeping segments of the Calaveras 

and San Andreas faults. Dextral slip with convergence is imposed on a 

vertical segment to emulate the Loma Prieta rupture northwest of the 

creeping section. In all models, the San Andreas and Calaveras faults 

are permitted to slip freely in response to strain changes calculated for 

the rupture. 

will consequently overestimate strain changes associated 

with a rupture confined to depths of 4 to 20 km. 

The geometry of the faults near San Juan Bautista have 

been interpreted differently in various maps of the region, 

and there is no consensus on the relationship between 

surface faults and active dislocations at depth (Mavko, 

1982). We use a creeping section represented by a series 

of contiguous, straight, vertical segments corresponding 

to line segments listed in Bilham and King (1989b). The 

displacements calculated for that model and for the current models are those that would result if aseismic andor 

seismic slip extended throughout the elastic crust and completely released the calculated fault driving stresses. Our 

models consequently provide an upper limit to the slip 

anticipated from the Loma Prieta rupture. 

In two boundary element models (fig. 61, slip on the 

creeping faults is driven only by the Loma Prieta rupture. 

In one model the Calaveras fault is locked, while in the 

other model the Calaveras fault slips freely and intersects 

the San Andreas fault. Whether or not slip occurs on the 

Calaveras fault, we find that the resulting characteristic 

slip distribution on the creeping section of the San Andreas fault reaches a maximum of 30 cm, 10 to 20 km 

southeast of San Juan Bautista. From that point southeastward, the potential slip with the Calaveras fault locked 

decays nearly linearly over the length of the creeping section. If the Calaveras fault is introduced as a freely slipping dislocation that nowhere approaches closer than 10 

km to the San Andreas fault, then maximum slip on the 

San Andreas is slightly reduced (25 cm), and the location 

of maximum slip is skewed 10 km to the northwest (model 

results not shown). If a slipping Calaveras fault contacts 

the San Andreas fault, then this abruptly reduces potential 

slip southeast of their intersection and significantly increases slip to the northwest. This result is counter-intuitive and contrasts with the finding reported by Mavko 

U, 1111111111111111 

30- - 

2 - - 

E 8 20- 

2 - 

U, 4 10.- u 

& - San Andreas-Calaveras 

E point of intersection 

E 50 100 150 

DISTANCE FROM END OF CREEPING SECTION, KILOMETERS 

Figure 6.-Slip induced on a freely slipping San Andreas fault by strain 

changes accompanying the Loma Prieta rupture. Model results are shown 

for a system with both a locked and a freely slipping Calaveras fault, 

which intersects the San Andreas. For the slipping Calaveras model, an 

abrupt decrease in potential slip occurs southeast of the intersection of 

the San Andreas and Calaveras faults. 
INCREASED SURFACE CREEP RATES ON THE SAN ANDREAS FAULT SOUTHEAST OF THE MAIN SHOCK Dl87 

(1982) that slip increases on the San Andreas fault southeast of its junction with the Calaveras fault. The reason 

for this difference is that our model is driven by stresses 

associated with the Loma Prieta rupture northwest of the 

Calaveras-San Andreas fault intersection, whereas slip in 

Mavko's (1982) model is driven by remote antisymmetric 

shear displacements. 

In models where the Calaveras fault is permitted to slip 

freely, we find that the mean slip on that fault is approximately 13 cm in a left-lateral sense, varying from a few 

centimeters to more than 20 cm at the latitude of the 

Loma Prieta rupture. Extending the fault model to include 

the southern 30 km of the Hayward fault permits an additional mean left-lateral slip of 5 cm on the Calaveras fault, 

but it increases the slip on the northwestern San Andreas 

fault by less than a few percent. Creepmeters and alignment arrays on the southern Calaveras recorded a reduction in right-lateral creep rates following the Loma Prieta 

main shock but no significant left-lateral slip (Galehouse, 

this chapter). If, prior to the earthquake, the Calaveras 

fault behaved as a frictionless surface and entirely released the applied dextral strain induced by plate motions, 

then we might anticipate that left-lateral slip induced by 

the main shock would be manifest on the fault. The absence of sinistral slip indicates that the dextral shear strain 

applied to the Calaveras fault prior to the earthquake exceeded the sinistral shear component applied to the fault 

by the earthquake. Creep retardation, a decrease in slip rate 

relative to long-term slip rates, presumably represents the 

expression of these combined strain fields. Our two-dimensional elastic models indicate that left-lateral slip 

would also be expected on the southern Hayward fault. 

As much as 1 cm of left-lateral slip has been manifest on 

that fault since the Loma Prieta earthquake (Lienkamper, 

oral commun,, 1992), indicating that dextral strain applied to the Hayward fault prior to the earthquake was 

less than the sinistral strain generated by the earthquake. 

Our models yield surprisingly large estimates for the 

amplitude of induced slip anticipated on the San Andreas 

fault, exceeding 20 cm in the northwestern 20 km of the 

creeping section. Using three-dimensional models in which 

coseismic strains fall off with distance much faster (-R-~) 

than they do in two-dimensional models (-R-~), 

Breckenridge and Simpson (this chapter) estimate maximum values of approximately 6 cm for induced rightlateral slip on the northwestern creeping section, with 

proportionately smaller slip induced on surrounding faults. 

Strain changes induced within the surrounding crustal volume by an earthquake evolve with time due either to viscous relaxation of the lower crust, or to time-dependent 

deep slip on faults within an otherwise elastic lithosphere 

(for example, Stein and others, 1992). The amplitude of 

slip on faults driven by such strains should evolve as well. 

Two-dimensional models of this system yield high estimates because in some sense they represent an extreme 

case in which all of the strain induced by the earthquake 

is manifest as slip. The introduction of a third dimension 

requires numerous other parameters to be estimated that 

we avoid in a two-dimensional approach; precise values 

computed in Breckenridge and Simpson's three-dimensional models are sensitive not only to which surrounding 

faults are assumed to slip, but also to the depth to which 

slip is assumed to occur. Nevertheless, our analysis permits the interpretation that less than 25 percent of the slip 

estimated by the two-dimensional models has occurred on 

the northwestern 20 km of the creeping San Andreas fault, 

and perhaps as little as 10 percent. At current creep rates, 

the slip modeled in two dimensions would not be manifest for decades, and if creep rates decrease over time, 

then potential slip will be delayed further. 

Frictional and viscous processes moderate the rate at 

which the estimated postseismic induced displacement will 

be manifest on the fault. However, since creep rates depend in part on the strain applied to the fault, we anticipate that observed postseismic slip rates will be 

proportional to the modeled potential slip amplitudes. In 

figure 7, observed surface creep rate changes in the northwestern 50 km of the creeping section are superposed on 

curves fit to two-dimensional model displacements calculated for three geometries of Calaveras fault/San Andreas 

fault interaction. Although the relative positions of the 

creepmeter data are exact, there is some uncertainty about 

the northwestern terminus of creep entered in the models, 

and in the location and proximity of approach of the Calaveras and San Andreas faults. For these reasons, we should 

not expect a precise correspondence between predicted 

slip data and slip-rate observations. Nevertheless, the predicted displacements and observed creep rate increases 

are reasonably close, suggesting our models may represent appropriate simplifications of the regional slip mechanism. Indeed, the fit of the creep data to the slip estimates 

in figure 7 suggest a best-fit model in which aseismic slip 

on the Calaveras fault may have occurred fewer than 3 

km from the creeping San Andreas fault. It would appear 

that more detailed models may provide a better estimate 

of the connectivity (Bilham and Bodin, 1992) between the 

San Andreas and Calaveras faults. Although our twodimensional methodology provides insight into the behavior of faults with approaches that are closer than the thickness of the elastic plate, in practice, three-dimensional 

subsurface geometry and analysis methods will more accurately characterize the detailed behavior of such systems. 

CREEPMETER AND BOREHOLE 

STRAINMETER DATA 

The mean creep rates at XSJ2 and Nyland Ranch for 

the year preceding the Loma Prieta main shock were 3.2 

to 3.5 mmlyear and 7 mdyear, respectively. From Octo- 
Dl88 AFTERSHOCKS AND POSTSEISMIC EmCTS 

ber 1990 to September 1992, the mean rates were approximately 12.7 mdyear and 13.1 mdyear, respectively. 

Thus, although recent post-event slip rates at these two 

creepmeters are similar, they differ before the main shock 

by a factor of two. Moreover, a change in the characteristics of slip at XSJ2 occurs at the time of Loma Prieta in 

that during nearly 3 years of observed postseismic displacement there have been as many as eight creep events 

(episodic creep), whereas only two creep events were recorded in the 3 years preceding the main shock and none 

during the immediately preceding 18 months. 

In view of the limited data available to study the event, 

models of the fault-slip process must inevitably remain 

non-unique. We use a simple two-dimensional analysis to 

test whether the strain and creep rates at the northwestern 

end of the fault are mutually consistent, A triaxial borehole strainmeter (SJT) operating 1.13 km from the fault, 

at 200 m depth, and within 3 km along strike from XSJ2 

and Nyland Ranch recorded a shear strain rate of 1.2kO. 1 

pstraidyear during the 10 months prior to the earthquake 

and 1.9kO.l pstraidyear from January 1990 through at 

least July 1991 (Gladwin and others, 1991; Gwyther and 

others, 1992). If we assume that observed strain rates (ad 

at) result from annual antisymmetric displacements applied remotely to the region and that they are monitored 

by a surface strainmeter y km from the fault (fig, 81, then 

we may estimate the remotely applied strain load 01p and 

the depths, b and H-a, between which the fault is locked 

(Tse and others, 1985): 

The surface slip, Au, on a fault subject to the same 

applied strain load is 

na nb 

OH 2+cos--cosAu=2-cosh-l[ nP 

cos - < +cos - $1 

(Tse and others, 1985).l 

Thus, the surface strain rate and surface creep rates 

provide certain limits to the applied strain loading rate 

and the locked depths of the fault. We use a mean postearthquake creep rate of 12.9 mdyear for XSJ2 (1 1.4-14 

mdyear) and Nyland Ranch (12.9-13.4 mmlyear) in the 

The cosine terms in this expression were incorrectly excluded in 

Bilham and Behr (1992) 

following estimates. For y = 1.13 km (neglecting the 200- 

m depth of the borehole strainmeter SJT), a surface creep 

rate of 12.9 mdyear is consistent with a fault locked 

below 1.57 km, subject to an applied shear strain rate of 

3.2 pstrainlyear. However, the observed strain and creep 

rates are compatible with a range of increased depths for 

subsurface slip (H-a) with correspondingly reduced loading rates as shown in figure 9. For elastic plate thicknesses (H) greater than 10 km, it is evident that the depth 

of surface slip is only marginally sensitive to slip at depth, 

varying by less than 10 percent for subsurface creep extending from the base of the elastic plate to within 3 km 

of the surface. The lowest applied shear-strain signal consistent with a creep rate of 12.9 mmlyear and a strain rate 

of 1.9 pstrairdyear observed 1.13 km from the fault is 0.7 

pstraidyear (fig. 9A). This value occurs for creep in a 15- 

km-thick elastic plate with a narrow region locked between 1.2 km and 2 km depth. Thus, from figure 9A, we 

deduce that post-earthquake shallow creep can occur to 

depths of 1.2 to 1.57 km for applied strain rates of 0.7 to 

3.2 pstraidyear, respectively, depending upon whether or 

not the fault creeps below 2 km depth. 

Assuming that the creep and strain are driven by remote loading of the fault, the difference in creep rates 

between Nyland Ranch (7.0k2.0 mmlyear) and XSJ2 

(3.4k2.0 mdyear) in the years prior to the main shock 

require different depths for slip at these two locations. 

The rate of shear strain parallel to the fault (Gamma 1 of 

Gladwin and others, 1991) recorded by SJT after mid1986 was near zero until one year before the main shock 

when strain rates increased to 1.2 pstraidyear. An observed strain rate of 1.2 pstrainlyear is consistent with the 

observed creep rate at Nyland Ranch for a depth of shallow creep of ~1.4 km and an applied strain loading rate of 

1.9 pstraidyear (see fig. 9B). Again a suite of possible 

combinations of creep at depth and applied strain rate can 

satisfy the observed creep and borehole strain data, one of 

which requires a low strain loading rate (0.41 pstrainl 

year) for a model in which a narrow zone remains locked 

between 1.1 and 2 km (fig. 9B). During this time, however, the lower XSJ2 creep rate required a significantly 

shallower depth of creep compared to that inferred for 

Nyland Ranch (fig. 90. The range of acceptable applied 

strain rates is not changed much given this reduced creep 

rate, but the range of shallow creep depths is reduced to 

330 to 500 m, depending on the applied loading rate (fig. 

9C). Although other interpretations are possible, a plausible explanation is that the depth of shallow creep increased by approximately a factor of three at XSJ2 at the 

time of the Loma Prieta earthquake (from 400 m to ~1.5 

km), but remained substantially unchanged at Nyland 

Ranch (from ~1.5 km to = 1.4 km depth). 

These inferred changes for depth of shallow creep are 

reasonably well indicated by the observed combination of 
INCREASED SURFACE CREEP RATES ON THE SAN ANDREAS FAULT SOUTHEAST OF THE MAIN SHOCK Dl89 

IV) 30 

s 

w 

i- 20 G 

u 

CL,- 

1-1 IV) ^ 10 

P 

I 

2 

...... Calaveras at 3 km 

............................................. 

DISTANCE FROM NORTHERN END OF CREEPING SECTION, KILOMETERS 

Figure 7.-Estimates of potential slip and observed creep rate changes Andreas to within 3 km, Calaveras intersecting San Andreas at kilomeon the northwestern 50 km of the San Andreas fault. Potential slip is ter 27.5. Estimates of the change in creep rate following the Loma Prieta 

estimated using three boundary element models involving different Ca- earthquake (column 9 of table 1) are shown as vertical bars. 

laveras fault geometries: Calaveras locked, Calaveras approaching San 

shear 

from 

\ remote 1 

.'] shear 

strainmeter 

antisymmetric 

displacement 

Figure 8.-Block diagram representing an elastic plate of thickness H undergoing antisymmetric shear 

displacement with a locked patch between depths b and H-a, after Tse and others (1985). This model is 

used in the analysis of surface creep and strain rates observed near San Juan Bautista. 
Dl90 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

ciple, by estimating the elastic strain changes induced by 

the Loma Prieta event at Nyland and XSJ2, it is possible 

to determine the increment in loading rate and thereby 

limit the possible range in creep depth on the subsurface 

fault. We do not pursue this here because the starting 

geometry is determined by the background loading rate, 

which is insufficiently established. Certainly additional 

strain measurements would permit a substantial improvement in our ability to interpret the ranges of fault slip at 

depth. 

An important feature in the XSJ2 data is the absence of 

episodic creep events in the data in the 18 months prior to 

the main shock. Episodic creep is defined as an abrupt 

increase in slip velocity with a duration of a few hours to 

a few days, commonly manifest as a transient acceleration 

followed by a gradual decay in slip rate. In previous years, 

the slip rate at XSJ2 was formed from a combination of 

episodic creep and background creep, a slow, continuous 

slip of the surface fault. Bilham and Behr (1992) proposed that steady background creep and episodic creep 

may be generated at different depths on a creeping fault, 

and that the ratio of the rates of background to episodic 

creep is approximately proportional to the ratio of the 

depths to which these processes occur. The ratio of cumulative background creep to cumulative episodic creep recorded by XSJ2 is 0.29 for the period 1983-1989, and 

0.37 after the Loma Prieta main shock, yielding a mean 

value of 0.33. That is, the long-term contribution to creep 

from episodic events is three times greater than the contribution from background creep. This implies that the 

depth of episodic creep is three times greater than the 

depth of background creep. If one of these depths is known, 

it is possible to estimate the other. From the preceding 

estimate of depth of shallow creep at XSJ2, the lack of 

episodic creep in the 18 months prior to the main shock 

suggests a maximum depth of 400 m for background 

creep. Following the main shock, episodic creep resumed 

and we inferred a creeping depth of "1.5 km. Thus, the 

ratio of depths derived from the long-term ratio of episodic and background creep agrees with the ratio of depths 

determined from reconciling the strain and creep rates 

before and after the main shock. We conclude that the 

creep retardation near San Juan Bautista in the late 1980's, 

as noted by Sylvester and others (1990) (fig. 4), was caused 

by the cessation of episodic creep in the depth range 0.5 

to 1.5 km. We interpret the return of episodic creep at 

XSJ2 after the Loma Prieta main shock as the resumption 

of surface creep to a depth of 1.5 km. Several factors 

affecting the fault zone could be responsible for arresting 

episodic creep prior to the main shock, including an inAPPLIED STRAIN, MICROSTRAIN PER YEAR - 

0 1 2 3 0 ,l,,,,~f,#r~ll 

surface creep (b) -.-.- .- 

-12 I &=1.9 pstrain/year 

- 15 L creep 1 2.9 mmlyear 

- 

llllllllll 

surface creep (b) 

e= 1.2 pstrainlyear 

creep 7.0 mrnlyear 

(A) RECENT NYLAND AND XSJ2 (B) NYLAND PRE-MAINSHOCK (C) XS J2 PRE-MAINSHOCK 

Figure 9.-Analytic solutions for applied strain and depths of surface (b) In 9A, a shear strain loading rate of 1.32 pstrainlyear generates creep of 

and subsurface (H-a) creep that satisfy observed creepmeter and bore- 12.9 mmlyear with strain of 1.9 pstrainlyear observed 1.13 km from the 

hole strain rates (indicated). The three curves in each graph correspond fault. The same load rate generates similar creep and strain rates in a 10- 

to the depth of surface creep (b) and the upper depths of a subsurface km-thick elastic plate with a narrower locked region. Examples B and C 

creeping zone (H-a) for crustal thickness (H) of 10 and 15 km. The correspond to the pre-Loma Prieta creep rates observed at Nyland Ranch 

vertical bar in 9A corresponds to a fault creeping above and below a and XSJ2, respectively, and an observed strain rate of 1.9 pstrainlyr. 

region locked between 1.49 km and 4 km in a 15-km-thick elastic plate. 
INCREASED SURFACE CREEP RATES ON THE SAN ANDREAS FAULT SOUTHEAST OF THE MAIN SHOCK Dl91 

crease in fault normal stress, a decrease in pore pressure, 

or other physical changes in the rheology of the fault zone. 

Despite the persuasive consistency of the above results 

and the potential existence of bimodal slip at XSJ2, several alternative explanations for the observed strain and 

creep data may be considered. Alignment array data recorded between 1968 and 1977 at Mission Vineyard Road, 

site of the XSJ2 creepmeter, indicated a diffuse shear zone 

(Burford and Harsh, 1980) in which the primary slip zone 

spanned by XSJ2 (13-m fault normal aperture) exhibited 

a creep rate of 7.6 mmlyr, while the alignment array (79- 

m aperture) recorded an average of 13.4 mrnlyr of shear 

deformation. It is therefore possible that the observed preevent slip deficit at XSJ2 results from a lack of coverage 

of the entire deformation zone. Whether or not this is the 

case, XSJ2 has been recording all slip within k 40 m of 

the primary fault trace since Loma Prieta. This seems 

especially to be the case during a period of significantly 

increased surface slip from December 1992 through June 

1993, when records from XSJ2 and the Mission Vineyard 

Road alignment array indicate 16.8 mm and 16.9 mm of 

right-lateral shear, respectively (Galehouse, oral commun., 

1993). The suggestion that XSJ2 was not monitoring the 

entire deformation zone prior to Loma Prieta, but has 

been since, requires a change over time in the location, 

or the nature, of principal shear deformation at that latitude. 

Additional alternatives may be assessed by examining 

transient strain and creep data at the time of episodic 

creep events to determine the regions where such events 

are generated. The very low strain rates recorded by the 

borehole strainmeter prior to 1988 require a different depth 

of slip to be applicable at that time. Strain rates recorded 

by SJT are lower than the 0.3 pstrainlyear measured geodetically by the USGS in the region northwest of San 

Juan Bautista (Lisowski and Prescott, 1981; Prescott and 

others, 1981). If we assume that the applied strain rate 

from 1986 to 1988 was 0.3 pstraidyear, then an estimated maximum observed borehole strain rate of 0.1 

pstraidyear requires the fault at Nyland Ranch to slip 

above 4.8 krn and below 5.1 km, whereas under the same 

conditions, the slip rate of 3.3 mmlyear at XSJ2 implies 

that the fault is locked below 4.8 km. These large depths 

for surface creep appear to us implausible in the extreme 

northwestern end of the creeping section because the surface fault is completely locked less than 3 km to the northwest of Nyland Ranch. We conclude that the borehole 

strain rates reported by Gwyther and others (1992) may 

signify changes in the fault zone which are more complex 

than those modeled here. In the 6 months prior to the 

Loma Prieta event, slip was apparently confined to shallow depths. The transient responses of the creepmeters in 

the months following the main shock confirm that the 

fault is responding differently along strike to strain changes 

induced by the Loma Prieta earthquake. 

CONCLUSIONS 

Aseismic slip rates on the northwestern part of the San 

Andreas fault creeping section in central California increased substantially following the Loma Prieta earthquake 

and continue, three years after the main shock, to be above 

their mean pre-earthquake values. Farther southeast in central California, creep rates are substantially unaltered by 

the earthquake. Our two-dimensional elastic models indicate that maximum displacements of 20 cm on the northwestern creeping section will result from strain changes 

accompanying the Loma Prieta main shock. Although these 

models differ from the creeping San Andreas and Calaveras fault systems in that they are simulated as frictionless, vertical planes, they suggest that southeast of the 

San Andreas fault's intersection with the Calaveras fault, 

displacements induced on the creeping section should be 

significantly reduced in comparison to those northwest of 

the intersection. This finding qualitatively agrees with observed creep-rate increases. At current surface creep rates, 

the induced fault displacements predicted by two-dimensional models of the Loma Prieta earthquake will take a 

minimum of two decades to eliminate by creep processes 

alone, although three-dimensional models indicate that 

most, if not all, of the induced slip will occur within the 

next few years. Another possibility is that slip predicted 

for the northwestern creeping section, which has not yet 

occurred aseismically, may occur during future seismic 

activity. 

Our two-dimensional analysis of strain and creep data 

suggests that slip was apparently confined to shallow 

depths beneath creepmeter XSJ2 during the 6 months prior 

to the Loma Prieta earthquake. High strain rates reported 

near the northwestern end of the creeping section by 

Gwyther and others (1992) can be reconciled with high 

creep rates by permitting a substantial fraction of the San 

Andreas fault to slip below a locked zone from 1 to 2 km 

depth. However, we consider that the substantial aseismic 

slip at depth that is implied by these models is inconsistent with the setting at the northwestern end of the creeping section in that surface creep is not observed more than 

a few kilometers northwest of Nyland Ranch. We conclude that the observed high strain rates recorded near 

San Juan Bautista must be related to a local strain adjustment process, perhaps caused by the concentration of strain 

near the creeping fault as it adjusts to the new strain conditions resulting from the Loma Prieta main shock. 

ACKNOWLEDGMENTS 

This work was funded by the U.S. Geological Survey 

grant 14-08-001-G1876. We wish to thank three anonymous reviewers for their perceptive, insightful, and help- 
Dl92 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

ful comments. We thank Mr. Frank Avilla and his family 

for the use of the pasture and barn at Nyland Ranch. 

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Bilham, R., and Behr, J., 1992, A two-layer model for aseismic slip on 

the Superstition Hills fault, California: Bulletin of the Seismological Society of America, v. 82, no. 3, p. 1223-1235. 

Bilham, R., and Bodin, P., 1992, Fault zone connectivity: slip rates on 

faults in the San Francisco Bay area, California: Science, v. 258, 

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Burford, R.O., 1972, Continued slip on the Coyote Creek fault after 

the Borrego Mountain earthquake: U.S. Geological Survey Professional Paper 787, p. 105-1 1 1. 

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Burford, R.O., Allen, S.S., Lamson, R.J., and Goodreau, D.D., 1973, 

Accelerated fault creep along the central San Andreas fault after 

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Burford, R.O., and Harsh, P.W., 1980, Slip on the San Andreas fault in 

central California from alignment array surveys: Bulletin of the 

Seismological Society of America, v. 70, no. 4, p. 1233-1261. 

Burford, R.O. and Schulz, S.S., 1988, A current retardation rate of 

aseismic slip on the San Andreas fault at San Juan Bautista, California: E 0 S, Transactions of the American Geophysical Union, 

v. 69, no. 44, p. 1424. 

Cohn, S.N., Allen, C.R., Gilman, R., and Goulty, N.R., 1982, 

Preearthquake and postearthquake creep on the Imperial fault and 

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Mavko, G.M., Schulz, S., and Brown, B.D., 1985, Effects of the 1983 

Coalinga, California, earthquake on creep along the San Andreas 

fault: Bulletin of the Seismological Society of America, v. 75, no. 

2, p. 475-489. 

Nason, R.D., and Tocher, D., 1970, Measurement of movement on the 

San Andreas fault, in Earthquake Displacement Fields and Rotation of the Earth, L. Mansinha and others, eds., Riedal, Dordrecht, 

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Prentice, C.S., and Schwartz, D.P., 1991, Re-evaluation of 1906 surface faulting, geomorphic expression, and seismic hazard along 

the San Andreas fault in the southern Santa Cruz Mountains: Bulletin of the Seismological Society of America, v. 81, no. 5, p. 

1424-1479. 

Prescott, W.H., Lisowski, M., and Savage, J.C., 1981, Geodetic measurement of crustal deformation across the San Andreas, Hayward, 

and Calaveras faults near San Francisco, California: Journal of 

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fault subsequent to the 1966 Parkfield earthquake: Bulletin of the 

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THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKE OCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP 

RATES IN THE SAN FRANCISCO BAY REGION 

CONTENTS 

By Jon S. Galehouse, 

San Francisco State University 

Abstract ............................................................... Introduction ............................................................... San Andreas fault ........................................................ 

Hayward fault ............................................................ 

Calaveras fault ........................................................... Concord fault ............................................................. Green Valley fault ....................................................... Other faults ............................................................... Summary and conclusions ............................................. 

Acknowledgments ....................................................... 

References ............................................................... 

Page 

Dl93 

193 

195 

195 

199 

202 

204 

204 

204 

206 

206 

ABSTRACT 

We have been measuring creep on San Francisco Bay 

region faults since 1979. Virtually no creep preceded the 

Loma Prieta earthquake (LPEQ) on the San Andreas fault 

northwest of the aftershock zone and virtually none has 

occurred since. In contrast, the post-LPEQ creep rate of 

13.5 mdyr near the northwestern end of the creeping 

segment of the San Andreas fault near San Juan Bautista 

is about twice the pre-LPEQ average. About 5 mm of 

coseismic right slip was triggered here by the LPEQ. Most 

of the Hayward fault creeps at a long-term rate of about 5 

mdyr. The median rate at five sites along the Hayward 

fault that we had been measuring for the decade before 

the LPEQ was 4.9 mdyr. The median rate at these same 

sites for the 3.4 yr following the LPEQ is 3.6 mdyr. The 

rate has been particularly low at a site near the northwestern end of the fault in San Pablo and at sites near the 

southeastern end in Fremont. In fact, left-lateral slip of a 

few centimeters associated with the LPEQ may have occurred in southern Fremont along this previously rapid, 

right-lateral-creeping segment of the fault. Small amounts 

of left-lateral slip have continued there since the LPEQ 

(through March 1993). The average creep rate on the 

southern Calaveras fault in the Hollister area before the 

LPEQ was 6.4 mdyr at one site and 12.2 mmlyr at another. More than 10 mm of right-lateral slip was triggered 

here by the M=6.3 Morgan Hill earthquake (MHEQ) in 

1984 and again in 1989 by the M=7.1 LPEQ. Since the 

LPEQ-triggered slip, one site has slowed to 4.2 mdyr 

and the other site has stopped creeping (the one that was 

creeping at 12.2 mmlyr), which has resulted in a slip deficit of more than 2 cm. The Concord fault creeps at an 

overall long-term average rate of about 3 mmlyr and the 

Green Valley fault at about 5 mmlyr. Both move episodically and both have post-LPEQ rates of about 1-2 mdyr 

slower than the longer-term averages. No creep has occurred at a site on the northern Calaveras fault in San 

Ramon either before or after the LPEQ. The same is true 

for sites on the noncreeping Seal Cove-San Gregorio, 

Rodgers Creek, West Napa, and Antioch faults. The 

changes in creep rates on San Francisco Bay area faults 

after the LPEQ are consistent with the static shear stress 

changes estimated for the LPEQ. 

INTRODUCTION 

My student research assistants and I have been measuring creep (aseismic slip) on active faults throughout the 

greater San Francisco Bay region since 1979. Over the 

past 13.5 yr (September 1979 to March 1993) we have 

made about 1,325 creep measurements at about two dozen 

sites (fig. 1). About 850 measurements were made in the 

decade preceding the Loma Prieta earthquake (LPEQ) and 

about 475 were made in the 3.4 yr following the quake. 

This data set enables us to determine the detailed creep 

rates and characteristics at each site and to compare measurements made after the LPEQ with those made previously in order to determine the effects, if any, of the 

LPEQ. An earlier summary of the effects of the LPEQ on 

creep rates during the first 6 months after the quake is in 

Galehouse (1990). 

We use a theodolite triangulation method (Galehouse 

and others, 1982), which allows us to determine the amount 

of strike-slip surface creep by noting changes in angles 

between sets of measurements taken across a fault at 
Dl94 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

different times. Most of the sites span a fault width of 50- 

225 m, but a few must span a greater width because of 

site considerations. The precision of the measurement 

method is such that we can detect with confidence any 

movement more than 1-2 mm between successive measurement days. 

The M=7.1 LPEQ is the largest earthquake to have occurred in the San Francisco Bay region since 1906 (Plafker 

and Galloway, 1989). Before 17 October 1989, we had 

been remeasuring sites about once every 2-3 months. Because of augmented funding following the LPEQ, we were 

able to increase the measurement frequency at sites on the 

San Andreas, Hayward, and southern Calaveras faults to 

about once every 5-6 weeks through 1990. Since then we 

are again occupying all sites about once every 2-3 months. 

Results are presented in figures 2-5, which give the 

average creep rate at each site as determined by the slope 

of the least-squares line. These figures also indicate the 

fault width spanned (W) and the time of the LPEQ (shown 

as a vertical line). All creep rates presented in this paper 

should be considered minimum rates because undetected 

creep on additional active fault traces could be occurring 

outside of the fault width spanned. Most sites are located 

in low-relief areas, so creep due to mass movement is 

probably not significant. Even though our data include 

some of the area's wettest and driest years, the long-term 

Figure 1 .-Map of San Francisco State University creep measurement sites (numbered dots). Epicenters and 

magnitudes are indicated for the 6 August 1979 Coyote Lake earthquake (CLEQ), the 24 April 1984 Morgan 

Hill earthquake (MHEQ), and the 17 October 1989 Loma Prieta earthquake (LPEQ). 
EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP RATES IN THE SAN FRANCISCO BAY REGION Dl95 

trends in creep rates do not seem to be significantly affected by these surficial weather conditions. We think that 

at most sites, the measured creep rate closely approximates the right-lateral, strike-slip component of the tectonic creep rate. All creep and slip mentioned in this paper 

are strike slip with right-lateral slip defined as positive. 

SAN ANDREAS FAULT 

We presently measure six sites on the San Andreas fault, 

five of which are shown in figure 1. Measurements for the 

past 8.0 yr at Site 14 at the Point Reyes National Seashore 

Headquarters and for the past 12.9 yr at Site 10 in South 

San Francisco suggest that the segment of the San Andreas fault is locked between these two sites that are about 

135 km and 80 km northwest of the LPEQ epicenter. 

Virtually no creep was occurring prior to 17 October 1989 

and virtually none has occurred thus far in the 3.4 yr 

since the quake (fig. 2). The same is true for our most 

northerly site (18) on the San Andreas fault in the Point 

Arena area (not shown in fig. 1). 

Langbein (1990) detected a few millimeters of 

postseismic slip within a small, multi-kilometer-long geodetic network spanning the San Andreas fault at the northwestern end of the LPEQ aftershock zone in the 7-200 

day interval following the quake. In November 1989 in 

order to detect any post-LPEQ creep closer to the epicentral area, we began measuring a previously established 

U.S. Geological Survey (USGS) site in Woodside (Site 

22) slightly northwest of the aftershock zone. Our results 

for the past 3.3 yr, when compared to unpublished USGS 

measurements in 1977 and 1980 (R. Burford, oral 

commun., 1989), show that little creep occurred at this 

site before, during, or after the LPEQ (through 13 February 1993). The rate that we measure is el mmlyr (fig. 2). 

We also established a new site (23) on the San Andreas 

fault near the southeastern end of the LPEQ aftershock 

zone just northwest of San Juan Bautista. Virtually no 

post-LPEQ creep has occurred at this site either (through 

20 February 1993). 

In July 1990, we established Site 25 on the creeping 

segment of the San Andreas fault just southeast of San 

Juan Bautista and the LPEQ aftershock zone. This site 

spans the location of USGS creepmeter XSJ2, which 

showed a slip of 5.2 mm triggered by the LPEQ shaking 

(Schulz, 1989). Our measurements give a creep rate of 

13.5k1.0 mdyr for the past 2.6 yr (through 20 February 

1993). This rate is about the same as the creepmeterdetermined rate of 12.4 mdyr following the LPEQ and 

considerably faster than the pre-LPEQ longer-term 

creepmeter rate of about 7 mmlyr (Schulz, 1989; 

Breckenridge and Burford, 1990; Gladwin and others, 

199 1 ; Breckenridge and Simpson, 1992). Bilham (1992) 

reports that the creepmeter-determined rate at Nyland 

Ranch, between our Site 23 and Site 25, was 20 mdyr in 

the 2 years following the LPEQ which is approximately 3 

times the mean creepmeter rate at this site from 1968-85 

(Schulz, 1989). 

In summary, the San Andreas fault at five sites (18, 14, 

10, 22, 23) along the locked portion of the fault both 

northwest and southeast of the LPEQ aftershock zone does 

not appear to have been affected by the LPEQ. Virtually 

no surface creep has been detected thus far. In contrast, 

the northwestern portion of the creeping segment of the 

fault at Site 25 near San Juan Bautista had about 5 mm of 

slip triggered by the LPEQ and has continued creeping at 

a post-LPEQ rate that is about twice the longer-term, preLPEQ creepmeter average. 

HAYWARD FAULT 

We have been measuring creep at five sites along the 

Hayward fault for about 13 years and have determined 

that the average long-term rate is slightly less than 5 mml 

yr (fig. 3). Although creep characteristics (steady or episodic) differ from site to site, the average rates for all 

sites are similar. Table 1 gives the relationship between 

creep on the Hayward fault and the LPEQ and is a summary of much of the following discussion. The Hayward 

fault at Site 17 at Contra Costa College in San Pablo was 

creeping at a rate of 4.7 mdyr for the 9 years before the 

LPEQ. Since the LPEQ, however, the rate has been only 

2.2 mmlyr, which has brought the overall average down 

to 4.2 mdyr for the past 12.6 yr. 

The Hayward fault at Site 13 at Rose Street in Hayward 

has also moved episodically at an overall average rate of 

4.9 mdyr since June 1980; the same as the average in the 

9.3 yr before the LPEQ. With J. Lienkaemper of the USGS, 

we remeasured curb offsets and old city of Hayward arrays at Rose Street in late 1992 and determined that the 

overall creep rate there since 1930 is 5 mmlyr. A leastsquares line through our theodolite-measured data points 

collected after the LPEQ, however, indicates a rate of 6.8 

mdyr. This higher rate may be partly due to a statistical 

phenomena because of the episodic nature of creep events 

at this site. The overall rate of 4.9 mdyr that includes all 

the theodolite data is the same as the pre-LPEQ rate. The 

total displacement since the LPEQ, however, is 20.1 mm 

in 3.37 yr, which gives an average of 6.0 mndyr, closer to 

the long-term average rate but still about 1 mmlyr higher. 

The Hayward fault at Site 12 at D Street in Hayward has 

two active traces and moves fairly steadily and uniformly, 

unlike the episodic creep at Rose Street just 1.3 km to the 

northwest. At D Street the fault was creeping at a rate of 

4.9 mdyr for the 9.3 yr before the LPEQ. Since the LPEQ, 

however, the rate has dropped to 3.6 mdyr, which has 
Dl96 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

SF-18 POINT ARENA AREA (Alder Creek) 

20 

LPEQ 

1.0 Â 0.1 mdyr for 12.0 yrs 

W = 267.4 meters n - LJ 3 - 0 

I I I I I I I I I I I I I -1 0 

W = 70.6 meters I02 w 

-1 0 s w 

20 g 

I 20 

SF-10 SOUTH SAN FRANCISCO (Duhallow Way) 

SF-1 4 POINT REYES NATIONAL SEASHORE HEADQUARTERS 

0.6 Â 0.1 mdyr for 8.0 yrs LPEQ 

1 -0.1 Â 0.0 mdyr for 12.9 yrs LPEQ 1- 10 

â -10 w 

W = 207.5 meters 

0 

1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 

20 

SF-22 WOODSIDE (Roberta Drive) 

0.9 Â 0.3 mmlyr for 3.3 yrs 10 

W = 91.2 meters 

Virtually no movement between 16 Feb 77 and 4 Nov 89 0 

20 fc 

SF-23 CANNON ROAD (San Juan Bautista area) _1 

-0.2 Â 0.3 mmlyr for 3.3 yrs 

W = 88.0 meters 

SF-25 MISSION VINEYARD ROAD (San Juan Bautista area) 

1 3.5 + 1.0 mdyr for 2.6 yrs Note: Different vertical 

W = 134.2 meters scale on this graph. 

- 10 

I I I I I I I I I 

Figure 2.-San Andreas fault displacement measurements at six sites square. A straight line represents the least-squares fit to all the pre- and 

between 1980 and 1993. LPEQ = Loma Prieta earthquake of 17 October post-LPEQ measurements at each site and its slope indicates the overall 

1989. Squares represent individual measurements at each site. The *l average creep rate. W is the width of the fault zone spanned by the 

standard deviation for each measurement almost always falls within its theodolite measurements. 
EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP RATES IN THE SAN FRANCISCO BAY REGION 

80 

60 

4.2 Â 0.1 mmlyr for 12.6 yrs 

W = 106.8 meters 40 

20 

0 

80 

4.9 Â 0.1 mdyr for 12.7 yrs 60 

W = 153.9 meters 40 

20 

0 

80 

60 

4.6 Â 0.1 mdyr for 12.7 yrs 40 W = 136.2 meters 

20 

0 

I I I I I I I I I 1 I I I I I : -20 

1979 1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 

80 SF-24 FREMONT (Camellia Drive) 

- 60 -0.9 Â 0.2 mdyr for 3.0 yrs 

W = 88.6 meters LPEQ - 40 F 

Figure 3.-Hayward fault displacement measurements at six sites between 1979 and 1993. See fig. 2 for further explanation, 
Dl98 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 1.-Hayward fault creep rates measured by theodolite between 1979 and 1993 

[See fig. 1 for site locations.] 

All 

Pre-Loma Prieta Post-Loma Prieta Measurements 

Measurement Average Measurement Average Total Average 

First Lasta Rate Firsta Last Rate Years Rate 

Site (YO (YO (mm/~r) (YO (YO (mm/~r) (mm/~r) 

17 1980.609 1 989.597 4.7k0.3 1 989.866 1 993.200 2.a0.4 12.6 4.2k0.1 

San Pablo 

Hayward 

Rose St. 

12 1980.478 

Hayward 

D st. 

2 1979.729 

Union City 

1 1979.726 1989.595 5.420.1 1989.808 1993.178 1 .1*0.5 13.5 4.W0.1 

Frem on t 

Rockett Dr. 

24 1967 1987 9.S0.3 b 1990.1 15 1993.123 -0.9k0.2 3.0 -0.9k0.2 

Fremont 

Camellia Dr. 

27 1968 1982 8.2k0.2 1992.262 1993.123 -4.Oa.8 0.9 -4.Ok3.8 

Fremont 

Parkmeadow Dr. 

aThe Loma Prieta earthquake occurred at 1989.795. 

b Pre-Loma Prieta rate from Lienkaemper and others (1991) 

from curb offsets for the 20 years from 1967-1 987. 

Pre-Loma Prieta rate from Harsh and Burford (1 982) from 

alignment array data from 1968-1 982 at a site 0.3 km 

southeast of Parkmeadow Drive in Fremont. 

brought the overall average down to 4.6 mmlyr for the The Hayward fault at Site 2 at Appian Way in Union 

past 12.7 yr. It is interesting to note but difficult to ex- City also moves fairly steadily, though not as uniformly 

plain why the Rose Street and D Street sites that are so as at D Street in Hayward. The LPEQ appears to have had 

close together and had the same pre-LPEQ creep rate little effect on the fault at Site 2. Even though the postshould be so different in the 3.4 yr following the quake, LPEQ rate is 0.3k0.4 mmlyr higher than the average bewith Rose Street apparently speeding up and D Street fore the LPEQ, the overall average rate has remained at 

slowing down. 4.7 mmlyr since September 1979. The Hayward fault at 
EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP RATES IN THE SAN FRANCISCO BAY REGION Dl99 

Site 1 at Rockett Drive in Fremont moves episodically 

and was creeping at a rate of 5.4 mdyr for the decade 

prior to the LPEQ. With J. Lienkaemper of the USGS, we 

measured the amount of curb offset along Rockett Drive 

in January 1993 and determined that the overall creep rate 

since 1964 is 5.3 mmlyr. This projects to a rate of about 

5.8 mdyr for the 25 years prior to the LPEQ, because the 

extremely low creep rate at this site since the LPEQ has 

brought the overall average down. The creep rate at Rockett 

Drive has been only 1.1 mmlyr since the LPEQ, which 

has reduced the average rate for the past 13.5 yr to 4.6 

mdyr. 

Although the creep data are equivocal because of their 

scatter and the normal episodic nature of movement at 

most sites, there is a possibility that a slowdown in creep 

on the Hayward fault began before the LPEQ. A crude 

estimate of the time when a slowdown may have begun 

can be made by noting the time of the last previous (before the LPEQ) significant creep event at each site. Even 

though the results are highly subjective and other investigators could pick different times for the onset of the possible slowdown, the following times were picked by 

"eyeball" from figure 3. The creep rate may have slowed 

about 0.3 yr before the LPEQ at Site 17, about 3.6 yr 

before at Site 13, about 1.4 yr before at Site 12, about 0.7 

yr before at Site 2, and about 1.4 yr before at Site 1. 

Although these eyeball estimates are suggestive that a 

slowdown in creep on the Hayward fault may have begun 

a few months to a few years before the LPEQ, more objective methods of evaluating the onset times of rate 

changes and a statistical measure of the confidence level 

of the proposed changes need to be developed before any 

pre-LPEQ slowdown can be considered likely. 

At the more recently installed Site 24 at Camellia Drive 

about 4 km southeast of Rockett Drive (Site 1) in Fremont, creep has only been measured subsequent to the 

LPEQ (first measurement on 11 February 1990). For the 

past 3 years, the fault has shown 0.9 mmlyr of left-lateral 

creep. Measurements for the past year at Parkmeadow 

Drive in Fremont (Site 27) only 0.4 km southeast of Camellia Drive also indicate a small amount of left-lateral 

creep (not shown on fig. 3). Sites 24 and 27 are along a 4- 

km-long segment in Fremont near the southeastern end of 

the fault that before the LPEQ had been creeping right 

laterally at about 8-1 1 mdyr since at least the 1920's 

(Harsh and Burford, 1982; Burford and Sharp, 1982; 

Lienkaemper and others, 199 1 ; Lienkaemper and 

Borchardt, 1992). Lienkaemper and Borchardt (1992) consider this rate significant because they think it may reflect 

the long-term surficial slip rate and the deep slip rate that 

controls the recurrence between large earthquakes on the 

Hayward fault. Recent measurements of an offset fence 

along Union Street in Fremont (about 0.9 km southeast of 

Rockett Drive and 3.2 km northwest of Camellia Drive) 

by J. Lienkaemper and us suggest that a few centimeters 

of left-lateral slip may have occurred here in conjunction 

with the LPEQ. 

In summary, our measurements at five sites (17, 13, 12, 

2, 1) along the Hayward fault show that the right-lateral 

creep rate ranged from 4.7 to 5.4 mmlyr for the decade 

preceding the LPEQ. Equivocal data suggest the possibility that a slowdown in creep may have occurred at these 

sites a few months to a few years before the LPEQ. No 

slip events appear to have been triggered at any of these 

sites by the quake. Since the LPEQ, however, rates have 

changed significantly at four of these five sites (see table 

1). The pre- to post-LPEQ rate decreased in San Pablo 

from 4.7 to 2.2 mdyr, at D Street in Hayward from 4.9 

to 3.6 mdyr, and at Rockett Drive in Fremont from 5.4 

to 1.1 mmlyr. The rate increased at Rose Street in Hayward 

from 4.9 to 6.8 mdyr and stayed nearly constant at 4.7 

mmlyr in Union City. In addition, it appears that a few 

centimeters of left-lateral slip associated with the LPEQ 

in southern Fremont may have occurred along thepreviously rapid, right-lateral-creeping southeastern segment of 

the fault. Slow, left-lateral creep at 0.9 mdyr has continued since the LPEQ along this segment of the fault at Site 

24, and Site 27 has also shown left-lateral creep since we 

began measurements in April 1992. 

Historic creep rates for the Hayward fault determined 

over the years by various investigators from alignment 

arrays and offset cultural features have been summarized 

recently by Lienkaemper and others (1991). The rate of 

right-lateral creep for the 40-50 years before the LPEQ 

has been between 3.5 and 6.5 mdyr for most of the 

Hayward fault, with a higher rate of about 9 mmlyr along 

a 4-km-long segment in southern Fremont. All of our preLPEQ measurements fall within the 3.5-6.5 mdyr range. 

As pointed out above, however, several sites along the 

fault are now creeping at rates below this range, including 

slow left-lateral creep in southern Fremont. This indicates 

that the LPEQ has had a significant effect on creep rates 

along the Hayward fault. 

CALAVERAS FAULT 

San Francisco State University began measuring creep 

at Sites 4 and 6 on the southern Calaveras fault in the 

Hollister area (fig. 1) in 1979. During the 10 years of 

measurements before the LPEQ, creep at both sites was 

quite episodic, with intervals of fast creep typically lasting a couple months or less, alternating with longer periods when little creep occurs (fig. 4). 

More specifically, in the decade before the LPEQ, Site 

4 along Seventh Street in the city of Hollister had nine 

episodes of fast creep of about 5 mm or more. This mode 

of movement had been noted previously by Langbein 

(198 I), based on data from a precision multi-wavelength 

distance-measuring instrument. He reported that four strain 
D200 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

episodes of creep of about 4-6 mm per episode occurred earthquakes. At Site 4, times of relatively rapid moveover time intervals of about 2 months or less during 1975 ment since at least 1979 alternate with intervals of little 

and 1976 along the Calaveras fault in the Hollister area. net movement typically lasting about 8-12 months, with 

He concluded that aseismic slip was the dominant mecha- one lasting 2 years between January 1986 and January 

nism for strain release, rather than slip associated with 1988. The LPEQ apparently triggered 14 mm of right0.2 Â 0.1 mmlyr for 12.2 yrs 

W = 11 1 .I meters 

160 

160 

140 

10.7 Â 0.3 mm/yr for 13.4 yrs 120 

W = 51.7 meters 100 

80 

60 

40 

20 

0 

-20 

1979 1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 

SF-19 SAN RAMON (Corey Place) 

Figure 4.-Calaveras fault displacement at three sites between 1979 and 1993. MHEQ = Morgan Hill earthquake of 24 April 1984. See 

fig. 2 for further explanation. 

- 140 
EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP RATES IN THE SAN FRANCISCO BAY REGION D201 

lateral slip at Site 4 that marked the abrupt end of an 

interval of slow creep that had persisted for about a year. 

Total cumulative displacement from the time of the first 

measurement in 1979 went from 67 to 81 mm. Following 

this triggered slip, Site 4 has been creeping at a rate of 

4.2 mdyr for the past 3.4 yr (through 13 March 1993). 

This is considerably lower than the pre-LPEQ rate of 6.4 

mrnlyr. 

Creep at Site 6 along Wright Road, located 2.3 km 

northwest of Site 4, included 11 episodes of fast creep of 

about 5 mm or more in the 10 years before the LPEQ. 

This faster creep alternates with intervals of little net movement that typically last about 3-12 months at Site 6, with 

one lasting 18 months between June 1985 and December 

1986. The LPEQ marked the end of an interval of slow 

creep that had persisted for about a year at Site 6, similar 

to the situation at Site 4. The earthquake apparently triggered 12 mm of right-lateral slip, and cumulative displacement increased from 119 to 13 1 mm. Following this 

triggered slip, the fault at Site 6 has virtually stopped 

creeping for the past 3.4 yr (through 13 March 1993). 

This has brought the pre-LPEQ average rate of 12.2 mm/ 

yr down to 10.7 mmlyr for the past 13.4 yr. As shown in 

figure 4, the fault at Site 6 now has a "slip deficit" of 

more than 2 cm. 

Creepmeter data also show that relatively long periods 

of time with little net movement have occurred along the 

Calaveras fault in the Hollister area. Evans and others 

(1981) note that the USGS creepmeter across the Calaveras fault at Shore Road 10 km north of Hollister showed 

virtually no net creep for a 3-year period from July 1976 

to July 1979. They point out that this period of no movement ended with a small slip event coincident with the 6 

August 1979 M=5.9 Coyote Lake earthquake (CLEQ) on 

the Calaveras fault about 12 km northwest of the 

creepmeter site. Accelerated afterslip in the form of a 

cluster of creep events followed the small slip event (Raleigh and others, 1979). 

Our theodolite data alone cannot prove that the LPEQ 

on the nearby San Andreas fault coseismically triggered 

aseismic slip on the Calaveras fault. Our last measurement before the earthquake was on 19 August 1989 and 

our first measurement after the quake was on 21 October 

1989. Strictly speaking, we can only say that 12-14 mm 

of displacement in the Hollister area occurred between 

these two dates. Other evidence, however, suggests that 

the timing of the slip was at least partially coseismic. We 

detected fresh en echelon cracks on Highway 25 near 

Hollister on the morning of 21 October 1989. These cracks 

extended from the asphalt into the dirt shoulder of the 

road. Fortunately, we were able to photograph them before they were obliterated by rainfall that occurred that 

afternoon. The cracks had to have formed sometime after 

the last previous rainfall on 29 September 1989 (McClellan 

and Hay, 1990) and probably formed shortly after the 

LPEQ (see the following discussion). 

The strongest evidence for coseismic slip on the Calaveras fault due to the LPEQ comes from the USGS 

creepmeter at Shore Road, 8.5 km northwest of Wright 

Road (Site 6). The LPEQ occurred at 1704 hours, Pacific 

Daylight Time (Plafker and Galloway, 1989). The Shore 

Road creepmeter recorded slip of 5.0 mm within a 10- 

minute interval from 1700 to 1710 hours (Schulz, 1989). 

It is unlikely that slip on the Calaveras fault preceded the 

LPEQ. It is more likely that shaking from the LPEQ triggered slip on the Calaveras fault at the exact time of the 

earthquake or within a few minutes after it. Additional 

evidence that slip on the Calaveras fault was triggered by 

the LPEQ comes from field observations made by 

McClellan and Hay (1990), who observed fresh en echelon cracks and offset cultural features that indicated at 

least 5 mm of movement along 17 km of the Calaveras 

fault from the city of Hollister northwest to Highway 152. 

Table 2 shows a comparison of Calaveras fault creep 

rates in the Hollister area determined by USGS creepmeters 

and San Francisco State University (SFSU) theodolite measurements. Both data sets span 10 years or more and overlap from the late 1970's to mid 1980's. The SFSU data 

exclude the LPEQ coseismic slip. Even though the measurement methods and time intervals are different, the 

creep rates determined by the USGS and SFSU are quite 

similar. Both indicate that the long-term creep rate at 

Wright Road (Site 6) is about 6 mdyr faster than the rate 

at nearby Seventh Street (Site 4). Either the creep rate on 

the Calaveras fault decreases significantly from Wright 

Road southeast to Seventh Street or undetected surface 

movement is occurring outside our 89.7-m-long survey 

line at Seventh Street, which encompasses the former 

USGS creepmeter site. Since the slip triggered by the 

LPEQ, the creep rate at both sites has decreased and the 

sites with faster and slower creep have reversed. Site 4 is 

now creeping at 4.2 mmlyr, whereas Site 6 has stopped 

creeping. 

In contrast to the evidence for creep and triggered slip 

along the southern segment of the Calaveras fault near 

Hollister, Site 19 in San Ramon near the northwestern 

terminus of the Calaveras fault has shown virtually no 

creep since measurements began in 1980 (fig. 4). 

Previous investigators have determined that significant 

earthquakes occurring on particular faults in California 

have triggered coseismic slip or afterslip on other nearby 

faults. Examples of such earthquakes include the Borrego 

Mountain earthquake of 9 April 1968 (Allen and others, 

1972), the Imperial Valley earthquake of 15 October 1979 

(Fuis, 1982; Sieh, 1982), the Livermore Valley earthquakes 

of late January 1980 (Harsh and Burford, 1982), the 

Coalinga earthquake of 2 May 1983 (Mavko and others, 

1985), the Tres Pinos earthquake of 26 January 1986 
D202 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 2.-Calaveras fault creep rates in the Hollister area measured by 

creepmeter and theodolite between 1970 and 1993 

[See fig. 1 for site locations.] 

Average creep rate 

Site (m ~IY r) 

Wright Road (SF-6) 

USGS (1 971 - 1 983)a 13.1 

SFSU (1 979 - LPEQ)' 12.2 

SFSU (LPEQ - 1993)' 0.0 

Seventh Street (SF-4) 

USGS (1970 - 1987)a 6.8 

SFSU (1 979 - LPEQ)' 6.4 

SFSU (LPEQ - 1993)b 4.2 

a U. S. Geological Survey (USGS) creepmeter data from Schulz 

and others (1982) and Schulz (1989). 

b San Francisco State University (SFSU) theodolite data before and 

after the Loma Prieta earthquake (LPEQ) from this paper. 

(Simpson and others, 1988), the North Palm Springs earthquake of 8 July 1986 (Williams and others, 1988), and the 

Superstition Hills earthquake of 24 November 1987 (Sharp, 

1989). 

In central California, USGS creepmeters on the San 

Andreas, Hayward, and Calaveras faults have recorded 

small, abrupt movements during or shortly after nearby 

moderate earthquakes as well as increased or decreased 

creep rates either before or after some of these quakes 

(Schulz and others, 1982). King and others (1977) point 

out that during 197 1-73, creepmeters along the creeping 

portion of the San Andreas fault in central California often recorded minor coseismic steps (<I mm) at times of 

local earthquakes of magnitudes 4-5. Although most of 

the nine moderate earthquakes that occurred locally during this time interval were on the San Andreas fault, some 

were not. In contrast, many larger creep events ( ^1 mm) 

occurred during this same interval at times of no local 

seismic events. King and others (1977) concluded that 

during 1971-73, coseismic slip was negligible compared 

to slip by episodic and continuous creep processes. They 

also concluded that creep episodes often follow seismic 

events at depth, suggesting that some creep may be afterslip 

delayed by different fault rheology near the surface. 

In our observations, some phases of rapid creep on the 

Calaveras fault in the Hollister area are apparently related 

to nearby seismic events, but others are not. Our data 

suggest that some significant nearby earthquakes trigger 

slip on the Calaveras fault, but others do not. In- addition 

to the relationship between the LPEQ and triggered slip 

on the Calaveras fault discussed above, there also appears 

to have been a period of rapid creep on the fault near 

Hollister that was related to the 1984 M=6.3 Morgan Hill 

earthquake on the Calaveras fault. However, no immediate surface displacement was observed at either of the 

Hollister sites when they were remeasured on 25 April 

1984, the day after the MHEQ (Galehouse, 1987), even 

though surface displacement had occurred 2.7 km northwest of Site 6 (Galehouse and Brown, 1987). However, 

within the following 2.5 months, Site 4 showed 11 mm of 

creep and Site 6 showed 10 mm of creep. This relatively 

large amount of post-seismic creep was about the same as 

the slip (13 mm) recorded in the 18 hours following the 

earthquake at the USGS creepmeter at Shore Road 8.5 km 

northwest of Site 6 (Schulz, 1987). 

Since detailed monitoring began more than 20 years 

ago, the southern Calaveras fault has had several intervals 

with relatively large amounts of creep (^ 1 cm) occurring 

within a relatively short time interval (months), but these 

intervals were apparently unrelated to any nearby significant earthquakes. Fast creep was recorded by two USGS 

creepmeters in Hollister in May-August 1977 (Evans and 

others, 1981). This rapid creep occurred during a 3-year 

interval when the Calaveras fault at Shore Road, 10 km to 

the northwest, was in a period of no movement. It also 

occurred with no apparent relationship to any significant 

earthquake(s) in the area. 

Another example of distinctly aseismic rapid creep occurred in the 14-month interval between June 1980 and 

August 1981 (fig. 4). The Calaveras fault at Site 4 showed 

creep of 19 mm during this interval, more than twice the 

predicted slip of about 8 mm obtained for this 14-month 

interval by extrapolating the pre-LPEQ least-squares average rate (6.4 mdyr). The fault at Site 6 showed creep 

of 23 mm, also considerably larger than the expected slip 

of 14 mm, based on the average rate (12.2 mdyr) for this 

interval. This rapid creep began 10 months after the M=5.9 

CLEQ on the Calaveras fault (epicenter about 30 km northwest of Hollister) and ended about 32 months before the 

M=6.3 MHEQ on the Calaveras fault (epicenter about 55 

km northwest of Hollister). Thus, this interval of rapid 

creep in the Hollister area appears to be unrelated to either large regional earthquakes (CLEQ, or MHEQ, or 

both), or any significant local earthquakes. 

CONCORD FAULT 

Sites 3 and 5 are in the city of Concord on what is 

known as the Concord segment of the Concord fault 

(Sharp, 1973). A detailed discussion of creep characteristics on the Concord fault is in Galehouse (1992). In general, the Concord fault creeps episodically, with intervals 
EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP RATES IN THE SAN FRANCISCO BAY REGION D203 

of fast creep of about 7-10 mm over aperiod of a few average creep rate on the Concord fault for at least the 

months alternating with intervals of slow creep of about past 13.4 yr is about 3 mm/yr (3.4 at Site 3 and 2.6 at Site 

1-2 mmlyr over a period of several years (fig. 5). This 5). Because of the episodic nature of creep on the Conpattern of episodic creep was slightly more pronounced cord fault and the lack of any triggered slip, it is difficult 

during the first eight to nine years of measurements. The to assess any effect(s) of the LPEQ. Both sites completed 

SF-20 GREEN VALLEY FAULT (near Cordelia) 

5.2 Â 0.3 mrnlyr for 8.6 yrs 

W = 335.8 meters 

I I I 

60 

3.4 Â 0.1 mmlyr for 13.4 yrs 50 

W = 130.0 meters 40 

30 

20 

10 

0 

-1 0 

Figure 5.-Concord fault displacement at two sites between 1979 and 1993 and Green Valley fault displacement at one site between 

1984 and 1993 . See fig. 2 for further explanation. 
D204 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

their latest phase of fast creep in February 1988 and overall have been creeping at about their longer-term average 

rate since then (through 7 March 1993). A least-squares 

line through only the post-LPEQ data points, however, 

shows a rate at both sites that is about 1-1.5 mdyr lower 

than the long-term average. However, given the episodic 

nature of the fault movement recorded over the years at 

Sites 3 and 5, we cannot definitely state that the rate has 

slowed due to the LPEQ. 

GREEN VALLEY FAULT 

Frizzell and Brown (1976) and Helley and Herd (1977) 

suggest that the southern extension of the Green Valley 

fault connects to the northern extension of the Concord 

fault by a relatively short (<3 km) right step (fig. 1). The 

overall average rate of creep for the past 8.6 yr on the 

Green Valley fault is 5.2 k 0.3 mdyr. This is virtually 

the same rate (5.4 mrnlyr between 1922 and 1974) that 

Frizzell and Brown (1976) determined from offset power 

transmission lines about 3 km south of Site 20. Although 

the creep rates are different, creep characteristics of the 

Green Valley fault are crudely similar to those of the 

Concord fault (fig. 5), although there is more "noise" in 

the Green Valley data, which is at least partly due to the 

relatively large fault width being measured (335.8 m). A 

detailed comparison is in Galehouse (1991 and 1992). In 

general, the Green Valley fault also creeps episodically, 

with intervals of fast creep of 10 mm or more over a few 

months alternating with intervals of slower creep up to 

about 3 mmlyr over a few years. In the 8.6 yr since we 

began our measurements, there have been two intervals of 

fast creep and three of slow creep. The most recent fast 

creep interval began sometime between 6 August 1989 

and 2 December 1989. Although there is a possibility that 

shaking from the 17 October 1989 LPEQ triggered the 

onset of this fast creep as it apparently did in Hollister, it 

is probably only a coincidence that the Green Valley fault 

began a period of faster creep at a time that can be crudely 

related to the quake that was centered more than 125 krn 

away. Our measurement intervals on the Green Valley 

fault were too far apart to pin down the exact time when 

rapid creep started, and we know of no other evidence 

suggesting a relationship. The latest phase of slow creep 

is continuing, and in fact there has been no net slip on the 

Green Valley fault for the past 2.5 yr (through 30 January 

1993), giving a post-LPEQ rate of only 2.9 mdyr. 

OTHER FAULTS 

We also have measurement sites on the Seal Cove-San 

Gregorio, Rodgers Creek, West Napa, and Antioch faults 

(fig. 1). None of these faults 

of fault creep and none have 

LPEQ thus far. 

SUMMARY AND 

show unequivocal evidence 

shown any effects from the 

CONCLUSIONS 

Reasenberg and Simpson (1992) point out that because 

the major San Francisco Bay area faults are probably 

loaded by right-lateral shear, the stress changes due to the 

LPEQ would either increase the shear load if the change 

was right-lateral or decrease the shear load if the change 

was left-lateral. Consequently, a right-lateral static stress 

change would tend to be associated with a post-LPEQ 

increase in seismicity rate and an increase in creep rate. A 

left-lateral change would tend to be associated with a decrease in seismicity rate and either a decrease in creep 

rate or even a reversal in creep direction if the static stress 

change was particularly significant. Reasenberg and 

Simpson (1992) found that their calculations of stress 

changes agreed closely with changes in the regional seismicity rate after the LPEQ. The data we have been collecting since the LPEQ are exactly what are needed to 

evaluate further the Reasenberg and Simpson (1992) calculations for static stress changes. For sites which clearly 

showed proposed stress changes, we compared our data 

with the Reasenberg and Simpson (1992) model. None of 

the stress changes were evidently large enough to cause 

noncreeping segments of faults to begin creeping. Whether 

the stress change was right-lateral or left-lateral, none of 

our measurement sites with pre-LPEQ rates less than 1 

mmlyr increased in rate after the LPEQ. In other words, 

fault segments that were virtually locked before the LPEQ 

remained locked after the quake. This suggests that none 

of the noncreeping areas that experienced an increase in 

right-lateral stress were loaded to their breaking (fault 

slip) stress level. Those that received a left-lateral change 

probably had their next time of surface rupture delayed. 

The amount of static stress change and degree of shaking 

from the earthquake were not enough to induce movement at any of these noncreeping fault sites. In areas that 

were already creeping, we can compare the creep rates 

before and after the LPEQ to see if the changes correspond to the Reasenberg and Simpson (1992) model. In 

11 of 13 cases, the changes in creep rates are consistent in 

sign with their model. In one additional case the data are 

consistent with the model, and in only one case is the 

change inconsistent. 

The San Andreas fault northwest of the LPEQ aftershock zone was not creeping before the quake, did not 

have any surface slip during the quake, and is still not 

creeping. There also has been no post-LPEQ creep at a 

site (23) just southeast of the aftershock zone. In contrast, 

about 42 km southeast of the epicenter, approximately 5 
EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP RATES IN THE SAN FRANCISCO BAY REGION D205 

mm of slip was triggered by LPEQ shaking near the northwestern end of the creeping segment of the fault near San 

Juan Bautista and the post-LPEQ creep rate at this site 

(25) is about twice the pre-LPEQ average. It appears that 

right-lateral shear stress from the Lorna Prieta region was 

transferred to the region southeast of the rupture following the LPEQ. 

Most of the Hayward fault creeps at a long-term rate of 

about 5 mmlyr. Five sites along the fault that we had been 

measuring for the decade before the LPEQ had rates between 4.7 and 5.4 mdyr, with a median rate of 4.9 mrnl 

yr. These same sites have a median rate of only 3.6 md 

yr during the 3.4 yr since the quake. The post-LPEQ rate 

is particularly low in San Pablo and Fremont, at either 

end of the fault. A few centimeters of left-lateral slip 

associated with the LPEQ may have occurred in southern 

Fremont, and small amounts of left-lateral creep are continuing. This left-lateral displacement is along the segment of the Hayward fault that was creeping right-laterally 

at about 9 mdyr for at least 50 yr before the LPEQ. 

Simpson and others (1988) point out two possible explanations for earthquake-related steps in creep (slip) along 

faults. Either the slip was triggered by earthquake shaking 

or the slip occurred in response to changes in static stress 

fields accompanying the earthquake. This left-lateral slip 

of a few centimeters in southern Fremont almost certainly 

would not have occurred as triggered slip. It is unlikely 

that the segment of the Hayward fault that had been creeping right-laterally at about twice the rate on the rest of the 

fault would have been loaded left-laterally and been primed 

to move in that direction. The left-lateral slip was almost 

certainly due to the nature of the static stress changes 

imposed upon the fault by the LPEQ, although stress 

changes due to unknown factors unrelated to the LPEQ 

cannot be completely ruled out. A decrease in right-latera1 static stress on the fault following the LPEQ was 

suggested by Reasenberg and Simpson (1992) based on 

their computer simulation of the earthquake and the postLPEQ decrease in microseismicity along the Hayward 

fault. Their model assumes low values for the apparent 

coefficient of friction along the fault and agrees well with 

our measurements of low creep rates following the LPEQ. 

The segment of the Hayward fault in Fremont that moved 

and is continuing to move left-laterally, however, probably had a larger left-lateral stress change than indicated 

in the Reasenberg and Simpson (1992) model. The postLPEQ slowdown in right-lateral creep over much of the 

rest of the Hayward fault is fully consistent with the sign 

of the predicted LPEQ stress change in their model. The 

model, however, does not help explain the puzzling postLPEQ increase in creep rate at Rose Street in Hayward or 

a pre-LPEQ slowdown along the Hayward fault, if that 

indeed occurred. 

The average creep rate on the southern Calaveras fault 

in the Hollister area before the LPEQ was 6.4 mdyr at 

Site 4 and 12.2 mrn/yr at Site 6. The LPEQ triggered 12- 

14 mm of slip here, less than 50 km from the epicenter. A 

similar amount of slip occurred at these sites after the 

MHEQ in 1984. It appears that the rapid slip that occurred in the Hollister area on the Calaveras fault in conjunction with the MHEQ in 1984 and the LPEQ in 1989 

were of different character. The rapid slip associated with 

the MHEQ occurred during the 2.5 months following the 

quake, but no slip had occurred when measurements were 

taken the first day after the quake. This rapid slip was 

most probably in response to a change in the static stress 

field. In contrast, the rapid slip associated with the LPEQ 

almost certainly was triggered by earthquake shaking. For 

decades, the typical mode of movement has been episodic 

in the Hollister area, with longer periods of slower creep 

alternating with shorter periods of faster creep. When no 

significant earthquakes occur in the Hollister area over a 

relatively long period, a rapid phase of movement will 

occur with no association with any seismic event(s). When 

a significant earthquake occurs on either the Calaveras or 

another nearby fault and the Hollister area segment of the 

Calaveras fault is primed to move anyway, the earthquake 

will probably trigger a relatively rapid phase of movement. This may occur as coseismic slip triggered by shaking (as with the LPEQ) or as accelerated creep when a 

change in the static stress field caused by the earthquake 

is right-lateral (as with the MHEQ). On the other hand, 

when the fault is not primed, a nearby earthquake will not 

trigger any slip. Because the lengths of time of the slower 

and faster phases are so variable, predicting whether or 

not the southern Calaveras fault is "primed and ready to 

go" is extremely difficult. At the present time (March 

1993), there appears to be a "slip deficit" of more than 2 

cm at Site 6 in the Hollister area (fig. 4). A post-LPEQ 

decrease in the rate of right-lateral creep on the Calaveras 

fault in the Hollister area is qualitatively consistent with 

the static stress change following the LPEQ proposed by 

Reasenberg and Simpson (1992). 

The observed decrease in creep rates on the Concord 

and Green Valley faults is also qualitatively consistent 

with the stress changes estimated for these sites with the 

Reasenberg and Simpson (1992) model. The noncreeping 

northern Calaveras, Seal Cove-San Gregorio, Rodgers 

Creek, West Napa, and Antioch faults were apparently 

not affected by the LPEQ. 

In summary, changes in creep rates on San Francisco 

Bay region faults after the LPEQ are consistent with stress 

changes estimated for the LPEQ by Reasenberg and 

Simpson (1992). At one measurement site (25) on a creeping fault where Reasenberg and Simpson (1992) show a 

right-lateral static shear stress increase, the creep rate has 

doubled since the LPEQ. Among 12 sites on creeping 

faults where Reasenberg and Simpson (1992) calculate a 

left-lateral change (a decrease in right-lateral shear load), 

the post-LPEQ creep rate slowed significantly (>1.2 mml 
D206 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

yr) at ten of these sites (1, 3, 4, 5, 6, 12, 17, 20, 24, 27), - 1990, Effect of the Loma Prieta earthquake on surface slip 

stayed virtually the same at one (2), and increased at only along the Calaveras fault in the Hollister area: Geophysical Reone (13). As Simpson and others (1988) conclude, it is search Letters, v. 17. no. 8, p. 1219-1222. 

1991, Present-day creep on the Green Valley fault, in Figuers, not yet possible to explain unambiguously all the observed S.H., ed., Field Guide to the Geology of Western Solano County: 

creep on certain faults during and after earthquakes on Northern California Geological Society and Association of Engiother faults. We do, however, seem to be making some 

progress in understanding and predicting these relationships. 

ACKNOWLEDGMENTS 

I thank the many geology students at San Francisco 

State University who have been instrumental in helping 

collect this theodolite data since 1979. I am particularly 

grateful to Carolyn Garrison, Oliver Graves, Theresa Hoyt, 

and Carl Schaefer, who are presently working with me on 

this ongoing research effort. I also thank J. Lienkaemper, 

P. Reasenberg, and R. Simpson of the USGS for helpful 

suggestions and thoughtful reviews of the manuscript. The 

present work is supported by the USGS, Department of 

the Interior, under award number 14-08-0001 -G 1992. 

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Allen, C.R., Wyss, M., Brune, J.N., Grantz, A., and Wallace, R.E., 

1972, Displacements on the Imperial, Superstition Hills and the 

San Andreas faults triggered by the Borrego Mountain earthquake, 

in The Borrego Mountain earthquake of April 9, 1968: U.S. Geological Survey Professional Paper 787, p. 87-104. 

Bilham, R., 1992, Creepmeters on California faults: U.S. Geological 

Survey Open-File Report 92-258, p. 153- 156. 

Breckenridge, K.S., and Burford, R.O., 1990, Changes in fault slip 

near San Juan Bautista, California, before the October 17, 1989, 

Loma Prieta earthquake-a possible precursor? [abs.]: Eos (American Geophysical Union Transactions), v. 71, no. 43, p. 1461. 

Burford, R.O., and Sharp, R.V., 1982, Slip on the Hayward and Calaveras faults determined from offset powerlines, in Hart, E.W., 

Hirschfeld, S.E., and Schulz, S.S., eds., Proceedings - Conference 

on Earthquake Hazards in the Eastern San Francisco Bay Area: 

California Division of Mines and Geology, Special Publication 62, 

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Evans, K.F., Burford, R.O., and King, G.C.P., 198 1, Propagating episodic reep and the aseismic slip behavior of the Calaveras fault 

north of Hollister, California: Journal of Geophysical Research, v. 

y[21-3735. 

Frizz 11 V.A., Jr., and Brown, R.D., Jr., 1976, Map showing recently 

acti e breaks along the Green Valley fault, Napa and Solano 

Counties, California: U.S. Geological Survey Miscellaneous Field 

Studies Map MF-743, scale 1 :24,000. 

Fuis, G.S., 1982, Displacement on the Superstition Hills fault triggered 

by the earthquake, in The Imperial Valley, California, earthquake 

of October 15, 1979: U.S. Geological Survey Professional Paper 

1254, p. 145-154. 

Galehouse, J.S., 1987, Theodolite measurements, in Hoose, S.N., ed., 

The Morgan Hill, California, earthquake of April 24, 1984: U.S. 

Geological Survey Bulletin 1639, p. 12 1- 123. 

neering Geologists Guidebook, p. 12- 16. 

1992, Creep rates and creep characteristics of eastern San Francisco Bay area faults: 1979-1992, in Borchardt, G. and others, 

eds., Proceedings of the Second Conference on Earthquake Hazards in the Eastern San Francisco Bay Area: California Division 

of Mines and Geology, Special Publication 113, p. 45-53. 

Galehouse, J.S., and Brown, B.D., 1987, Surface displacement near 

Hollister, California, in Hoose, S.N., ed., The Morgan Hill, California, earthquake of April 24, 1984: U.S. Geological Survey Bulletin 1639, p. 69-72. 

Galehouse, J.S., Brown, B.D., Pierce, B., and Thordsen, J.J., 1982. 

Changes in movement rates on certain East Bay faults, in Hart, 

E.W., Hirschfeld, S.E., and Schulz, S.S., eds., Proceedings - Conference on Earthquake Hazards in the Eastern San Francisco Bay 

Area: California Division of Mines and Geology, Special Publication 62, p. 239-250. 

Gladwin, M.T., Breckenridge, K.S., Hart, R.H.G., and Gwyther, R.L., 

1991, Recent acceleration of characteristic creep-strain events at 

San Juan Bautista [abs.]: Eos (American Geophysical Union 

Transactions), v. 72, no. 44, p. 484. 

Harsh, P.W., and Burford, R.O., 1982, Alinement-array measurements 

of fault slip in the eastern San Francisco Bay area, California, in 

Hart, E.W., Hirschfeld, S.E., and Schulz, S.S., eds., Proceedings - Conference on Earthquake Hazards in the Eastern San Francisco 

Bay Area: California Division of Mines and Geology, Special 

Publication 62, p. 251-260. 

Helley, E.J., and Herd, D.G., 1977, Map showing faults with Quaternary displacement, northeastern San Francisco Bay region, California: U.S. Geological Survey Miscellaneous Field Studies Map 

MF-881, scale 1: 125,000. 

King, C.Y., Nason, R.D., and Burford, R.O., 1977, Coseismic steps 

recorded on creep meters along the San Andreas fault: Journal of 

Geophysical Research, v. 82, p. 1655-1662. 

Langbein, J.O., 1981, An interpretation of episodic slip on the Calaveras fault near Hollister, California: Journal of Geophysical Research, v. 86, p. 4941-4948. 

1990, Post-seismic slip on the San Andreas fault at the northwestern end of the 1989 Loma Prieta earthquake rupture zone: 

Geophysical Research Letters, v. 17, no. 8, p. 1223-1226. 

Lienkaemper, J.J., and Borchardt, G., 1992, Hayward fault: Large 

earthquakes versus surface creep, in Borchardt, G. and others, 

eds., Proceedings of the Second Conference on Earthquake Hazards in the Eastern San Francisco Bay Area: California Division 

of Mines and Geology, Special Publication 1 13, p. 101 - 1 10. 

Lienkaemper, J.J., Borchardt, G., and Lisowski, M., 1991, Historic 

creep rate and potential for seismic slip along the Hayward fault, 

California: Journal of Geophysical Research, v. 96, no. Bll, p. 

18, 261-18, 283. 

Mavko, G.M., Schulz, S.S., and Brown, B.D., 1985, Effects of the 

1983 Coalinga, California earthquake on creep along the San Andreas fault: Seismological Society of America Bulletin, v. 75, p. 

475-489. 

McClellan, P.H., and Hay, E.A., 1990, Triggered slip on the Calaveras 

fault during the magnitude 7.1 Loma Prieta, California, earthquake: Geophysical Research Letters, v. 17, no. 8, p. 1227-1230. 

Plafker, G., and Galloway, J.P., eds., 1989, Lessons learned from the 

Loma Prieta, California, earthquake of October 17, 1989: U.S. 

Geological Survey Circular 1045, 48 p. 

Raleigh, C.B., Stuart, W., and Harsh, P., 1979, Creep on the Calaveras 
EFFECT OF THE LOMA PRIETA EARTHQUAKE ON FAULT CREEP RATES IN THE SAN FRANCISCO BAY REGION D207 

fault near Coyote Lake [abs.]: Eos (American Geophysical Union 

Transactions), v. 60, p. 890. 

Reasenberg, P.A., and Simpson, R.W., 1992, Response of regional 

seismicity to the static stress change produced by the Loma Prieta 

earthquake: Science, v. 255, p. 1687-1690. 

Schulz, S.S., 1987, Response of U.S. Geological Survey creepmeters 

near Hollister, California, in Hoose, S.N., ed., The Morgan Hill, 

California, earthquake of April 24, 1984: U.S. Geological Survey 

Bulletin 1639, p. 111-115. 

1989, Catalog of creepmeter measurements in California from 

1966 through 1988: U.S. Geological Survey Open-File Report 89- 

650, 193 p. 

Schulz. S.S., Mavko, G.M., Burford, R.O., and Stuart, W.D., 1982, 

Long-term fault creep observations in central California: Journal 

of Geophysical Research, v. 87, p. 6977-6982. 

Sharp, R.V., 1973, Map showing recent tectonic movement on the 

Concord fault, Contra Costa and Solano Counties, California: U.S. 

Geological Survey Miscellaneous Field Studies Map MF-505, 

scale 1 :24,000. 

1989, Pre-earthquake displacement and triggered displacement 

on the Imperial fault associated with the Superstition Hills earthquake of 24 November 1987: Seismological Society of America 

Bulletin, v. 79, p. 466-479. 

Sieh, K.E., 1982, Slip along the San Andreas fault associated with the 

earthquake, in The Imperial Valley, California, earthquake of October 15, 1979: U.S. Geological Survey Professional Paper 1254, 

p. 155-159. 

Simpson, R.W., Schulz, S.S., Dietz, L.D., and Burford, R.O., 1988, 

The response of creeping parts of the San Andreas fault to earthquakes on nearby faults: two examples: Pageoph, v. 126, nos. 2-4, 

p. 665-685. 

Williams, P.L. and others, 1988, Triggered slip along the San Andreas 

fault after the 8 July 1986 North Palm Springs earthquake: Seismological Society of America Bulletin, v. 78, p. 11 12-1 122. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA 

EARTHQUAKE FROM REPEATED GPS MEASUREMENTS 

By Roland Biirgmann, University of California, Davis; 

Paul Segall, Stanford University; and 

Mike Lisowski and Jerry L. Svarc, U.S. Geological Survey 

CONTENTS INTRODUCTION 

Page 

D209 

209 

21 1 

21 1 

21 1 

225 

227 

227 

230 

233 

234 

234 

236 

237 

24 1 

24 1 

ABSTRACT 

The postseismic velocity field in the epicentral region 

of the Loma Prieta earthquake differs significantly from 

displacement rates measured in the two decades preceding the event. The post-earthquake displacement rates along 

the Black Mountain profile, which crosses the San Andreas fault 44 km northwest of the Loma Prieta epicenter, 

do not differ significantly from those determined from 20 

years of trilateration measurements. However, station velocities along the Loma Prieta profile, which passes 

through the epicentral region, significantly exceed preearthquake rates within 20 km of the Loma Prieta rupture. 

There is also significant San Andreas normal shortening 

centered near Loma Prieta. Aseismic right-lateral strike 

slip at a rate of 0.10 m/yr on the Loma Prieta rupture and 

reverse slip at a rate of 0.12 mlyr on a reverse fault of the 

Foothills thrust belt can explain the observed postseismic 

deformation. Slip within the Foothills thrust belt appears 

to have been triggered by the stressing during the Loma 

Prieta main shock. 

Geodetic observations have shown accelerated strain 

rates adjacent to faults in the years to decades following 

some large earthquakes (for example, Thatcher, 1986). 

Following the 1906 San Francisco earthquake, which ruptured a 400 km segment of the San Andreas fault (Lawson, 

1908), shear strain rates near Point Reyes and Point Arena 

were 2 to 3 times greater than rates measured in the 1970's 

(Thatcher, 1974). Postseismic transients have also been 

observed following some subduction-zone earthquakes in 

Japan (Kasahara, 198 1 ; Thatcher, 1986). 

Transient postseismic deformation has generally been 

explained either by viscous relaxation of a ductile 

(asthenospheric) layer underlying an elastic (lithospheric) 

plate, or by downward propagation of aseismic slip along 

an extension of the fault zone into the lower crust (Bott 

and Dean, 1973; Nur and Mavko, 1974; Thatcher, 1974; 

Anderson, 1975; Budianski and Amazigo, 1976; Rundle 

and Jackson, 1977; Savage and Prescott, 1978; Cohen, 

1979; Lehner and others, 1981; Thatcher, 1983; Li and 

Rice, 1987). Unfortunately, there is a lack of adequate 

data to test and differentiate among the various kinematic 

models. 

Earthquakes have been observed to propagate along fault 

zones, rupturing adjacent fault segments over the time 

span of several years. Examples involving strike-slip faults 

have been reported in northeast China preceding the 1975 

Haicheng earthquake (Scholz, 1977), and along the North 

Anatolian (Toksoz and others, 1979), the Calaveras 

(Oppenheimer and others, 1990), and the San Andreas 

(Savage, 1971; Wood and Allen, 1973) fault zones. In the 

19th century, earthquakes along the San Andreas fault 

and the Hayward fault in the San Francisco Bay area occurred in clusters spread over several years (Ellsworth, 

1990). These observations suggest the existence of transient loading processes that occur with characteristic time 

scales of several years. However, without more data on 

time-dependent deformation following large earthquakes, 

we can only speculate on the mechanics of postseismic 
D210 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

deformation and the relation between strain transients and 

the clustering of earthquakes. 

The Loma Prieta earthquake was the largest earthquake 

in California in 40 years and the largest earthquake in the 

San Francisco Bay area since 1906. The horizontal deformation field in the epicentral region prior to the 1989 

earthquake was already well established on the basis of 

20 years of Geodolite measurements (Lisowski and others, 199la) and several years of Global Positioning System (GPS) measurements (Davis and others, 1989; 

Lisowski and others, 1990b). Because of the extensive 

pre-earthquake monitoring, it is possible to identify fea-I 

tures in the present-day deformation field that are attributable to the earthquake. Furthermore, the source 

characteristics of the Lorna Prieta earthquake itself are 

well established by both seismic and geodetic measurements (Beroza, 1991; Hartzell and others, 1991; Steidl 

and others, 1991; Wald and others, 1991; Lisowski and 

others, 1990a; Marshall and others, 1991; Williams and 

others, 1993; Arnadottir and Segall, 1994). The magnitude of the earthquake, together with the extensive preand co-seismic geodetic monitoring, make the Lorna Prieta 

earthquake an important source of data for the study of 

postseismic processes. The possibility that accelerated 

postseismic straining could advance the occurrence of other 

earthquakes in the San Francisco Bay area adds additional 

importance to the study. 

Postseismic slip at shallow depths along the trace of the 

San Andreas fault was generally less than 1 cm in the first 

year following the Loma Prieta earthquake (Behr and others, 1990; Langbein, 1990; Rymer, 1990). The cumulative 

slip of aftershocks on the Lorna Prieta rupture and adjacent faults is about 5 mm (King and others, 1990), which 

is too small to be detected geodetically. Behr and others 

Figure 1.- Location map of San Francisco Bay area showing major fault zones and localities discussed in text. 

BeF = Berrocal fault, SgtFZ = Sargent fault zone, ShF = Shannon fault, ZF = Zayante fault. 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D2 1 1 

(this chapter) find that slip rates on the creeping section 

of the San Andreas fault near San Juan Bautista accelerated from a pre-earthquake rate of 7 to 8 mdyr to a rate 

of about 13 mmlyr following the earthquake. About 2 to 3 

cm of excess slip occurred on the northernmost 15 km of 

the creeping section in the 3 years following the earthquake, probably down to a depth of 1 to 3 km (Behr and 

others, this chapter; Gwyther and others, 1992). Surficial 

creep rates on the southern Hayward fault have decreased 

to about half the pre-earthquake rates of 3 to 9 mdyr. 

Near the southeast end of the Hayward fault, where the 

pre-earthquake-creep rate was highest, the fault has been 

slipping left-laterally since the earthquake (Lienkaemper 

and others, 1992) with a possible reversal in early 1993 

(Lienkaemper, oral commun., 1993). 

In this study we report on post-earthquake GPS measurements from the Black Mountain profile, which crosses 

the San Andreas fault 44 km northwest of the Loma Prieta 

epicenter. We analyze these data together with GPS data 

from a similar profile (the Loma Prieta profile) which 

crosses the SAF through the epicentral region (Savage 

and others, 1994), as well as from the San Francisco Bay 

area monitor network (fig. 2). The post-earthquake network geometry was chosen to measure postseismic transients away from and along the San Andreas fault trace. 

Twenty years of trilateration measurements prior to the 

earthquake constrain the preseismic displacement field 

(Lisowski and others, 199la). The five-station San Francisco Bay area monitor network and two VLBI (Very Long 

Baseline Interferometry) sites had also been repeatedly 

measured for several years prior to the earthquake 

(Lisowski and others, 1990b; Clark and others, 1990). 

Our objectives are to characterize the deformation of 

the crust following the Loma Prieta earthquake, to determine the sources of the postseismic deformation, and to 

assess the implications of these results for earthquake hazards in the San Francisco Bay region. 

METHODS: GPS DATA COLLECTION, 

PROCESSING, AND ERROR ANALYSIS 

DATA COLLECTION 

A total of 17 occupations during 11 campaigns of the 

Black Mountain profile (fig. 2) were carried out with dual 

frequency CIA-code Trimble 4000, P-code TI 4100, and 

CIA-code Ashtech LM-XI1 receivers. The results we 

present here are derived from 9 campaigns between January 1990 and January 1993 (fig. 3). We also interpret data 

from 9 occupations of the Loma Prieta profile during the 

first 2 years following the Loma Prieta earthquake 

(Lisowski and others, 1991b; Savage and others, 1994), 

and 12 to 14 occupations of the San Francisco Bay area 

monitor network baselines (fig. 4). Table 1 shows the 

names and site coordinates of the postseismic GPS stations. Table 2 details the occupation dates and receiver 

type used in the Black Mountain profile. The Loma Prieta 

profile and the monitor network were surveyed by the 

U.S. Geological Survey with TI 4100 receivers until August 1991 and Ashtech receivers since that time. 

TI 4100 and Ashtech carrier phase and pseudorange 

data were collected at 30-s intervals for approximately 6 

hours, whereas the Trimble receivers recorded data at 15- 

s intervals for 6 to 14 hours. Occupations exceeding 8 

hours in the May 1992 campaign appear to have contributed significantly to the precision (see below) of this survey. Survey times were chosen to maximize the number 

of satellites visible during an experiment. 

DATA ANALYSIS 

Our analysis of the GPS data follows the methods described by Davis and others (1989). Doubly differenced, 

ionosphere-free carrier phase observations were processed 

to estimate station coordinates, tropospheric zenith delays, satellite orbit parameters, and integer phase ambiguities. We used the Bernese GPS analysis software (Versions 

3.2 and 3.3) for parameter estimation (Beutler and others, 

1987; Rothacher and others, 1990). 

Errors in the orbit information broadcast by the GPS 

satellites can be significantly reduced by monitoring the 

satellites from global tracking networks with well known 

station coordinates (for example, Larson and others, 1991). 

We improved GPS satellite orbits using data from fiducial 

stations of the CIGNET network; usually Mojave, California, Westford, Massachusetts, and Richmond, Florida. 

The fiducial sites were constrained to their SV6 coordinates (M. Murray, written commun., 1992). During the 

June 1991 campaign we added data from tracking stations 

at Penticton and Yellowknife in Canada, and from Kokee 

in Hawaii to compensate for the lack of data from the 

three North American stations. In three campaigns (January and September, 1990 and February, 1991) we have 

not been able to produce reliable improved orbits with the 

tracking data from the CIGNET stations. In these campaigns, broadcast orbits with inferred precision of about 5 

parts in lo7 (Hager and others, 1991) were used. 

For the Trimble and Ashtech data, we determined receiver clock corrections at each measurement point using 

CIA-code pseudorange data. For the TI 4100 data we estimated corrections with a polynomial in time using P-code 

pseudorange data (Davis and others, 1989). The data were 

processed with an automatic program to fix cycle slips, 

and then all doubly differenced carrier-phase residuals were 

visually inspected to remove any remaining cycle slips 

and outliers. 
D212 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

All data were processed with a satellite elevation cutoff angle of 15O, because higher cut-off angles appear to 

reduce the short-term repeatability of the station coordinates in the June 1991 survey. Carrier phase data from 

GPS satellites are known only up to an integer number of 

wavelengths. We resolve these so-called integer ambiguities using phase and pseudorange measurements (Blewitt, 

1989). We initially resolved "wide-lane" ambiguities-a 

linear combination of the Ll and L2 carrier frequencieswith a wavelength of 86 cm. Using the resolved widelane ambiguities, we fixed the remaining linearly 

independent integer-cycle ambiguities using the ionosphere-free linear combination. This combination of the 

two L-band frequencies (L3) eliminates the first-order effects of the dispersive ionospheric delays. Practically all 

integer-cycle ambiguities in the doubly differenced phase 

Figure 2.- Trilateration network, GPS, and VLBI station locations in the south San Francisco Bay-Monterey Bay area. Pre-earthquake 

data are available from the trilateration network, VLBI measurements at Presidio and Ft. Ord, and GPS occupations of the Bay area 

monitor network (underlined station names). Postseismic measurements are available from the VLBI sites, the Bay area GPS monitor 

network, the Black Mountain profile, and the Loma Prieta profile. 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATFiD GPS MEASUREMENTS D213 

m Bend - Pigeon Pt. m 

m Bend - Halmaur m 

Figure 3.- Relative changes in the 

north, east, and vertical baseline components and line length changes in the 

Black Mountain profile with respect 

to station Bend as a function of time. 

The line is a least-squares fit to the 

data. Error bars are k 1 sigma based 

on the observed residuals to the linear 

fit. Standard errors are 6 mm in the 

north and 7.1 mm in the east component. The line fit in fig. 31 is based on 

the data collected after a benchmark 

disturbance at Bayshore in 1990. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

m Bend - Haul m 

Bend - True 

Figure 3.- Continued 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D215 

m Bend - Pawt m 

Bend - BMI 0 m 

Figure 3.- Continued 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Bend - Foothill m 

m Bend - Antelynx m 

Figure 3.- Continued 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Lp1 - Cliff m 

-25569.42 

Normal 

m Lp1 - Gregor m 

I Normal \ 

Figure 4.- Relative changes in the 

north, east, San Andreas fault-normal 

and fault-perpendicular baseline components in the Loma Prieta profile with 

respect to station LP1 as a function of 

time. The line is a least-squares fit to 

the data. Error bars are Â 1 sigma 

based on Bernese formal errors scaled 

by a factor of 5. Significant fault-norma1 motions occur at sites within about 

20 km of the Loma Prieta rupture. 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D219 

m Lp1 - Fire m 

Figure 4.- Continued 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

rn Lpl - Lp2 m 

- End m 

- -947.54 

Normal : 

-5428.94 

i -5428.95 

{ -5428.96 

< -5428.97 

Figure 4.- Continued 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D221 

m Lpl - Mazzone m 

5752.51 

Normal 

m Lpl - Calero m 

Figure 4.- Continued 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

m Lpl - coy m 

,,,,,l,L,,,l,,r,,ll,,,3[,,,,,~,,,,, ,,,,,~,,,,,,l,,,ll,,,,l,,,,,,~,,,,c 

m Lpl - Hamilton m 

L 1 Normal 1 

Figure 4.- Continued 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D223 

m Lpl - Mt. OSO m 

Figure 4.- Continued 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

m Lpl - Eagle Rk. m 

m Lpl - Allison m 

A Normal 

Figure 4.- Continued 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D225 

measurements were reduced to integer values in the local 

data. Final coordinate estimates of the local stations were 

determined by processing the combined local and fiducial 

data using the ionosphere-free linear combination of the 

Ll and L2 phases and the resolved integer biases. We 

simultaneously estimated atmospheric propagation delay 

parameters (solving for a constant zenith delay parameter 

for each station for 8-hour periods), satellite orbit parameters, and clock synchronization errors (Davis and others, 

1989). 

DATA PRECISION AND ACCURACY 

The short-term precision of GPS measurements depends 

on error sources related to satellite and receiver clocks, 

propagation delays, satellite orbit and constellation characteristics, length and time of observation, the geographic 

location of a network, receiver types, post-processing software, and set-up procedures (for example, Larson and 

Agnew, 1991). By using optical plummets and a careful 

centering routine with several redundant height measurements, we estimate the absolute antenna set-up errors to 

be less than 1 to 2 mm in all coordinate components. Four 

receiver station days were left out in the final computations because operators noted significant shifts in antenna 

positions caused by bovine impact or tripod instability 

during the surveys (occupations of Halmaur in May 1992 

and January 1993, Antelynx in May 1992, and Bayshore 

in January 1993). Benchmark occupations usually lasted 

about 6 hours, except in the May 1992 survey when occupation times of up to 14 hours were used. 

Station Bayshore was disturbed by vehicle impact in 

the fall of 1990. Following this, a tie between the old 

Bayshore benchmark and a new reference mark (Bayshore 

RM 1) was measured in December 1990. In May and June 

1991 the original Bayshore benchmark was re-surveyed. 

Figure 37 shows the relative displacement before and after 

the benchmark disturbance determined by subtracting the 

tie from the later measurements of Bayshore RM 1. Because the tie was measured after the disturbance, we do 

not know the true offset. The station appears to have been 

offset to the north in the disturbance (fig. 34. Therefore, 

we estimated the displacement rate of Bayshore using only 

data collected after the incident, from February 1991 to 

January 1993 (fig. 30, which results in a larger error associated with the velocity vector at this site. 

During the June 1991 and May 1992 campaigns, most 

of the baselines were measured two to four times to estimate the precision of the survey results. Furthermore, the 

Lp1 - Brush 2 

Figure 4.- Continued 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Table 1 .-Coordinates and names of benchmarks surveyed following the 1989 Lorna Prieta earthquake 

Station Stamping Latitude Longitude 

0 I 'I 0 I I' 

Black Mountain profile 

Pigeon Point 

Halmaur 

Haul 

True 

BM10 

PAWT 

Foothill 

Bend 

Bay shore 

Bayshore RM 1 

Antelynx 

PROP. COR. 1973 

HALMAUR 

HAUL 1989 

TRUE 1990 

BMT10RM1 1977 

PAWT ERC BM 1 

FOOTHILL 1 989 

BEND SURVEY PT 1987 

BAYSHORE 1989 

RM 1 

ANTELYNX 1989 

Bay Area monitor network 

Eagle Rock 

Allison ALLISON V A 2659 

Hamilton 'unstamped" 

LP 1 LOMA PRIETA 1 

Brush 2 BRUSH 2 1978 

Loma Prieta profile 

Cliff CLIFF 193 1 

Gregor GREGOR 1989 

Fire FIRE 1989 

LP2 SKYLAND 2 

LP4 BUZZARD 

End END 1989 / nd* 

LP 1 LOMA PRIETA 1 

Mazzone MAZZONE 1989 

Calero CALERO 1989 1 ca* 

COY COY 1989 

Hamilton 'unstamped" 

Mocho MOCHO 1875 1969 

Mt. Oso MT OSO NO. 1 1931 

* Calero and End benchmarks were damaged and replaced in 1991 

profile was measured within the same week with TI 4100 

and Trimble receivers in January, April, and September 

of 1990 (table 2). The average, unweighted repeatabilities 

of the measured baselines in the June 1991 survey are 5.3 

mm in the north, 5.8 mm in the east, and 20.1 mm in the 

vertical component. The repeatabilities of baselines formed 

to Foothill, which was a commonly occupied site prior to 

June 1991, are about twice that of the other baselines in 

the June 1991 campaign. We believe these high 

repeatabilities are due to the poor location of Foothill 

near a fence and high vegetation. The average precision 

of repeated station estimates in the May 1992 campaign is 

2.4 mm in the north, 5.5 mm in the east, and 15.8 mm in 

the vertical components. Occupations of up to 14 hours 

clearly improved the precision of these measurements; 

however, other factors, such as satellite coverage and operator errors may have affected the reported precision. 

The TI 4100 and Trimble results in April of 1990, which 

were reduced with good tracking data, agree at the 3.9- 

mm, 3.3-mm, and 13-mm level in the north, east, and 

vertical components, respectively. The day-to-day precision of repeated measurements is nearly independent of 

the length of the orbital arcs we estimate (1 to 4 days). 

The short-term repeatability is also about the same for 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D227 

Table 2.- Occupation schedule and receiver types used during occupations of the Black Mountain profile between January 1990 and January 

1993 

[Short term repeatabilities given in the text are derived from multiple occupations in April 1990, June 1991 (each station occupied at least three times), and May 1992 

(each station occupied at least twice)] 

Jan 90 Apr 90 Sep 90 Feb 91 May 91 June 91 Mar 92 May 92 Jan 93 

Pigeon Point SDT+TI SDT+TI SST+TI TI SST SST AS SST AS 

Halmaur SDT+TI SDT SST+TI TI SST SST AS 

Haul SDT+TI SDT+TI SST+TI TI SST SST AS 

True SDT+TI SDT+TI SST+TI TI SST SST AS 

BM10 SDT+TI SDT+TI SST+TI TI SST SST AS 

PAWT SDT+TI SDT+TI SST+TI TI SST SST AS 

Foothill SDT+TI SDT+TI SST+TI TI SST SST AS 

Bend SDT+TI SDT+TI SST+TI TI SST SST AS 

Bay shore SDT+TI SDT+TI SST+TI SST SST 

Bayshore RM 1 TI AS 

Antelynx SDT+TI SDT+TI SST+TI TI SST SST AS 

SDT - Trimble 4000 SDT 

SST - Trimble 4000 SST 

TI -TI4100 

AS - Ashtech LM XI1 

broadcast and improved orbits; however, the broadcast 

orbit solutions may contain rotational or scale biases at 

the few parts in the lo7 level. 

RESULTS 

PRE-LOMA PRIETA EARTHQUAKE 

DISPLACEMENT FIELD 

Trilateration measurements have been conducted 

throughout the San Francisco Bay area during the last two 

decades (Lisowski and others, 199la). Except for the effects of the 1979 Coyote Lake and the 1984 Morgan Hill 

earthquakes, the velocity field appears to have been steady 

during that time period. Figure 5 shows the velocity field 

in the region computed from: (1) 20 years of trilateration 

data (1970 to 1989), (2) several years of GPS data for five 

Bay area monitor stations (Lisowski and others, 1990b), 

and (3) 6 years of Very Long Baseline Interferometry 

(VLBI) data collected at Presidio, Fort Ord, and Fort Ord 

S (Clark and others, 1990; D. Caprette, written commun., 

1992). The velocity field presented here differs from that 

of Lisowski and others (199la) primarily in the inclusion 

of GPS and VLBI data. The trilateration and GPS networks are linked through common benchmarks (or ties 

between co-located trilateration and GPS benchmarks) at 

Loma Prieta, Eagle Rock, Allison, Hamilton, and Brush 2 

(fig. 2). The GPS baseline velocities constrain the rigid 

body rotation of the trilateration network, while the VLBI 

SST 

SST 

SST 

SST 

SST 

SST 

SST 

SST 

SST 

data constrain the translation of the network with respect 

to the North American plate. 

The velocities of two VLBI stations in the region (Fort 

Ord and Presidio, fig. 2) are shown relative to a North 

American reference frame defined by a fixed origin at the 

Westford, Massachusetts VLBI station and a fixed azimuth from Westford to the VLBI station at Richmond, 

Florida (Clark and others, 1990). Figure 6 shows the timedisplacement histories of the VLBI data. Pre-Lorna Prieta 

VLBI measurements at Fort Ord S, which is located near 

the GPS station Brush 2, did not begin until 1988. The 

pre-seismic velocity of Fort Ord S is therefore poorly determined. There are, however, five years of VLBI data at 

Fort Ord, which is located 8.9 km north of Fort Ord S and 

Brush 2. Within uncertainties the velocity of the two VLBI 

stations, Fort Ord and Fort Ord S, are indistinguishable. 

Thus, we compute the velocity field in figure 5 assuming 

that Fort Ord and Brush 2 moved at the same rates prior 

to the earthquake. 

The velocity field in the southern San Francisco Bay 

area exhibits 38 k 5 mmlyr of right-lateral shear across a 

120-km-wide zone, as described by Lisowski and others 

(199la). Figure 5 shows that the easternmost stations of 

the trilateration network are moving at about 1 cmlyr with 

respect to North America, indicating that most of the 4.8- 

cmlyr relative motion between the Pacific and North 

American plates (DeMets and others, 1987) is contained 

within the network. The velocity vectors are approximately 

parallel to the regional trend of the San Andreas fault 

system and do not appear to be influenced by the local 

left bend in the San Andreas fault trace through the Santa 
D228 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Cruz Mountains. The displacement field broadens towards 

the northwest. Whereas the creeping segment of the San 

Andreas fault southeast of San Juan Bautista accommodates 35 mmlyr of slip, deformation becomes distributed 

between the San Andreas, Hayward, and Calaveras faults 

in the San Francisco Bay area. Deviations from the simple 

shear pattern are apparent along the southern Calaveras 

fault in the closely spaced Hollister network. Stations just 

east of the Calaveras fault appear to move towards the 

fault at an oblique angle. Even though we removed 

coseismic line-length changes and postseismic transient? 

due to the 1979 Coyote Lake and the 1984 Morgan Hill 

earthquakes along the Calaveras fault, these deviations 

persist. It is possible that the complex series of creeping 

and locked fault segments along the Calaveras fault causes 

the more complicated displacement-rate vectors 

Figure 5.- Pre-Loma Prieta earthquake site velocities (solid arrows with 95-percent error ellipses) in the southern San 

Francisco Bay region determined from VLBI, trilateration, and GPS measurements. The GPS and trilateration station 

Brush 2 is assumed to move with the VLBI determined velocity of Fort Ord. All velocity vectors are shown relative to 

North America. Dashed arrows show vectors computed from a kinematic model of the principal San Francisco Bay area 

faults (fig. 7). 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D229 

(Oppenheimer and others, 1990). Alternatively, the southern Calaveras or the southern Sargent faults may accommodate convergence by oblique slip (Matsu'ura and others, 

1986). 

Pre-Loma Prieta displacement rates are available for 

only a few of the stations monitored after the earthquake, 

so that to compare pre- and post-earthquake velocities we 

must interpolate the pre-seismic velocity field between 

stations. We utilize a kinematic model of the preseismic 

velocity field using three-dimensional elastic dislocations 

(Erickson, 1987; Okada, 1985). Uniform-slip boundary 

elements in an elastic, homogenous, and isotropic half 

space represent the Bay area faults. Interseismic shear 

about a locked upper crustal strike-slip fault is approximated by slip on a buried fault plane below the seismic 

zone at a rate comparable to the average geologic sliprate (for example, Lisowski and others, 1991a). Surface 

creep on the Hayward, Calaveras, and the central San 

Andreas fault is modeled by shallow fault elements. Our 

objective here is not so much to find the correct kinematic 

m Fort Ord 

Station 

offset 

removed 

0.3 I,...... 

Fort Ord S 

r~ll~.ll~,lr~, 

: Earthquake 

offset 

: removed 

East 

Figure 6.- Time-displacement plots of pre- and post-Loma Prieta earthquake VLBI measurements in the San Francisco Bay area (Presidio, Fort 

Ord, and Fort Ord S), relative to a North American reference frame. The 

displacements of Fort Ord and Fort Ord S located 8.9 km to the south are 

shown on the same plot. Loma Prieta earthquake offsets and the offset 

introduced in the 1988 change of stations (Fort Ord to Fort Ord S) have 

model of San Francisco Bay area tectonics, as it is to 

develop a close-fitting approximation of the preseismic 

velocity field that can be used to compute the expected 

velocities at stations whose preseismic velocities are not 

available. 

We developed a best-fitting model using a linear inversion method (Du and others, 1992) that inverts trilateration 

and GPS data for right-lateral strike-slip magnitudes on 

the individual segments of the San Francisco Bay area 

faults shown in figure 7. All faults are assumed to be 

vertical with no dip-slip component. A priori constraints 

on the starting model come from geologically determined 

slip rates (closely based on a compilation of Lienkaemper 

and others, 1991). Our results (fig. 7) include (1) deep 

(>I5 km) slip on the Peninsular segment of the San Andrea~ fault at about 20-25 mmlyr (prior was 20k3 mmlyr), 

(2) deep slip (>I1 km) on the Hayward fault at about 14 

mmlyr (prior was 9&2 mdyr), (3) deep slip (>lo km) on 

the northern Calaveras fault at about 10-12 mmlyr (prior 

was 6k4 mmlyr) and at up to 28 mdyr (prior was 17k5 

been removed from the data (Clark and others, 1990). The line fits are 

based on the pre-earthquake data at Presidio and on the Fort Ord data 

between 1983 and early 1988. The preseismic and postseismic displacement rates at Presidio and at Fort OrdIFort Ord S are not significantly 

different. 
D230 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

mm/yr) on the southern Calaveras fault, and (4) shallow 

creep on the East Bay and southern San Andreas faults at 

rates comparable to the a priori values. The deep slip rate 

on the southern Calaveras fault appears too high and may 

be related to effects of the Coyote Lake and Morgan Hill 

earthquakes that occurred during the measurement period. 

Figure 5 shows a comparison of the measured and modeled station velocities. To facilitate the comparison, the 

model velocities are adjusted so that Loma Prieta moves 

at the observed velocity. We note that most velocities are 

fit well within their 95-percent uncertainties except for 

several of the sites near the southern Calaveras fault. The 

model also slightly underpredicts the velocities southwest 

of the San Andreas fault south of Loma Prieta. 

The kinematic model shown in figure 7 can be used to 

compute the predicted preseismic velocities of all the stations monitored after the earthquake (table 3). Figure 8 

shows the measured postseismic velocities in comparison 

with pre-earthquake velocities, either measured at these 

sites or derived from the kinematic model. 

POST-LOMA PRIETA EARTHQUAKE 

DISPLACEMENT FIELD 

Figure 3 shows the changes with time of the coordinates of the stations in the Black Mountain profile relative to station Bend. The least-squares estimates of the 

relative station velocities are shown as solid lines. Within 

the measurement errors we see no evidence for changes in 

velocity during the surveyed 3-year period following the 

earthquake. The average root-mean-square (rms) residuals about the best fit line are 6.0 mm and 7.1 mm in the 

north and east, respectively. These residuals are slightly 

larger than the short-term precision estimates, presumably 

because of the more poorly determined broadcast orbit 

solutions and the fact that some error sources (for example, atmospheric and orbital) are correlated in the dayto-day comparisons (Davis and others, 1989). The 

estimates of the vertical displacement component do not 

show a trend above the data noise; the average rms residual about the mean is 3.3 cm. 

A tie between the two profiles is provided by four measurements of the Bend-to-LPl baseline (fig. 34. Timedisplacement plots of stations in the Loma Prieta profile 

relative to station LP1 are shown in figure 4. We note that 

several sites show a significant convergence normal to the 

San Andreas fault. There is some indication that faultnormal rates decayed during the 2-year period (Savage 

and others, 1994) at some sites (for example, figs. 4G and 

H). However, we assume constant postseismic velocities 

in our analysis. The solution for the baseline to Brush 2 

shows a large scatter in the east component (fig. 40), 

resulting in a weak tie to the VLBI reference frame. 

Figure 8 shows the postseismic velocity field in the 

region based on (1) nine occupations of the Black Mountain profile between January 1990 and January 1993, (2) 

nine occupations of the Loma Prieta profile between October 1989 and December 1991 (Savage and others, 1994), 

(3) 12-14 occupations of the San Francisco Bay area monitor network between October 1989 and December 1991, 

(4) four occupations of ties between LP1 and Bend, and 

locked 4 12 20 28 36 mmfyr 

Figure 7.- Slip rates on segments of the San Francisco Bay area fault system used to model the preseismic velocity field. 

The fault elements that end at 30 km depth in the plot continue to 1000 km depth in the model. 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D23 1 

Table 3.- Preseisrnic and postseismic relative station velocities 

[The pre-earthquake horizontal velocities were computed from a kinematic model of the preseisrnic deformation (fig. 7). The postseismic 

velocities and their uncertainties (1 sigma) were determined from a least-squares fit to the GPS data and their full-position covariances] 

Baseline Preseismic (cmlyr) Postseismic (crn/yr) 

From To North East North +I- East +I- Up +/- 

GPS 

Bend - Pigeon Point 

Bend - Halmaur 

Bend- Haul 

Bend - True 

Bend- BM10 

Bend- PAWT 

Bend - Foothill 

Bend - Bay shore RM 1 

Bend - Antelynx 

Bend- LP1 

LP1 - Cliff 

LP1 - Gregor 

LP1 - Fire 

LP1- LP2 

LP1- LP4 

LP1 - End 

LP1 - Mazzone 

LP1 - Calero 

LP1- Coy 

LP1 - Hamilton 

LP1 - Mocha 

LP1 - Mt. Oso 

LP 1 - Eagle Rock 

LP1 - Allison 

LP1 - Brush 2 

VLBI 

Westford Presidio 

Westford Fort Ord 

Westford Fort Ord S 0.18 1 1.36 

(5) Very Long Baseline Interferometry (VLBI) data collected at Fort Ord S and at Presidio between October 

1989 and July 1991 (fig. 6, D. Caprette, written commun., 

1992). The GPS station velocities were determined using 

a least squares algorithm that computes station velocities 

and covariances relative to a local reference station from 

the station coordinates and the associated covariance matrices of the individual campaigns (M. Murray, written 

commun., 1991). Formal errors from the Bernese GPS 

solutions were scaled by a factor of five, as they appear to 

underestimate the observed scatter of the residuals about 

the solution by about that factor (M. Murray, oral 

commun., 1992). Table 3 lists the computed relative GPS 

station velocities and their 1-sigma uncertainties. 

The velocities of the VLBI stations Fort Ord S and 

Presidio are shown relative to fixed North American sites 

(fig. 8, Clark and others, 1990). Because the tie to the 

VLBI network through Brush 2 is noisy (fig. 40), we 

present the velocities of the GPS sites holding the velocity vector at Mt. Oso, located about 60 km northeast of 

Loma Prieta, at its pre-earthquake rate (fig. 8). 

Several features in the postseismic velocity field (fig. 

8) are worthy of note. Stations east of the Calaveras fault 

as well as those in the Black Mountain profile have velocities consistent with their pre-earthquake velocities. 

Many of the Black Mountain profile sites have moved 

slightly more northeasterly since the earthquake (fig. 8). 

This may be caused by the weak tie between Bend and 

LP1, which is based on only four surveys. Over the 60- 

km-long Black Mountain profile, from the Pacific coast 

(Pigeon Pt.) to just east of the Hayward fault (Antelynx), 

right-lateral shear at 22k5 mmlyr occurred since the earth- 
D232 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

quake. The postseismic displacement rates do not differ liquely towards the fault at an accelerated rate. The most 

significantly from the pre-earthquake rate of 20k4 mm significant difference between the pre- and post-earthquake 

over approximately the same region. velocity fields is the fault-normal contraction centered 

In contrast, some stations in the Loma Prieta profile northeast of Loma Prieta mountain. Stations Fire to Coyshow dramatic changes in velocity after the earthquake ote show significant fault-normal motion (see also fig. 4) 

(Lisowski and others, 1991b; Savage and others, 1994). that would not have been anticipated from the pre-earthStations near Santa Cruz and at Eagle Rock moved ob- quake data. 

Figure 8.- Postseismic velocity field determined from GPS measurements. The velocities (solid arrows) and associated 

errors were computed relative to LP1. A rigid body translation is then added so that Mt. Oso displaces at its pre-Loma Prieta 

velocity. Postseismic velocities relative to stable North America are shown for the two VLBI stations in the area. The 

discrepancy between the GPS-derived vector at Brush 2 and the VLBI vector at Fort Ord S, which is not significant at 95- 

percent confidence, is likely due to the poor results in the Brush 2 - LP1 baseline (fig. 40). Dashed arrows show the 

preseismic velocities, either measured with error ellipses or computed from the kinematic model (fig. 7), without error 

ellipses. 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D233 

There is a large difference between the postseismic velocity vectors at the GPS station Brush 2 and the VLBI 

station at Fort Ord S, less than 100 m away. The discrepancy, which is not significant at the 95-percent confidence level, is almost certainly associated with errors in 

the GPS baseline determinations of LP1 to Brush 2. The 

large scatter in the east component of Brush 2 in the timedisplacement plots (fig. 40) may be caused by incorrectly 

resolved integer phase ambiguities. 

Whereas only minor changes have been observed in the 

Black Mountain profile 44 krn northwest of the 1989 epicenter, measurements along the Loma Prieta profile show 

significant changes in displacement rates from pre-earthquake values. The difference between the post- and preearthquake velocities (fig. 9) represents the component of 

the measured signal that exceeds the secular deformation 

field. The displacement rates predicted by the preseismic 

model (fig. 7) were subtracted from the postseismic rates, 

assuming that the motion of Mt. Oso did not change after 

the Loma Prieta earthquake. The error ellipses were determined under the assumption that the errors in the preand post-earthquake velocities are equal. 

DATA INTERPRETATION AND MODEL 

We model the residual site velocities with rectangular 

uniform-slip displacement discontinuities embedded in a 

homogeneous elastic half-space (Okada, 1985). Using a 

forward-modeling approach, we conclude that two distinct fault planes are required to fit the data. A single fault 

plane, at any depth, does not produce the observed faultFigure 9.- Excess postseismic station velocities (secular displacement fault in the Foothills thrust system. The two shaded rectangles show the 

rates subtracted from postseismic rates) assuming the velocity of Mt. projections of the faults used to compute the modeled station displaceOso was unaffected by the earthquake. Error ellipses for the sites near ments. The first fault lies in the Loma Prieta aftershock zone, from 6 to 

Lorna Prieta are not shown for clarity. They are the same order of 17 km depth, dipping 70' SW, with a strike-slip displacement rate of 

magnitude as the error ellipses at other sites. Also shown are the veloci- 0.10 dyr. The second fault slips 0.12 dyr on a reverse fault dipping 

ties (filled arrows) computed from a dislocation model involving slip on 47' SW, striking subparallel to the San Andreas fault, from 1.4 to 5.8 

two faults; strike slip on the Loma Prieta fault and thrusting on a reverse km depth. 
D234 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

normal and fault-parallel displacements at the surface. As 

an example, we show the predicted displacement fields 

from uniform slip on a 5-km-wide, down-dip extension of 

the Loma Prieta rupture. Figure 10 compares the measured and predicted excess velocities computed from models involving strike slip, dip slip, and oblique slip. The 

magnitude of the slip is adjusted to provide the best fit (in 

a least-squares sense) to the residual displacement rates. 

The models involving a deep source do not fit the data 

along the Loma Prieta profile well and predict a significant acceleration along the Black Mountain profile that is 

not observed in the data. Table 4 lists the model parameters and misfits to these models. 

We use a quasi-Newton algorithm (Arnadottir, 1993; 

Arnadottir and Segall, 1994) to estimate the geometry and 

position of the dislocation that best fits the data. The algorithm allows upper and lower bounds on all of the dislocation parameters. The best-fitting one-fault model with 

the strike and dip of the Loma Prieta rupture is located at 

a depth of 13 to 15 km (table 4). This suggests that transient slip may have occurred within rather than below the 

rupture plane. However, the misfit, (dm, table 4), is 

2.3, which suggests that the model is inadequate. More 

importantly, the spatial distribution of residuals demonstrates that even the best-fitting one-fault model does not 

provide an adequate fit to the data. 

We used the quasi-Newton method to investigate the 

possibility that the data could be satisfactorily fit with 

multiple faults. In this calculation, one dislocation is constrained to have the strike and dip of the Loma Prieta 

rupture, while the other dislocation is not constrained. We 

found a better fit to the data with two slipping faults 

(table 4). One dislocation is roughly coincident with the 

Lorna Prieta coseismic rupture, the second is located northeast of the San Andreas fault near the Foothills thrust 

belt. The surface projections of the two model slip planes 

and the predicted velocity vectors are shown in figure 9. 

The best-fit model, ( ,,/RSS/df= 1.8), includes a combination of strike-slip at a rate of 0.10 dyr, from 6 to 17 km 

depth on the Loma Prieta fault and reverse slip at a rate of 

0.12 m/yr on a 47' southwest-dipping thrust plane at shallow depth (1.4 to 5.8 km) northeast of the San Andreas. 

The velocities predicted by the model are shown in figure 

9. While the model by no means provides a perfect fit, it 

does explain both the accelerated fault-parallel motion and 

the fault-normal convergence. 

The depth of postseismic slip on the Loma Prieta fault 

is primarily constrained by the narrowness of the zone of 

accelerated deformation and the absence of increased displacement rates on the Black Mountain profile. There is 

some trade-off between the down-dip fault width and the 

amount of slip on the faults. For example, models with 

somewhat wider faults and lower slip rates would produce similar surface displacements. It is possible that limited aseismic slip or ductile flow occurred below the 1989 

rupture zone, but that this was masked by shallower deformation sources. Our inversion results indicate pure 

strike-slip motion on the Loma Prieta rupture plane since 

the earthquake. We believe that significant dip slip has 

not occurred on the Loma Prieta fault as any reverse component would cause the sites southwest of the San Andreas to move oceanward (fig. 105 and C), opposite to 

the observed trend. Finally, the location, strike, and dipparallel rake of the thrust fault is constrained by the nearly 

San Andreas fault-perpendicular contraction northeast of 

the fault. 

DISCUSSION 

ALTERNATIVE MODELS 

The postseismic displacement pattern following the 

Loma Prieta earthquake has also been interpreted by Linker 

and Rice (1991, and this chapter) and by Lisowski and 

others (1991b). Linker and Rice (this chapter) model the 

displacement field resulting from the interaction of the 

earthquake rupture with a linear viscoelastic zone that 

represents the deep aseismic portion of the fault. They 

also developed a non-linear model that assumes rate- and 

state-dependent friction on the down-dip aseismic region 

of the fault zone. Their models, involving a relatively 

deep deformation source below the coseismic rupture, predict the fault-parallel velocities of the Loma Prieta profile 

well, but do not predict the observed fault-normal contraction and appear to overpredict the rates in the Black 

Mountain profile. Figure 10D shows the displacements 

predicted by model 1 of Linker and Rice (this chapter, 

figure 6A) for comparison. As with the deep-fault oblique-slip model in figure 10C, their relaxation model is 

not capable of producing fault-trace normal contraction. 

Both deep models also overpredict displacement rates in 

the Black Mountain profile. On the other hand, some of 

the increased fault-parallel shear observed in the Loma 

Prieta profile may be explained by the transfer of stress to 

a ductile region below the fault zone, as proposed by 

Linker and Rice (this chapter). 

The model proposed by Lisowski and others (1991b) 

involves strike-slip shear on the Loma Prieta rupture plane 

(0.83 m/yr at 4 to 11 km depth), as well as fault-normal 

compaction of the fault zone at a rate of 0.11 mlyr (3 to 7 

km depth). The compaction is required to explain the faultnormal site-velocity vectors. This model successfully predicts the horizontal displacement measurements. Lisowski 

and others (199 1 b), however, offer no satisfactory physical model of the proposed fault-zone collapse at decimeter rates. 

Our model involves contemporaneous pure strike-slip 

and thrust faulting at high rates on two faults with similar 

strikes. The inferred slip on the Loma Prieta fault may 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D235 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

represent creep around high-slip asperities that released 

most of the earthquake moment (Beroza, 1991; Hartzell 

and others, 1991; Steidl and others, 1991; Wald and others, 1991; Arnadottir and Segall, 1994). We do not have a 

unique explanation for the dominance of the strike-slip 

component on the Loma Prieta fault. One possibility is 

that postseismic fault creep may be concentrated around 

the southern half of the rupture that underwent mostly 

strike slip during the earthquake (Beroza, 1991; Arnadottir 

and Segall, 1994). A large postseismic strain anomaly, 

measured with a borehole tensor strain meter near the San 

Andreas fault about 30 km southeast of the Loma Prieta 

epicenter, may be further evidence of continued aseismic 

slip in this region (Gwyther and others, 1992). 

THRUSTING NORTHEAST OF THE SAN 

ANDREAS FAULT 

Northeast of the San Andreas fault lie several sub-parallel southwest-dipping faults of the Foothills thrust belt 

(fig. 1, and Aydin and Page, 1984), with a dominant component of thrust motion (McLaughlin, 1990). Some of 

these faults appear to have been active in the Quaternary, 

as evidenced by offset young alluvial deposits and soil 

horizons (Haugerud and Ellen, 1990; McLaughlin, 1990). 

Broad zones of contractional surface deformation during 

the Loma Prieta earthquake follow the trend of these fault 

zones intermittently over a total distance of about 20 krn 

(Haugerud and Ellen, 1990). Our model thrust fault appears to coincide with the Foothills thrust system. 

A cross section perpendicular to the San Andreas fault 

shows the mapped faults, the post-earthquake seismicity, 

and the model faults found from the inversion of the excess postseismic velocities (fig. 11). The modeled reverse 

fault lies close to the Berrocal fault zone and has a comparable dip (McLaughlin, 1990). This zone has experienced strongly increased earthquake activity since the 

Loma Prieta event (Reasenberg and Simpson, 1992). Focal mechanisms of all aftershocks greater than M=2.0 along 

the Foothills thrust belt indicate that these events occurred 

predominantly on reverse faults (fig. 115). The seismicity 

near the Loma Prieta rupture, on the other hand, is complex and includes all types of fault plane solutions 

(Oppenheimer, 1990). Note that the fault slip we model 

exceeds by several orders of magnitude the estimated slip 

accumulated in aftershocks and must be dominantly 

aseismic. 

Stress redistribution accompanying the Loma Prieta main 

shock increased the loading on thrust faults northeast and 

updip of the rupture (Reasenberg and Simpson, 1992). 

Figure 12 shows a cross section perpendicular to the San 

Andreas fault, with the contoured coseismic stress changes 

resolved on planes striking N60Â° and dipping 45' SW 

overlaid on a plot of the aftershocks. Shear and normal 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D237 

stress changes both affect the tendency for slip through a 

Coulomb Failure Function (Reasenberg and Simpson, 

1992) 

CFF = T~ + PO,, 

where T~ is the shear traction in the slip direction, on the 

fault-normal traction, and p, is the coefficient of friction. 

A coefficient of friction of 0.6 has been chosen for the 

calculation in figure 12. The slip distribution during the 

earthquake is important in determining the induced stress 

field in the near field of the earthquake rupture. In this 

calculation we used the heterogeneous fault slip distribution of Beroza (1991). We found that co-seismically induced static stress changes of 1 to 2 MPa that enhance 

thrust loading occurred northeast of the San Andreas fault. 

The stress changes on fault planes of somewhat different 

strikes and dips are comparable to those shown in figure 

12. We conclude that the increased micro-seismicity and 

the aseismic thrusting inferred from geodetic data reflect 

deformation in response to coseismic stress changes. Aftershock seismicity near the Foothills thrust belt is restricted to an approximately 15-km-wide zone along the 

northern half of the Lorna Prieta rupture. This may be 

related to the observation that most of the thrust component of the earthquake slip is concentrated northwest of 

the Lorna Prieta epicenter (Beroza, 1991; Steidl and others, 1991; Marshall and others, 1991; Arnadottir and 

Segall, 1993). 

IMPLICATIONS FOR SEISMIC HAZARD 

IN THE SAN FRANCISCO BAY REGION 

Continued slip at decimeter rates on faults in the San 

Francisco Bay region will change the loading of nearby 

segments of the San Andreas fault and the Hayward fault. 

We modeled the change ',in loading on northwest-striking 

fault planes due to the Loma Prieta earthquake and modeled postseismic fault slip with the method outlined above, 

assuming that fault-parallel shear stress as well as faultnormal stresses determine the fault response through a 

Coulomb failure criterion. A detailed analysis of coseismic 

static stress changes on San Francisco Bay area faults 

shows that micro-seismicity rates increased or decreased 

in accordance with the computed stress changes 

(Reasenberg and Simpson, 1992). Reduced creep rates, 

observed along the southern Hayward fault following the 

earthquake, are evidence of induced left-lateral shear stress 

across that segment of the fault (Lienkaemper and others, 

1 992). 

Figure 13 shows the contoured changes in frictional 

failure conditions for right-lateral strike-slip shear on northwest-striking fault planes at a depth of 10 km due to the 

coseismic rupture (fig. 13A), and from the predicted yearly 

postseismic fault slip (fig. 135). Secular tectonic loading 

is not included, and the modeled postseismic stresses indicate a relaxation or enhancement of fault loading above 

the background. The results are only directly applicable 

to N45OW-striking fault planes. However, deviations of 

fault strike by up to 10' do not change the pattern significantly. We find that the postseismic fault slip further increases the load on the San Andreas fault immediately 

northwest of the Loma Prieta rupture, while the combined 

coseismic and postseismic effects have a retarding effect 

on the Hayward fault zone and on most of the Calaveras 

and San Gregorio-Hosgri fault zones. The yearly 

postseismic stress changes are about an order of magnitude less than those induced coseismically. If the Loma 

Prieta earthquake stress changes advanced the occurrence 

of the next earthquake north of the rupture by 2-25 years 

(Reasenberg and Simpson, 1992), then the postseismic 

changes enhance this effect by about 10 percent per year. 

Similar postseismic adjustments occurred following the 

1979 Coyote Lake and the 1984 Morgan Hill earthquakes 

along the Calaveras fault (Oppenheimer and others, 1990). 

While the coseismic stress perturbations on neighboring 

fault segments induced by these moderate earthquakes were 

small (Du and Aydin, 1990), postseismic fault creep adjacent to the coseismic ruptures may have caused the observed south-to-north propagation of earthquakes along 

the Calaveras fault (Oppenheimer and others, 1990). 

The stress changes caused by coseismic and postseismic 

fault slip only add to stresses accumulated by background 

tectonic loading of the San Francisco Bay area faults and 

to stress changes produced by previous earthquakes (for 

example, the 1906 San Francisco earthquake). If a fault 

segment is near instability, the stress release associated 

with the Loma Prieta earthquake may only slightly retard 

the imminent failure. The January 15, 1993, Gilroy earthquake that occurred in a region of induced left-lateral 

shear east of the southern termination of the Loma Prieta 

rupture is a timely reminder of this. The Peninsular segment of the San Andreas fault, the updip extension of the 

Loma Prieta rupture (from about 8 km to the surface), and 

several thrust faults that parallel the San Andreas fault 

through most of the San Francisco Bay region appear to 

represent significant seismic hazards in the wake of the 

Loma Prieta earthquake. Continued reverse faulting may 

also relieve normal traction along the southern Hayward 

fault and increase the likelihood of faulting there, despite 

a decrease in right-lateral stress. 

Rapid fault slip along a fault of the Foothills thrust belt 

since the Loma Prieta earthquake emphasizes the potential earthquake hazards associated with these faults. While 

records of historic seismicity in the area reach only as far 

back as the early 19th century, they show that the last 80 

years have been marked by unusual seismic quiescence. 

In 1865, a M=6.5 earthquake may have ruptured a thrust 
San 

Andreas Sargent Berrocal 

fault zone Zayante fault zone fault 

fault 

Shannon ". rn n . w . -6.. 

fault fault zone 

Berrocal fault 

I-. 

MAGNITUDES 

0.0 + 

3.0 + 

4.0 + 

5.0 + llllll~l 

na Pr 

fault 

Figure 11.- A, Southwest-to-northeast cross section (Gregor to Coy in fig- and include all well-located events from one month after the Loma Prieta 

ure 2) of the geology (modified from McLaughlin, 1990) and Loma Prieta earthquake until December 1992. B, Map view of post-earthquake microaftershocks. Also shown are the two model faults. Note the strike slip on the seismicity and focal plane solutions of earthquakes greater than M = 2 northLoma Prieta fault and the dip slip on the Berrocal fault. The plotted earth- east of the San Andreas. 

quakes are from a 15-km-wide zone centered on the line of section (box in B) 
Figure 1 1 .- Continued 
10 20 

DISTANCE (KM) 

MAGNITUDES 

O.O+ 

1.0+ 

2.0+ 

3.0+ 

lllllllll 

Figure 12.- Contoured stress changes (in MPa) induced by the Loma Prieta earth- homogeneous halfspace. Both fault-normal stresses and dip-parallel shear stresses 

quake on 45' SW dipping fault planes. Dashed contours indicate decreases in CFF, effect the loading of the fault through a Coulomb failure criterion assuming a coeffisolid contours increases. The coseismic displacement distribution determined by Beroza cient of friction p = 0.6. 

(1991) was discretized into 272 uniform slip fault elements in an elastic, isotropic, and 
POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D241 

fault northeast of the San Andreas fault with two M=5 

foreshocks in the preceding year (Tuttle and Sykes, 1992). 

This earthquake was preceded by a M=7 break in 1838 

that involved the Peninsular San Andreas fault segment 

and maybe the Loma Prieta fault segment and followed in 

1868 by a large earthquake on the southern Hayward fault 

(for example, Tuttle and Sykes, 1992). An earthquake comparable to the 1865 event in the southern Santa Cruz Mountains constitutes a significant hazard to the south San 

Francisco Bay region. 

ACKNOWLEDGMENTS 

We thank Dave Oppenheimer and Greg Beroza for aftershock data and Doug Caprette for providing the VLBI 

data. Thora Arnadottir, Yijun Du, Jeff Freymueller, and 

Mark Murray wrote very useful programs used in our 

analyses. A large number of Stanford students and faculty 

helped with the GPS surveys of the Black Mountain profile. We benefited from discussions with Mark Linker. 

Ruth Harris and Mark Murray provided very helpful reviews. We acknowledge support from NSF grants EAR9003575, EAR-9002 164, and EAR-9 1 16 1 17, and from the 

U.S. Geological Survey. 

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Figure 13.- A, Horizontal stress changes (in kPa) at 10 km depth on Andreas fault, and the central Calaveras fault experienced increased 

N45OW-striking strike-slip faults caused by the Loma Prieta earthquake. right-lateral stress (positive), whereas the San Gregorio-Hosgri and the 

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POSTSEISMIC STRAIN FOLLOWING THE LOMA PRIETA EARTHQUAKE FROM REPEATED GPS MEASUREMENTS D243 

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p. 1199-1202. 

Oppenheimer, D.H., Bakun, W.H., and Lindh, A.G., 1990, Slip partitioning of the Calaveras fault, California, and prospects for future 

30' 20' 10' 122 O 50' 40' 30' 

Figure 13.- Continued 
D244 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

earthquakes: Journal of Geophysical Research, v. 95, p. 8483- 

8498. 

Reasenberg, P.A., and Simpson, R.W., 1992, Response of regional 

seismicity to the static stress change produced by the Loma Prieta 

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Rothacher, M., Beutler, G., Gurtner, W., Schildknecht, T., and Wild, 

U., 1990, Documentation for Bernese GPS Software Version 3.2: 

Bern, University of Bern. 

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Astronomical Society, v. 49, p. 575-591. 

Rymer, M.J., 1990, Near-fault measurement of postseismic slip associated with the 1989 Loma Prieta, California, earthquake: Geophysical Research Letters, v. 17, p. 1789-1792. 

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Savage, J.C., and Prescott, W.H., 1978, Asthenosphere readjustment 

and the earthquake cycle: Journal of Geophysical Research, v. 83, 

p. 3369-3376. 

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earthquake: Journal of Geophysical Research, v. 99, p. 13757- 

13765. 

Scholz, C.H., 1977, A physical interpretation of the Haicheng earthquake prediction: Nature, v. 267, p. 121-124. 

Steidl, J.H., Archuleta, R.J., and Hartzell, S.H., 1991, Rupture history 

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San Andreas fault: Journal of Geophysical Reseach, v. 88, p. 

5893-5902. 

1986, Cyclic deformation related to great earthquakes at plate 

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272. 

Toksoz, M.N., Shakal, A.F., and Michael, A.J., 1979, Space-time migration of earthquakes along the North Anatolian fault zone and 

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THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

SHALLOW, POSTSEISMIC SLIP ON THE SAN ANDREAS FAULT AT 

THE NORTHWESTERN END OF THE LOMA PRIETA EARTHQUAKE 

RUPTURE ZONE 

CONTENTS 

By John Langbein, 

U.S. Geological Survey 

ABSTRACT 

A small, 10 krn by 10 km geodetic network spanning 

the San Andreas fault was measured 7, 77, 157, 194, 422, 

547, and 756 days following the Loma Prieta earthquake. 

This network is located at the northwestern end of the 

rupture plane defined by the locations of numerous aftershocks. In the initial 70-day interval, the measured linelength changes revealed that 4.7k0.6 of right-lateral slip 

occurred within the network. However, during the later 2- 

year interval only 2.150.5 mmlyr of right-lateral slip could 

be detected. Thus, it appears that the measured slip is a 

typical response of a fault following a major shock in that 

the rate of slip decreases rapidly with time. However, the 

magnitude of the postseismic slip is less than 0.5 percent 

of the inferred co-seismic slip at depth. In addition, we 

estimate that secular strain is accumulating at 0.1950.05 

ppdyr of fault-parallel strain, -0.3750.06 ppdyr of faultnormal strain, and 0.2Ok0.05 ppmlyr of fault-tensor-shear 

strain which appears to be caused by postseismic, dextral 

slip on the rupture plane of the Loma Prieta earthquake. 

INTRODUCTION 

Immediately following the Loma Prieta earthquake, U.S. 

Geological Survey personnel installed and measured a 

small geodetic network on the northwestern end of a zone 

defined by the epicenters of aftershocks that occurred 

within 2 days of the main shock (fig. 1). The intent of 

these observations is to detect possible fault slip following the main shock and to monitor the deformation rate 

for anomalous changes that might precede a large and 

potentially damaging earthquake on the San Francisco peninsula segment of the San Andreas fault northwest of the 

Loma Prieta shock. The network was set up to the northwest rather than to the southeast for several reasons. With 

only a short segment between the southern end of the 

rupture zone and the creeping section of the San Andreas 

fault south of San Juan Bautista (Burford and Harsh, 1980), 

the potential for an earthquake larger than M=6 is low. 

Furthermore, there already were several geodimeter 

baselines (Lisowski and Prescott, 198 1; King and others, 

1981; Savage and others, 1979) and volumetric strain 

meters (Johnston and others, 1986; Johnston and others, 

1987) for detecting deformation on the segment to the 

southeast. Within the zone defined by the aftershocks, 

there was already geodetic coverage because several 

baselines are monitored using a central point at Loma 

Prieta (Savage and others, 1987). However, geodetic coverage just to the northwest of the aftershock zone has 

been poor because of dense tree growth and steep topography. Nonetheless, we were able to install a few monuments that could be measured with a two-color geodimeter. 

The advantage of this instrument over a global positioning system (GPS) (Prescott and others, 1989a) and the 

more traditional terrestrial distance measuring instrument 

(Savage and Prescott, 1973) is its higher precision of 0.10 

to 0.15 part per million for lengths between 2 and 10 km 

(Langbein, 1989), which makes it possible to detect small 

displacements and possible changes in rates. The data to 

be discussed here indicate displacements of 4 mm in the 

2.5-month period following the main shock, and these 

displacements can be interpreted as 5 to 6 mm of shallow, 

cumulative fault slip on this segment of the San Andreas 

fault between Black Mountain and Loma Prieta. However, the measurements during the next 2 years detected a 

similar amount of displacements, indicating a substantial 

reduction in rate. Although other geodetic measurements 

D245 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 
POSTSEISMIC SLIP ON THE SAN ANDREAS FAULT AT THE NORTHWESTERN END OF RUPTURE ZONE D247 

have detected significant deformation following the main Prescott, 1978), which may have a time scale of years 

shock (Lisowski and others, 1991; Burgmann, and others, rather than days, as is often found with shallow slip. 

1994), these measurements have not been able to unambiguously detect any decrease in rate that is usually characteristic of postseismic slip because of the lower precision THE DATA 

of these techniques (Savage and Prescott, 1973; Prescott 

and others, 1989b). It also is possible that these large, Within one week following the Loma Prieta earthquake, 

100-lun networks are more sensitive to slip at depth caused we installed and measured the lengths of the four baselines 

by visco-elastic relaxation of the lower crust (Savage and shown in figures 1 and 2. During the first week of JanuDEPTHS .I SARA ' TABLE / 

A 20.0+ 

MAGNITUDE 

Figure 2.-Detailed map showing the locations of the six baselines of shocks (Dietz and Ellsworth, this chapter) near and within the network 

the El Sereno network. The four baselines shown with solid lines were for the period spanning the surveys of the network. Most of the shocks 

measured starting 7 days after the Loma Prieta earthquake. The other occurred within a month of the main shock. Again, only earthquakes 

two baselines, shown with dashed lines, have measurements starting 7 with magnitudes greater than 2 are shown. 

days after the main shock. Also plotted are the locations of the after- 
D248 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

ary 1990, these baselines were surveyed again, and two 

additional baselines were measured for the first time. In 

mid-March 1990, three out of the original four baselines 

were measured for the third time. Finally, in late April 

1990, all six baselines were measured again. The three 

most recent surveys included measurements on all six 

baselines with the exception of the December 1990 survey, when the monument at GAP was not recovered. Because of difficult line-of-sight conditions, we needed to 

use two central monuments, BOHL and BAY, on El Sereno 

Ridge so that we could have a 360' view of the surroundings. Two of the six baselines cross the surface trace of 

the San Andreas fault at about 45' angles, and their length 

changes are sensitive to fault slip. A third baseline, from 

BAY to VASONA, does not cross the San Andreas fault 

but is oriented at 45' to the fault strike and provides a 

sensitive measure of off-fault shear strain. A fourth 

baseline crosses the San Andreas fault at a steep angle 

and is useful for detecting displacements perpendicular to 

the fault. The last two baselines, from BAY to SARA and 

TABLE, are useful for detecting off-fault deformation. 

The instrument used to measure the baseline lengths is 

a portable, two-color geodimeter (Slater and Huggett, 

1976), with a nominal precision between 0.4 mm and 0.9 

mm for lengths between 2 and 7 km (Langbein, 1989). 

For distances of less than 10 krn, this instrument has at 

least a factor of three better precision than any other geodetic techniques (Savage and Prescott, 1973; Prescott and 

others, 1989a). Because the instrument ranges on two 

wavelengths, it is able to detect differences in travel time 

due to dispersion in the atmosphere. Whereas the use of 

two optical wavelengths gives the instrument its high precision, it also limits the maximum range to that obtained 

with its long wavelength, blue. The scatter of the blue 

light during hazy conditions limits the use of the instrument to clear nights. The March 1990 survey was incomplete due to foggy conditions. 

The results of measuring the six baselines are shown in 

figure 3 and clearly show that two of the fault-crossing 

baselines changed length by 4 mm between the first two 

surveys and show some changes in length during the next 

2 years. The sense of the initial set of displacements is 

consistent with approximately 5 to 6 mm of right-lateral 

slip on the San Andreas fault, or localized displacement 

of the monument at BOHL to the southeast by the same 

amount. While the baselines to SUMMIT and GAP showed 

significant displacements during the initial 2.5 months, 

the other two baselines did not show any length changes 

within their one standard deviation error bars. With any 

instrument that measures distance, one needs to be careful 

about possible drift in the instrumentation which could 

manifest itself as a length change. Typically, the drift 

would cause a change in the length scale, thus causing all 

of the baselines to either extend or contract in length. 

However, since the lengths of two baselines remain unchanged and the other two baselines show displacements 

of equal magnitude but in the opposite sense, a change of 

the length scale in the instrumentation is unlikely. 

The 2- to 4-mm amplitude of displacement on the 

baseline BOHL-CENTRAL has a suspicious seasonal periodicity that has been seen elsewhere in central California (Langbein and others, 1990). The monument at 

CENTRAL is placed at the edge of a steep embankment 

next to a highway. Any soil dilation corresponding to the 

seasonal rainfall would contaminate the data from this 

baseline. Accordingly, the data from this baseline will be 

ignored in the following discussion. 

DISCUSSION 

Although a localized displacement vector to the southeast of the monument BOHL is consistent with the observations, the simplest tectonic model is 4.7k0.6 mm of 

shallow, right-lateral slip on the segment of the San Andrea~ fault located within the network during the 2.5 

months following the main shock (fig. 4). Because the 

network is limited spatially, we cannot place any definitive bounds on the extent of the slip plane. Accordingly, 

we use the simplest model of a throughgoing fault plane 

that splits the crust into two rigid blocks. The results of 

estimating right-lateral slip as a function of time on the 

basis of this simple model are shown in figure 4. However, in a few tests of varying the top and bottom edges of 

the dislocation surface, we could reject with a 99 percent 

confidence level those models that specify zero slip between the surface and 3 km depth. All other models with 

slip on the San Andreas fault provide an adequate fit to 

the observations, and it seems likely that slip did occur at 

shallow depths. 

The inferred value of slip for late 1990 appears anomalous. However, confidence for that particular value of slip 

is low because only one out of the two fault-crossing 

baselines was measured during that survey. The baseline 

to GAP was not measured. Accordingly, in the following 

discussion, the slip value from late 1990 is ignored. 

The observed 4.7 mm of fault slip should be interpreted 

as postseismic slip, which often occurs after large earthquakes (Smith and Wyss, 1968; Langbein and others, 

1983). As shown in figure 2, there are a number of aftershocks located within the El Sereno network. Just to the 

north of the network, the number of aftershocks decreases 

significantly. The cumulative moment of these tiny aftershocks would translate into slip much smaller than the 

inferred 4.7 mm. The presence of seismicity is evidence 

for deformation, but most of the deformation could be 

happening aseismically. 

While estimating the slip as a function of time, the 

three components of tensor strain rate are fit simultaneously 

to the data using the method discussed by Langbein (1989). 
POSTSEISMIC SLIP ON THE SAN ANDREAS FAULT AT THE NORTHWESTERN END OF RUPTURE ZONE D249 

The results of this simultaneous adjustment for both slip 

and strain rate indicate 0.19k0.05 ppdyr of fault-parallel 

strain, -0.37k0.06 ppmlyr of fault-normal strain, and 

0.20k0.05 ppdyr of tensor shear. The local strike of the 

fault is taken to be N54OW. Recall, however, that data 

from the five baselines are used to estimate four parameters, so these estimates are statistically co-dependent. 

There are three sets of geodetic measurements that have 

detected postseismic deformation. Two sets of these measurements use GPS to re-occupy a profile of monuments 

oriented perpendicular to the San Andreas fault and spanning a 50- to 80-h-wide zone. The profile of Lisowski 

and others (1991) spans the epicentral area of the Loma 

Prieta earthquake and the second profile of Burgmann 

and Segall (1991) is located approximately 10 km northwest of the two-color network. Lisowski and others' (1991) 

measurements show 8 mmlyr near surface slip, an additional fault-parallel 20-mdyr displacement distributed 

over a 20-km-wide zone centered on the San Andreas 

fault, and 18 mdyr contraction over 15-km-wide zone 

for which the San Andreas forms the southwestern boundary. These displacement rates translate into a 0.5 ppmlyr 

shear strain and 1.2 ppdyr fault-normal contraction. By 

comparison, the two-color network has detected lower rates 

by a factor of two to three than the GPS observations 

spanning the epicentral area. However, to the northwest 

of the two-color network the data from the GPS network 

of Burgmann and Segall (1991) show rates that can be 

explained by secular strain accumulation. Thus, the strain 

rates derived from the two-color data fit between the high 

rate within the epicentral area and the low rate to the 

northwest. 

Burgmann and others (this volume) have modeled the 

GPS data measurements from both the epicentral area and 

DISTANCE CHANGES 

TIME 

Figure 3.-Line-length changes observed on the six baselines of the El reflectors. Hence, the observations of length from a single evening are 

Sereno network using a two-color geodimeter. The error bars represent plotted with overlapping error bars. The data for the baseline BOHLthe one standard deviation level derived by Langbein (1989). Typically, CENTRAL were not used in the modeling that is discussed in the text. 

each baseline is measured twice during a single evening using different 
D250 AlTEiRSHOCKS AND POSTSEISMIC EmCTS 

the Black Mountain profile in terms of deep, dextral slip 

on the Loma Prieta rupture plane and thrust slip on a fault 

plane northeast and paralleling the San Andreas fault. Although the lateral extent of the thrust fault is poorly constrained from the GPS data, Burgmann and others model 

this fault as extending beneath the baseline from BAY to 

VASONA. Given Burgmann and others' slip rate of 120 

mdyr of thrust on this foothills fault, it would predict 

approximately 3.6 mdyr of extension on this baseline, 

which is clearly not the case. Instead, if thrust is indeed 

occurring, then it must be located southeast of the twocolor network. 

However, the 100 mdyr of dextral slip on the plane 

representing rupture of the Loma Prieta fault determined 

by Burgmann and others is broadly consistent with the 

two-color measurements. In fact, if we use the model of 

time-dependent slip on the San Andreas fault shown in 

figure 4, and slip on the two planes discussed in Burgmann 

and others, a satisfactory fit to the two-color data is obtained. The inferred value of postseismic slip on the 

Loma Prieta plane using the two-color data is 157k23 

mdyr, which is within 50 percent of the value estimated 

by Burgmann and others of 100 mmlyr. However, the 

model derived from the two-color data implies -14k6 

mdyr of thrust on the foothills fault, which is not consistent with the estimate of Burgmann and others of 120 

mdyr. 

Finally, the third set of postseismic measurements, which 

have been described by Rymer (1990), are a series of 

taped distance measurements on small-aperture (0.5 to 7.7 

m) quadrilaterals. One quadrilateral is located 3.4 km 

southeast of the station GAP straddling a subparallel trace 

of the San Andreas fault. At this site during the 6 months 

following the main shock, Rymer (1990) recorded a cumulative of 5k2 mm of right-lateral slip consistent with 

the observations described here. 

The unique feature of the two-color data is that we can 

infer a decrease in slip rate over the 2 years following the 

INFERRED SLIP 

TIME 

Figure 4.-Values of inferred slip as a function of time following the slip as a function of the logarithm of time for two differing decay times, 

Loma Prieta main shock. The error bars represent one standard devia- 0.5 and 20 days. The value of slip for December 1990 was not used in 

tion. Since the total amount of slip following the earthquake is un- estimating the parameters of the logarithmic curves. 

known, only the relative slip is plotted. The two curves show the predicted 
POSTSEISMIC SLIP ON THE SAN ANDREAS FAULT AT THE NORTHWESTERN END OF RUPTURE ZONE D25 1 

main shock. To demonstrate that slip has been occurring 

during the interval between January 1990 and November 

1991, a different function of slip with time is fit to the 

observed line-length changes. For the interval between 

October 1989 and the January 1990 surveys, slip is assumed to occur as a single event at an unspecified time. 

However, for the observations from early January through 

November 1991, slip is assumed to occur at a constant 

rate. Estimates of these two parameters yield 4.7k0.6 mm 

of slip between early October 1989 and early January 

1990, and a significant slip rate of 2.1k0.5 mdyr for the 

period following January 1990. This rate appears to be 

marginally faster by 1.7 standard deviations than the apparent background rate of mdyr 1.2h0.2 mdyr estimated 

for the Black Mountain network (fig. 1) by Prescott and 

others (1989b). 

Although the slip rate recorded within days after the 

main shock is significant, the rate of postseismic slip appears to be decreasing inversely with time. Along with 

the estimated slip, figure 4 shows two curves representing 

postseismic slip as Uolog(t/~+l)+B, where t is the time of 

observation since the time of the main shock (Langbein 

and others, 1983). Here we have specified two differing 

values for the decay time T as 0.5 days and 20 days and 

estimated corresponding values for Uo and B by leastsquares fitting. With the exception of the inferred slip in 

late 1990, the comparison in figure 4 shows reasonable 

agreement between the inferred values of slip as a function of time and the postseismic slip predicted from the 

logarithmic functions. However, those functions having 

decay times of less than a month fit the observed slip data 

better than those functions with decay periods in excess 

of several months, Thus the short decay time (7) found 

here is consistent with that found in other cases of shallow, postseismic slip following the 1966 Parkfield and 

1979 Imperial Valley earthquakes (Smith and Wyss, 1968; 

Langbein and others, 1983). 

Breckenridge and Simpson (this volume) show data for 

changes in the rates of fault slip on other fault segments 

in response to the Loma Prieta earthquake, which they 

suggest is caused by the Loma Prieta earthquake perturbing the regional stress field. They demonstrate that the 

spatial pattern of stress changes predicted from their stress 

model was consistent with the changes in slip that they 

observed from creepmeters along the Hayward, San Andrea~, and Calaveras fault system. Furthermore, their 

model predicts an increase in rate of slip on the section of 

the San Andreas fault that bisects the two-color network 

immediately following the Loma Prieta earthquake. The 

results presented here support their model. However, their 

creep data show that the time constant is about a year for 

the postseismic slip rate to return to its secular rate. This 

observation is inconsistent with the time-constant for 

postseismic slip determined from the measurements discussed here. It could be that the time-constant of deformation detected by the creep data and the GPS observations is due to a deep-seated process with a long time 

constant, but that there is a short time constant associated 

with near-surface fault-slip. 

REFERENCES 

Burford, R.O., and Harsh, P.W., 1980, Slip on the San Andreas fault in 

central California form alinement array surveys: Bulletin of the 

Seismological Society of America, v. 70, p. 1233-1261. 

Burgmann, R., and Segall P., 1991, Postseismic GPS monitoring 

NW of the 1989 Loma Prieta rupture zone: Eos, v. 72, no. 44, p. 

119. 

Johnston, M.J.S., Linde A.T., Gladwin M.T., and Borcherdt, R.D., 

1987, Fault failure with moderate earthquakes: Tectonophysics, v. 

144, p. 189-206. 

Johnston, M.J.S., Borcherdt R.D., and Linde A.T., 1986, Short period 

strain (0.1-10-5 s): Near source strain field for an earthquake (ML 

3.2) near San Juan Bautista, California: Journal of Geophysical 

Research, v. 91, p. 11497-11502. 

King, N.E., Savage, J.C., Lisowski, M., and Prescott, W.H., 1981, 

Preseismic and co-seismic deformation associated with the Coyote 

Lake, California earthquake: Journal of Geophysical Research, v. 

86, p. 892-898. 

Langbein, J., McGarr, A., Johnston, M.J.S, and Harsh, P.W., 1983, 

Geodetic measurements of postseismic crustal deformation following the 1979 Imperial Valley earthquake, California: Bulletin of 

the Seismological Society America, v. 73, p. 1203-1224. 

Langbein, J., 1989, The deformation of the Long Valley caldera, eastem California, from mid-1983 to mid-1988; measurements using a 

two-color geodimeter: Journal of Geophysical Research, v. 94, p. 

3833-3850. 

Langbein, J., Burford, R.O., and Slater, L.E., 1990, Variations in fault 

slip and strain accumulation at Parkfield, California: Initial results 

using two-color geodimeter measurements, 1984-1988: Journal of 

Geophysical Research, v. 95, p. 2533-2552. 

Lisowski, M., and Prescott, W.H., 1981, Short-range distance measurements along the San Andreas fault system in Central California, 

1975-1979: Bulletin of the Seismological Society America v. 71, 

p. 1607-1624. 

Lisowski, M., Prescott, W.H., Savage, J.C., and Johnston, M.J.S., 

1990, Geodetic estimate of co-seismic slip during the 1989 Loma 

Prieta, California, earthquake: Geophysical Research Letters, v. 

17. 

Lisowski, M., Savage, J.C. and Prescott, W.H, 1991, Surface deformation after the Loma Prieta, California Earthquake: Eos, v. 72, 

no. 44, p. 119. 

Prescott, W.H., Davis, J.L, and Svarc, J.L, 1989a, Global positioning 

system measurements for crustal deformation: Precision and accuracy: Science, v. 244, p. 1337-1340. 

Prescott, W.H., Savage, J.C. and Lisowski, M., 1989b, Crustal strain, 

in National Earthquake Reduction Program, Summaries of Technical Reports (v. 29): U.S. Geological Survey Open-File Report 90- 

54. 

Rymer, M.J., 1990, Near-fault measurement of postseismic slip associated with the 1989 Loma Prieta, California, Earthquake: Geophysical Research Letters, v. 17, p. 1789-1792. 

Savage, J.C., and Prescott, W.H., 1973, Precision of geodolite distance measurements for determining fault movements: Journal of 

Geophysical Research, v. 78, p. 6001-6008. 

Savage, J.C., and Prescott, W.H., 1978, Asthenosphere readjustment 
D252 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

and the earthquake cycle: Journal of Geophysical Research, v. 83, 

p. 3369-3376. 

Savage, J.C, Prescott, W.H., Lisowski, M., and King, M, 1979, 

Geodolite measurements of deformation near Hollister, California, 

1971-1978:- Journal of Geophysical Research, v. 84, p. 7599- 

7615. 

Savage, J.C, Prescott, W.H, and Lisowski, M., 1987, Deformation 

along the San Andreas fault 1982-1986 as indicated by frequent 

Geodolite measurements: Journal of Geophysical Research, v. 92, 

p. 4785-4797. 

Slater, L.E., and Huggett, G.R, 1976, A multi-wavelength distancemeasuring instrument for geophysical experiments: Journal of 

Geophysical Research, v. 81, p. 6299-6306. 

Smith, S.W., and Wyss, M., 1968, Displacement on the San Andreas 

fault initiated by the 1966 Parkfield earthquake: Bulletin of the 

Seismological Society of America, v. 68, p. 1955-1974. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKE OCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

MODELS OF' POSTSEISMIC DEFORMATION AND STRESS TRANSFER 

ASSOCIATED WITH THE LOMA PRIETA EARTHQUAKE 

CONTENTS 

By M.F. Linker, 

Harvard University and US. Geological Survey; 

and J.R. Rice, Harvard University 

Page 

D253 

253 

255 

256 

258 

258 

260 

263 

263 

264 

266 

27 1 

27 1 

272 

272 

272 

273 

273 

ABSTRACT 

Recent geodetic observations indicate that the velocity 

field at the Earth's surface has been perturbed by the Loma 

Prieta earthquake. We interpret this change in terms of 

models in which transient postseismic slip occurs beneath 

the locked seismogenic portion of the San Andreas fault 

zone. In our three-dimensional finite element calculations, 

the deep aseismic region of the fault zone is either treated 

as linear Maxwell viscoelastic or is made to follow a 

steady-state version of the laboratory-derived rate- and 

state-dependent friction law, in which slip rate depends 

exponentially on the ratio of shear stress to effective normal stress. We refer to this second rheology as our "hotfriction" model. Comparison of model predictions to 

observed postseismic displacements provides a constraint 

on the ratio of Maxwell relaxation time, t,., for material in 

the deep aseismic region of the fault zone to fault-zone 

thickness, h. Neither quantity can be constrained independently. Best fit to the initial 1.3 years of data is obtained 

with t,/h = 0.3 yrlkm thickness. If we assume that 

postseismic slip has occurred in a fault zone of thickness 

hS1 km, then using a shear modulus of 30 GPa we obtain 

an estimate of the effective viscosity of material in the 

deep aseismic region of the fault zone, qenS3xl~17 Pa-s. 

Previous estimates made for the mater~al of the lower 

crust, based on laboratory measurements of steady-state 

dislocation or diffusion creep, exceed this value by at 

least an order of magnitude. The laboratory-derived hotfriction model can yield postseismic deformation of magnitude comparable to that observed, but only if ddd(1n V) 

is of the order 0.5 MPa, where T is the resistance to sliding and V is the slip rate. Laboratory measurements indicate that ddd(1n V) can be written approximately as co, 

where o is the effective normal stress equal to total normal stress minus pore pressure and c is on the order of 

0.015 for mid-crustal conditions. To be consistent with 

our models, the effective normal stress in the deep aseismic 

portion of the fault zone must, therefore, be extremely 

low, perhaps on the order of 30 MPa. In the context of the 

laboratory-derived hot-friction model, the occurrence of 

postseismic deformation may be evidence that pore pressure in the aseismic portions of the fault zone and perhaps 

the lower crust is near lithostatic. 

INTRODUCTION 

The occurrence of the Loma Prieta earthquake brought 

immediate concern that another large and damaging earthquake might occur in the San Francisco Bay region in the 

near future (Working Group, 1990). This concern rose 

from numerous examples of earthquake triggering and 

migration (see, for example, Richter, 1958; Mogi, 1968; 

Allen 1969; Yonekura, 1975; Scholz, 1977; Sykes and 

others, 1981; Doser, 1986; Stein and Ekstrom, 1992) in- 
D254 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

eluding two earthquake pairs in the San Francisco Bay 

region (Toppazada and others, 1981). In 1836 an M=7 

event occurred on the northern half of the Hayward fault 

and 2 years later an M=7 event occurred on the San Francisco peninsula. Similarly, in 1865 an Mz6.5 event occurred in the vicinity of Loma Prieta and 3 years later an 

M=7 event occurred on the southern segment of the 

Hayward fault. More recently, on the Calaveras fault three 

moderate events have occurred in spatial sequence, migrating towards the north (Oppenheimer and others, 1990). 

Previous workers have suggested that earthquake migration and triggering might be the result of stress transfer between earthquake source regions. Numerous attempts 

to relate static coseismic stress changes to the occurrence 

of future events (Stein and Lisowski, 1983; Mavko and 

others, 1985; Oppenheimer and others, 1988; Poley and 

others, 1987; Hudnut and others, 1989; Du and Aydin, 

1990; Seeber and Armbruster, 1990; Michael and others, 

1990; Michael, 1991; Reasenberg and Simpson, 1992) 

demonstrate that the magnitude of the static coseismic 

stress change is likely to be small, except perhaps in the 

region immediately adjacent to the main rupture. Though 

postseismic relaxation associated with inelastic deformation processes could act to increase the magnitude of the 

stress transfer relative to the static coseismic stress change, 

the total stress change in the far-field is still likely to be 

modest compared to typical earthquake stress drop (Rice 

and Gu, 1983). Furthermore, the rate of stress transfer 

resulting from postseismic relaxation may be small compared to the background loading rate, and so its contribution to the far-field stress state is likely to be negligible, 

unless postseismic relaxation occurs at a sufficiently high 

rate (Lehner and others, 1981). Nevertheless, the observations of earthquake triggering and migration are numerous, and so we have undertaken a study of the stress 

transfer associated with the Loma Prieta earthquake. 

In the first 50 years following the great 1906 San Francisco earthquake, the shear strain rate near the San Andreas fault decreased by about a factor of three (fig. 1) 

(Thatcher, 1983), and so we might expect similar transient postseismic deformation to follow the Loma Prieta 

earthquake. Indeed, geodetic data indicate that postseismic 

deformation has taken place at rates that exceed the preearthquake rate (Burgmann and others, 199 1, this chapter; 

Lisowski and others, 1991a, 1991b; Burgmann and others, 1992; Savage and others, 1994). Therefore, the Loma 

Prieta earthquake provides us an opportunity not only to 

study the stress transfer that could potentially lead to a 

future major earthquake, but more generally, to obtain 

constraints on the rheological properties of the crust, fault 

zone, and upper mantle. 

We use the finite element method to examine three dimensional models of postseismic deformation and stress 

transfer associated with the Loma Prieta earthquake. We 

use preseismic (Lisowski and others, 199 1 b) and preliminary postseismic measurements made during the initial 

1.3 years following the earthquake (Burgmann and Segall, 

1991; Lisowski and others, 199la; Burgmann and others, 

1992) to constrain parameters that control the relaxation 

process. Our models are additionally constrained by consideration of heat flow data, laboratory-based rheologies 

and constitutive parameters, seismic observations of preTIME SINCE 1906 (YEARS) 

Figure 1.Ã‘Geodeticall measured right-lateral shear strain rate near the San Andreas fault in northern California following the great 1906 San Francisco earthquake (from Thatcher, 1983). Inset map indicates the location 

of the networks. 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUAKE D255 

earthquake microseismicity, and seismological studies of 

the main shock. 

Our paper is organized as follows. First, we study the 

magnitude of stress transfer by examining models of the 

static coseismic stress field and long-time limits of 

postseismic relaxation. We then examine time-dependent 

models and compare predicted displacements to geodetic 

data in attempts to constrain the parameters that control 

the relaxation time-scale. Our first set of time-dependent 

models incorporate linear Maxwell viscoelasticity in the 

inelastically deforming regions of the model. Our second 

set of time-dependent models employ a steady-state version of the laboratory-derived rate- and state-dependent 

friction model in which aseismic slip rate depends exponentially on the ratio of shear stress to effective normal 

stress (Blanpied and others, 1991), where the effective 

normal stress is the total normal stress minus the pore 

pressure. We refer to the nonlinear model as our "hotfriction" model. 

Our principle conclusions are as follows. (1) The largest stress changes associated with the earthquake occur 

adjacent to the main rupture, for example, along the San 

Francisco Peninsula segment of the San Andreas fault. (2) 

In that region, postseismic relaxation processes could lead 

to an increase in shear stress by an amount that exceeds 

the static coseismic stress increase. (3) Northeast to north 

of Loma Prieta, both coseismic slip and postseismic relaxation reduce the right-lateral shear stress and the compressive stress along the Hayward and Calaveras faults. 

(4) Models in which deep aseismic fault slip occurs beneath the seismogenic zone are capable of producing strikeslip displacements consistent with the initial 1.3 years of 

postseismic deformation observed along the profile of GPS 

stations that crosses through the epicenter (Lisowski and 

others, 199 1 a). These deep-slip models, however, appear 

to be incapable of producing fault-trace normal contraction compatible with the observations there and an additional source of deformation may be required. (5) 

Northwest of Loma Prieta, the postseismic displacements 

predicted by our deep-slip models may exceed the observations, though a careful treatment of all of the pre-earthquake and post-earthquake geodetic data is required to 

better establish a measure of the change in the velocity 

field associated with the earthquake. (6) The time-scale 

for postseismic relaxation appears to be sufficiently long 

so that the average background tectonic stress rate is likely 

to exceed the stress rate resulting from relaxation, except 

perhaps within about 20 km of the edge of the coseismic 

rupture. (7) In linear viscoelastic models that incorporate 

postseismic fault slip, the parameter controlling the relaxation time is the ratio of fault-zone material relaxation 

time to the thickness of the fault zone, t/h. This parameter appears to have a value of approximately 0.3 years1 

km thickness. If relaxation takes place in a shear zone 

that is less than 1 km thick and if the shear modulus is 30 

GPa, then the effective viscosity in the deforming region 

must be less than about 3x1017 Pa-s. This value is at least 

one order of magnitude lower than any obvious interpretation of laboratory measurements of the steady-state creep 

of solid crustal rocks controlled by dislocation or diffusion creep. Transient creep response to a sudden change 

in stress has not been as well characterized in the laboratory and so probably cannot be ruled out as a possible 

mechanism. (8) The laboratory-derived hot-friction model 

can yield postseismic deformation of magnitude comparable to that observed, but only if ddd(1n V) is of the 

order 0.5 MPa, where T is the resistance to sliding and V 

is the slip rate. Laboratory measurements indicate that dd 

d(1n V) can be written approximately as cor (Stesky, 1975, 

1978; Dieterich, 198 1; Ruina, 1983), where or is the effective normal stress and c is of order 0.015 for mid-crustal 

conditions (Blanpied and others, 199 1). Therefore, to be 

consistent with our models, the effective normal stress 

must be extremely low, perhaps of order 30 MPa. (9) In 

the context of the laboratory-derived hot-friction model, 

the occurrence of postseismic deformation may be evidence that pore pressure in the aseismic portions of the 

fault zone and perhaps the nearby lower crust is near 

lithostatic. 

PHYSICAL MODELS OF POSTSEISMIC 

DEFORMATION 

Models of postseismic deformation have been motivated 

primarily by observations of deformation that followed 

large earthquakes in subduction zones and along the San 

Andreas fault. In these models deformation occurs as the 

result of fault slip on a down-dip projection of the seismic 

rupture or as distributed shear in a broad zone beneath the 

brittle seismic layer (Savage and Prescott, 1978; Savage, 

1983; Thatcher, 1983; Thatcher and Rundle, 1984; Li and 

Rice, 1987; Lyzenga and others, 199 1). Unfortunately, 

the displacement field at the earth's surface is rather insensitive to the location of the deformation source, and so 

generally it is not possible to reject either of these models 

on the grounds that they do not fit the deformation data 

(Savage and Prescott, 1978; Savage, 1990; Segall, 199 1). 

Therefore, physical models of postseismic deformation 

must be defended as well as motivated on broad geophysical and geological grounds. England and Molnar 

(1991), for example, propose that the upper continental 

crust is weak relative to the lower parts of the lithosphere 

and that crustal blocks passively follow the deformation 

of the substrate. Within this context, postseismic deformation could be caused by transient interaction between 

the individual crustal blocks as well as between the blocks 

and the substrate. This concept has been adopted either 

explicitly or implicitly in much of the modeling mentioned above. 
D256 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

We examine three models of the geometrical distribution of inelastic material that shears postseismically (fig. 

2). In Model 1, postseismic fault slip is allowed to occur 

down-dip from the seismogenic zone but is confined to 

the crust. In Model 2, slip is allowed to extend from the 

base of the seismogenic zone downward through the crust 

and into the upper mantle. In Model 3, slip is allowed to 

extend from the base of the seismogenic zone downward 

through the crust to where it terminates at its intersection 

with a horizontal detachment zone in the lower crust. 

In the Earth, immediately after the earthquake, stress 

changes and presumably slip rates are highest near the 

edge of the coseismic rupture. Therefore, Model 1 can be 

thought of as a short-time version of either Model 2 or 

Model 3. 

In Model 2, crustal blocks are thought of as effectively 

welded to the substrate, at least on the time-scale of the 

repeat-time for large crustal earthquakes. Postseismic motion is interpreted as the result of transient motion between the crustal blocks whose boundaries extend into the 

upper mantle. We refer to Model 2 as our throughgoing 

fault model. 

Model 3 is based on previous studies of postseismic 

deformation (for example, Li and Rice, 1987) that were 

motivated by the recognition that elevated temperatures 

in the lower crust may give rise to inelastic flow in response to the stress changes associated with earthquakes 

(Brace and Kohlstedt, 1980; Meissner and Strehlau, 1982; 

Sibson, 1982; Chen and Molnar, 1983). This class of models can be further motivated by compilations of seismic, 

geological, and geophysical data in central California (see, 

for example, Saleeby, 1986; Fuis and Mooney, 1990) and 

by electrical conductivity measurements, which indicate 

that horizontal seismic reflectors in the lower crust frequently coincide with horizontal layers of high conductivity and, by inference, with regions of elevated fluid content 

(Hyndman and Shearer, 1989) and fluid pressure (Rice, 

1992). All of these observations are in accord with the 

commonly held view that large-scale horizontal shearing 

or fault-slip occurs in the middle to lower crust. In Model 

3, motion of the shallow portion of the crust is coupled 

through this horizontal shear zone to the motion of the 

substrate that drives the system. We refer to Model 3 as 

our detachment model. 

Tse and Rice (1986) studied models based on rate- and 

state-dependent friction (Dieterich, 198 1 ; Ruina, 1983), 

in which the fault-zone rocks at depth are hot enough to 

be in the "velocity strengthening" range of frictional response. In this case, time-varying stress occurs naturally 

in the aseismic portion of the fault zone. In contrast, Li 

and Rice (1987), Fares and Rice (1988), and Ben-Zion 

and others (1993) treated the aseismic portion of the fault 

zone within the crust as freely slipping, so that only the 

shear stress in the basal shear zone varies with time. Their 

free-slip condition is consistent with an assumption that 

the relaxation time for postseismic crustal fault slip is 

considerably shorter than for basal shear. 

In this paper, we include the effects of time-varying 

stress in the aseismic portion of the fault zone as well as 

in the basal shear zone. We treat the region outside the 

fault zone and basal shear zone as purely elastic, with the 

thought that stress changes associated with earthquakes 

are too small to cause any appreciable inelastic deformation there. We additionally include the effects of finite 

rupture length and so our models are three dimensional. 

THREE-DIMENSIONAL FINITE 

ELEMENT MODELS 

We use the finite element method to model stress 

changes and deformation associated with the earthquake. 

Our modeling strategy is as follows. Coseismic slip is 

imposed on a segment of the fault zone that represents the 

coseismic rupture surface. This imposed slip produces a 

stress field throughout the body that represents the static 

coseismic stress change in the earth. We then use two 

approaches to model the postseismic deformation and stress 

transfer resulting from that coseismic stress field. First we 

examine coseismic and long-time fully-relaxed postseismic 

stress states. Then we examine time-dependent models in 

which stress and displacement are calculated while the 

system relaxes toward the fully-relaxed state. By comparing the computed time-dependent displacements to deformation data, we attempt to constrain the parameters that 

control the time-scale for relaxation. 

The solutions we present employ fixed-displacement 

boundary conditions along the outer surfaces of the model 

where it has been cut away from the surroundings. The 

remote boundaries of our model are located sufficiently 

far from the coseismic rupture so that fixed-displacement 

and stress-free boundary conditions give nearly the same 

solution in the regions of interest. 

We ignore the effects of the shallow creeping zone along 

the San Andreas fault that begins southeast of Loma Prieta 

near San Juan Bautista and only consider models that are 

geometrically symmetric along strike about the epicenter. 

We additionally ignore any deformation associated with 

slip on the Hayward and Calaveras faults. We assume a 

crustal thickness of 25 km, which is representative of the 

region near the San Andreas fault (Oppenheimer and Eaton, 

1984), and represent the horizontal shear zone in the lower 

crust with a 5-km-thick layer of elements that extends 

from 20 to 25 km in depth. 

Our finite element mesh represents a volume with dimensions 2 17 km along strike by 265 km perpendicular to 

strike by 62.5 km in depth (fig. 2A). Minimum elementdimensions correspond to 3.5x2.5x1.0 km3. Approximately 

10,000 nodes, or 30,000 degrees of freedom, comprise the 
A 3D Mesh for Lorna Prieta 

B Fault element detail near rupture zone 

(from Model #I) 

E Model #3 : 

Crustal fault + basal shear zone 

C Model#l: 

Crustal fault 

D Model #2: 

Through-going fault 

Figure 2.-Finite element (FE) models used to represent three distributions of 

regions undergoing viscoelastic relaxation. A, Three-dimensional FE mesh used 

to model the crust and upper mantle near Loma Prieta. B, Close-up view of 

fault-zone elements in and around the coseismic rupture zone for a particular 

model of postseismic relaxation, Model 1. Black region denotes core of coseismic 

rupture where the imposed coseismic slip is uniform. In the surrounding darkgray region the imposed coseismic slip tapers to zero away from the core of the 

rupture. Lightest gray region slips aseismically in response to the stress changes 

induced by coseismic slip. Gray regions above and below the aseismically 

slipping zone are locked. C, View of Model 1 in which the southwest side of 

model has been cut away to reveal the dipping fault zone. The aseismically 

slipping portion of the fault is confined to lie between the shallow seismogenic 

zone and the base of the crust. D, View of Model 2 in which the aseismically 

slipping portion of the fault zone extends through the Moho and into the upper 

mantle. E, View of Model 3 in which the aseismically slipping portion of the 

fault zone is confined to the crust but intersects a horizontal shear zone, or 

detachment zone. that lies in the lower crust. 
D25 8 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

mesh of eight-noded isoparametric-brick elements. We use 

the finite element code ABAQUS (Hibbitt and others, 

1991). Elastic three-dimensional and viscoelastic two-dimensional models were run on a SUN 4 workstation and 

require several hours of CPU time and 10 to 20 Mbytes of 

disk space for output. All computations for three-dimensional time-dependent models were performed on Cray 2 

and Cray YMP supercomputers at the National Center for 

Supercomputing Applications (Urbana, Illinois). The linear viscoelastic solutions typically require about 3 hours 

of Cray CPU time for 65 time increments using an explicit time-integration method to reach about 5 years of 

model time. The nonlinear hot-friction models use appreciably more CPU time, since the creep rate varies dramatically during the relaxation process. Typical run times 

for three dimensional nonlinear models range from 4 to 

10 hours on the Cray 2 to reach 2 years of model time. It 

is likely that some computation time could be saved if an 

implicit integration procedure was used, but we did not 

take that approach. 

In our models, the fault zone is represented by a thin 

layer of finite elements (fig. 2). The seismic portion of 

the fault zone is treated as elastic, and coseismic slip is 

imposed via a shear transformation strain, y T, which 

changes the stress-free shape of the material in the rupture zone (Aki and Richards, 1980; Rice, 1980). The 

equivalent seismic slip is hyT, where h is the thickness of 

the layer of fault elements. At distances greater than order 

h from the edge of the rupture, this technique produces 

results equivalent to those obtained by imposing slip on a 

surface. 

Any estimate of the stress field close to the edge of the 

coseismic rupture will be extremely sensitive to one's estimate of the slip distribution. One might incorporate estimates of the slip distribution determined from seismic 

records (Beroza, 199 1 ; Hartzell and others, 199 1 ; S teidl 

and others, 1991; Wald and others, 1991) or measurements of the static displacement field (Marshall and others, 1991), but we have not taken this approach. Future 

work on this problem may be warranted since the largest 

stress changes occur in this near-rupture region. Farther 

from the edge of the rupture, the stress field is less sensitive to the slip distribution and can therefore be estimated 

with greater confidence. 

In this spirit, we assume that the transformation strain 

and hence coseismic slip are uniform except in the outermost elements of the rupture zone (fig. 2B). In these perimeter elements, we specify that the shear modulus is 

only 1 percent of the nominal shear modulus, and Poisson's 

ratio is chosen so that the bulk modulus remains uniform. 

In addition, the transformation strain tapers to zero here. 

This procedure roughly simulates a freely slipping zone 

and has the effect of suppressing large local stresses that 

are poorly modeled without extensive mesh refinement 

and that have magnitudes which depend critically on the 

slip distribution near the edge of the rupture. 

The characteristics of our coseismic rupture were chosen to correspond to estimates obtained by other workers. 

Specifically, we used scalar moment = 3x1019 Nm, 

dip=64', and rake=145'. The rupture extends from 5 km 

to 17.5 km in depth (fig. 2B). The shape of the rupture 

perimeter was chosen to mimic the distribution of 

preseismic microseismicity (Olson, 1990) and initial aftershock distribution (Dietz and Ellsworth, 1990). The chosen moment and rake is that from Lisowski and others' 

(1990) preferred model, in which they specify uniform 

slip on a rectangular dislocation in a homogeneous elastic 

half space to model horizontal geodetic data. Seismic and 

other geodetic estimates of the moment generally range 

from 1 to 4x1019 Nm, dips range from 55' to 80Â° and 

rakes range from 115' to 155' (see summary by Marshall 

and others, 1991, table 7). 

MAGNITUDE OF COSEISMIC AND 

LONG-TIME POSTSEISMIC STRESS 

TRANSFER 

ANALYTICAL MODELS 

Before presenting results for the Loma Prieta earthquake, we review simple analytical models to obtain an 

understanding of how stress fields associated with generic 

models of stress transfer vary in space. Our analysis follows the work of Rice and Gu (1983). 

We consider two limiting states: (1) immediate 

coseismic and (2) fully-relaxed downward continuation of 

the crustal fault zone and underlying horizontal substrate. 

The latter corresponds to Model 3 (fig. 2E) at times sufficiently large that relaxation has gone to completion. We 

compare the far-field solution for a finite, vertical, strikeslip dislocation in an elastic half-space to the corresponding solution in an elastic plate (fig. 3). 

The static coseismic change in right-lateral shear stress 

on the still locked segment of the fault zone as a function 

of distance, r, from the center of the rupture is 

and the long-time state when the deep aseismic portion of 

the fault zone and the horizontal detachment are fullyrelaxed is 

where D is rupture depth, L is rupture length, H is the 

thickness of the elastic portion of the crust, and AT = 

2pAu/nD is the coseismic stress drop associated, via a 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUAKE D259 

crack model, with a long surface-breaking strike-slip rupture over depth D, with average slip Au, in a half-space 

with shear modulus p. These solutions are valid at distances from the center of the rupture that are large compared to rupture half-length. The ratio of long-time to 

short-time stress is ro/rw = (5/4)(H/D)(r/D). At distances 

from the rupture r >> D relaxation of the horizontal substrate yields postseismic stress changes that greatly exceed the coseismic change. 

For example, given an earthquake with average stress 

drop At, rupture length equal to twice the elastic plate 

thickness, L=2H, and rupture depth D=H/2, then at a distance from the center of the rupture equal to 2H the static 

coseismic stress change is approximately 0.0164~ while 

the stress change after complete relaxation of basal shear 

stress is approximately 0.164~. In this example, 

postseismic relaxation increases the far-field stress by an 

order of magnitude relative to the static coseismic stress 

- Average stress droo = AT 

Figure 3.-Dislocation models comparing (A) short-time coseismic static shear stress change, 

& to (5) long-time limit, -rd of Model 3 (fig. 2E) in response to coseismic stress drop, AT, on 

the main rupture surface of depth D and length L. In the first case stress is carried by the elastic 

half space, while in the latter case stress is carried by an elastic plate, with thickness H, floating 

on a traction-free substrate meant to represent a fully-relaxed horizontal shear zone in the lower 

crust. The solutions are valid at distances from the center of the rupture that are large compared 

to rupture half-length. The transition from short-time to long-time stress state corresponds to a 

change from 1/r to 1/r decay of stress with distance, r, from the center of the rupture. 
D260 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

change. Note, however, that the total magnitude of the 

stress change in the far-field is still only about 16 percent 

of the earthquake stress drop. 

COSEISMIC AND LONG-TIME FULL-RELAXED 

POSTSEISMIC STRESS STATES 

We estimate the coseismic and long-time fully-relaxed 

postseismic static stress changes associated with our three 

geometrical models of the Loma Prieta region (fig. 2). 

Note that in inhomogeneous viscoelastic systems, the fullyrelaxed-limit solution does not necessarily provide an upper bound on changes of stress during the relaxation 

process. 

The models are now purely elastic and postseismic relaxation is represented by introducing freely slipping regions that, in the earth, might deform in a time-dependent 

manner. This approach saves dramatically on computation time since only one "time-step" is need to achieve 

"relaxation." The states obtained with these elastic models should be interpreted as representing the long-time 

fully-relaxed limits of the corresponding viscoelastic models that are presented in later sections of this paper. 

The calculations are done using the finite element 

method as discussed previously. The "free-slip" regions 

are incorporated by specifying that the shear modulus of 

the "relaxed" material is reduced relative to the surroundings. To simulate free-slip with this procedure, one needs 

^faul/l' to be small compared to &rroundin8/H, where h 

is the thickness of the layer of fault elements and H is a 

scale-length for the surrounding region. We use a reduced 

shear modulus that is 1 percent of the nominal shear 

modulus, 30 GPa, and choose Poisson's ratio so that the 

bulk modulus is uniform. In the finite element mesh, the 

thickness of the fault zone and basal shear zone are 1 km 

and 5 km, respectively. Further reduction of the low shear 

modulus does not affect the stress distribution appreciably. The mantle is made stiffer than the crust by a factor 

of two. 

Profiles of the changes in stress along profiles corresponding approximately with the San Francisco peninsula 

segment of the San Andreas fault and to the Hayward and 

Calaveras faults (fig. 4) are plotted in figure 5 as components of the traction vector resolved onto vertical planes. 

The values plotted in figure 5 are in general agreement 

with those obtained by Simpson and Reasenberg (this chapter), who used shear traction-free rectangular dislocations 

to represent relaxed, deep fault zones and horizontal detachment faults. 

In the San Andreas profiles (figs. 5A and 5B), the values of stress represent averages over the bottom one-quarter of the locked seismogenic zone, which extends from 0 

to 10 km in depth in our models (fig. 2). Very close to the 

end of the rupture, computed stress changes are sensitive 

to the assumed spatial distribution of slip. We have not 

attempted to incorporate detailed estimates of the coseismic 

slip distribution in our models and have only plotted those 

stress changes that we consider to be relatively reliable in 

this context. 

Along the San Andreas profile, the coseismic change in 

static shear stress is largest close to the end of the rupture, 

as expected (fig. 5A). Here, relaxation can double or even 

quadruple the change in shear stress. Farther from the end 

of the rupture, at distances that are large compared to the 

17.5 km rupture depth, the ratio of postseismic to coseismic 

shear stress reaches values exceeding 10:1, but the total 

change in shear stress is quite small. These three observations are qualitatively consistent with the results obtained 

above with simple analytical models. Both coseismic slip 

and subsequent postseismic relaxation result in faultnormal compression relative to the pre-seismic state (fig. 

55). 

The enhancement of shear and normal stress on the San 

Andreas fault that results from postseismic relaxation can 

be rationalized if one considers the stress concentration in 

the vicinity of a shear crack in an elastic body. In our 

fully-relaxed models, the tip of this shear crack lies along 

the base of the seismogenic zone, adjacent to the points 

sampled in figures 5A and 55. Here, the stress field will 

be dominated by near-field terms so that the state of stress 

along the up-dip projection of the aseismically slipped 

zone will be approximately pure shear. Recall that 

coseismic slip was oblique, reverse plus right-lateral, on a 

steeply SW-dipping fault segment so that postseismic slip 

on the aseismic portion of the fault zone will likewise be 

oblique, reverse plus right-lateral. By Mohr-circle analysis it is simple to demonstrate that the resolved shear and 

normal stress on a vertical plane extending upward from 

the tip of the aseismically slipped zone will therefore be 

right-lateral and compressive, in agreement with figures 

5A and 5B, respectively. 

Stress changes along the HaywardICalaveras trend (fig. 

4) are plotted in figures 5C and 5D. There we plot the 

change in traction that occurs on a vertical plane trending 

21' clockwise from the strike of the modeled coseismic 

rupture plane versus position along strike measured from 

a point lying adjacent to the Lorna Prieta epicenter and 

coinciding, approximately, with the southeast end of the 

M=6.2, 1984 Morgan Hill rupture (Bakun and others, 

1984). 

Coseismic slip and postseismic relaxation lead to decreases in both right-lateral shear stress and in compressive stress along the HaywardlCalaveras trend. At positions 

0 to 7 km the coseismic decrease in right-lateral shear 

stress is exceeded by the decrease in compressive stress. 

The ratio of the two is about 0.1010.23 = 0.4 (figs. 5C and 

5D). Therefore, this portion of the fault would be moved 

toward failure if the effective coefficient of friction is 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUAKE D261 

greater than about 0.4 (see extensive discussion by Calaveras trend. In the far-field, postseismic relaxation of 

Reasenberg and Simpson (1992) and by Simpson and the lower crust can result in changes in shear stress that 

Reasenberg (this chapter)). are large relative to the static coseismic change (fig. 5C, 

Postseismic relaxation appears to have only a moderate Model 3) but, as noted previously, the total magnitude of 

effect on the overall state of stress along the Haywardl the stress change is rather small. 

Figure 4.-Overlay of finite element mesh onto a map of the San Francisco Bay region. The shoreline and major faults 

are indicated by light and heavy sinuous lines, respectively. The Loma Prieta epicenter (U.S. Geological Survey Staff, 

1990) is plotted as a square and geodetic stations occupied by the GPS are plotted as triangles. The perimeter of the 

finite element mesh is indicated by the dark bounding box and the profiles along which computed stresses and 

displacements are plotted in subsequent figures are indicated by solid lines. The relative position of the finite element 

mesh is chosen so that the strike of the model San Andreas fault is N44OW. The model center-line, normal to the fault 

trace, passes though the main shock epicenter and the model San Andreas fault overlies the mapped fault trace NW of 

Loma Prieta. 

MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUAKE D263 

A RELATIVE TIME-SCALE FOR POSTSEISMIC 

RELAXATION 

In later sections of this paper we analyze time-dependent viscoelastic models to constrain the absolute timescale for the relaxation process. Here, we construct a 

relative time-scale by estimating the amount of time that 

background tectonic loading requires to contribute a change 

in shear stress equal to that resulting from long-time 

postseismic relaxation. This relative time-scale can later 

be compared to estimates of the system relaxation time. 

If, for example, the average tectonic contribution to the 

rate of shear-stress accumulation is the ratio of a typical 

earthquake stress drop to an earthquake cycle time, then 

this tectonic loading rate is likely to be about 5 MPaJ200 

yr = 0.025 MPa/yr. Along the Peninsula segment of the 

San Andreas fault, 80 km from the epicenter, the increase 

in shear stress resulting from complete relaxation of the 

lower crust is about 0.05 MPa (fig. 5A, Model 3). This 

increase is then equivalent to about 2 years of tectonic 

loading. Postseismic relaxation is, therefore, unlikely to 

contribute significantly to the stress rate at distances beyond 80 krn, unless the time-scale for relaxation of the 

lower crust is less than a few years. In contrast, 40 km 

from the epicenter the corresponding increase in shear 

stress is about 0.4 MPa (fig. 5A, Model 3) and consequently represents about 16 years at the tectonic loading 

rate. Therefore, relaxation of the lower crust will contribute significantly to the stress rate in the region near the 

end of the Loma Prieta rupture, provided that the timescale for relaxation is on the order of 16 years or less. 

changes to the state that would have existed in absence of 

the earthquake. It can be shown that the postseismic problem can be treated exactly in this manner if one assumes 

that a preexisting steady-state of inelastic shearing existed 

in all aseismic regions of the model prior to the earthquake. 

The more traditional approach to this problem is to 

precondition the model by imposing a cycle of repeated 

characteristic earthquakes until the system responds in a 

periodic manner (Savage, 1983; Thatcher and Rundle, 

1984; Li and Rice, 1987; Dmowska and others, 1988; 

Fares and Rice, 1988; Lyzenga and others, 1991; BenZion and others, 1993; Reches and others, 1994). In our 

situation, this strategy may be difficult to apply with any 

confidence since there is some question as to how the 

Loma Prieta earthquake relates to characteristic events for 

the region (Anderson, 1990; Segall and Lisowski, 1990). 

We represent the aseismically shearing regions (fig. 2) 

with Maxwellian viscoelasticity as 

where E" is the elastic strain, 87s the inelastic creep 

strain, and ei is the total strain. We adopt the standard 

assumptions of elementary plasticity theory-that the inelastic creep strain is not affected by hydrostatic pressure 

and that the volumetric creep strain is zero (see, for example, Malvern, 1969). The creep law can then be written 

in terms of deviatoric stress, 5,. = 0,. - (1 / 3)akk6,, where 

cr,, is the stress tensor and we sum on repeated indices. In 

the creeping portion of the model 

TIME-DEPENDENT MODELS 

We analyze our three geometrical models of postseismic 

relaxation (fig. 2) while employing two different Maxwell 

viscoelastic rheologies. First, we use linear viscoelasticity 

to model the presumably nonlinear behavior of materials 

in the inelastically shearing portions of the fault zone and 

lower crust. With this approach, any estimate of relaxation time, obtained by fitting our models to data, must be 

interpreted as an effective relaxation time that physically 

reflects the response of a nonlinear creep process to stress 

changes resulting from the earthquake. Our second set of 

time-dependent models employ a nonlinear rheology that 

represents a steady-state creep version of the laboratoryderived rate- and state-dependent friction law (Stesky , 

1975, 1978; Dieterich 1981; Ruina, 1983; Blanpied and 

others, 1991). With this law the aseismic slip rate depends exponentially on the ratio of shear stress to effective normal stress. 

Using either viscoelastic rheology, we treat the static 

coseismic stress change and resulting deformation as 

where K and p are the elastic bulk and shear moduli, 

respectively. We adopt the von Mises formulation and 

write the flow law in terms of the second invariant of the 

deviatoric stress, J2 = S,S, 12. The Mises equivalent tensile stress, q, is defined so that, in a state of uniaxial 

tension, q is equal to the tensile stress: 

The creep law is then written as 

where &(q) is the scalar Mises equivalent tensile creep 

rate, defined so that ecr(q) = Â£ in response to a uniaxial 

load ai, =q. 
D264 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

LINEAR VISCOELASTIC MODELS 

In the case of linear viscoelasticity flq) = q/A, where 

A is a constant. In fluid mechanics, it is conventional to 

define a shear viscosity, q, such that in a state of pure 

shear T = qy, where T is the shear stress and y is the 

engineering-shear-strain rate. We adopt this convention 

so that A=3q and 

The linear creep law as employed in our finite element 

calculations is, therefore, 

The viscous resistance to shear across a thin layer depends on the ratio of viscosity to layer thickness, q/h. 

Dimensional analysis, therefore, tells us that the timescale for relaxation of the system will be proportional to 

where H is the effective length-scale of the elastic surroundings and t is the relaxation time of the viscoelastic 

material. In Models 1 and 2, H corresponds to the downdip dimension of the coseismic rupture while in Model 3, 

H corresponds to the thickness of the elastic portion of 

the crust. In our finite element calculations we treat H as 

a known quantity. By comparing model predictions to data 

we can, therefore, hope to constrain the ratio t/h but neither t or h independently. We include a 1-km-thick fault 

zone in our finite element models, and in Model 3 additionally include a 5-km-thick basal shear zone. We then 

make t/h uniform throughout the model by specifying 

that the viscosity of the material in the basal shear zone is 

five times that of the material in the aseismic 'portion of 

the fault zone. By this procedure, time is measured in 

units of tr kdh and t,. and /z are the unknown material 

relaxation time and shear-zone thickness in the earth. 

By comparing computed displacements to those observed geodetically we can hope to equate model-time atr 

km/h with time t and by doing so constrain t/h. Rewriting 

this expression, we obtain 

which demonstrates how the parameters used to constrain 

viscosity will trade off. Note that if (t/a)(h/km) is of order one year and p=30 GPa, then q=lol* Pa-s. In what 

follows we examine solutions at a=5 and a=20 and geodetic observations at t= 1.3 years. 

In figure 6 we plot calculated and observed postseismic 

displacements measured relative to extrapolation of the 

pre-earthquake velocity field. The displacements are plotted along two profiles that lie perpendicular to the strike 

of the model fault zone (fig. 4). The first profile crosses 

through the center of the rupture and is referred to as the 

Loma Prieta profile. The second lies 44 km to the northwest and is referred to as the Black Mountain profile. 

These two profiles correspond approximately with two 

profiles of GPS stations measured frequently since the 

earthquake (Burgmann and others, 1991, and this chapter; 

Lisowski and others, 1991a; Burgmann and others, 1992; 

Savage and others, 1994). We plot three components of 

displacement from Models 1, 2, and 3 (fig. 2) at two 

model times, t=5t kmlh in figure 6A and t=20tr kmlh in 

figure 65. 

At t=5t kmlh the computed displacements are nearly 

equal for the three models (fig. 6A). This indicates that at 

short-time all three models are dominated by aseismic 

fault slip in the down-dip portion of the fault zone confined above the Moho. At t=20tr kmlh the computed displacements for the three models differ to a larger degree 

(fig. 6B). Therefore, if one of these models is correct for 

Loma Prieta at long-time, there is some hope of resolving 

which one it is. 

Also plotted in figure 6 are measures of the observed 

postseismic displacement field 1.3 years after the earthquake, made relative to extrapolation of the pre-earthquake velocity field. The bulk of the pre-earthquake 

geodetic data are from ground-based laser ranging measurements (EDM) and so do not constrain the rigid body 

component of the pre-earthquake velocity field. In contrast, all of the post-earthquake data are from GPS measurements in which the rigid body motion is constrained 

by making simultaneous observations to remote stations. 

As of this writing, though there are some pre-earthquake 

GPS observations, there is no self-consistent estimate of 

the change in the velocity field-post-earthquake minus 

pre-earthquake-based on all of the geodetic data, and so 

we have plotted the data as we describe in the following 

paragraphs. 

The data plotted along the Loma Prieta profile represent the difference between estimates of post- and preearthquake velocity obtained by assuming that the velocity 

field far to the northeast of the San Andreas fault did not 

change (Lisowski and others, 199la, 1991b). This measure of the change in velocity is then multiplied by 1.3 

years to obtain an estimate of the postseismic displacement measured relative to extrapolation of the pre-earthquake velocity. The error bars assume no error in the 

estimate of the pre-earthquake velocity field and thus underestimate the actual uncertainty. 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUAKE D265 

A 

LOMA PRIETA PROFILE 

100Ll rlQ[rlql, 

BLACK MOUNTAIN PROFILE 

DISTANCE NORTHEAST OF FAULT TRACE (KILOMETERS) 

LOMA PRIETA PROFILE 

100L1, I I ,, - 1, 

(.. ..... $ Model 1 .............. Model 2 

3 

50 bLp ~odel - 3 .- 

0 v"^..... .................. 

0 

BLACK MOUNTAIN PROFILE 

DISTANCE NORTHEAST OF FAULT TRACE (KILOMETERS) 

Figure 6.-Profiles of observed and calculated postseismic displacements measured relative 

to extrapolation of the pre-earthquake velocity field. The Loma Prieta and Black Mountain 

profiles lie perpendicular to the strike of the San Andreas fault. The former passes through the 

center of the model rupture and also through the Loma Prieta epicenter, while the latter lies 44 

km to the northwest along strike (see fig. 4). Observed displacements are based on geodetic 

measurements and correspond to 1.3 years after the earthquake (Lisowski and others, 1991a, 

1991b; Biirgmann and others, 1992). Calculated displacements represent versions of Models 

1, 2, and 3 (solid, dotted, and dashed curves, respectively) in which the relaxing portions of 

the models are made linear Maxwell viscoelastic (see fig. 2). Figures A and B correspond to 

model times of 5tr kdh and 20t kdh respectively, where t is the Maxwell relaxation time 

of the viscoelastic material, and h is the thickness of the viscously deforming region. 
D266 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Analysis of the Black Mountain data indicates that the 

change in velocity there-post- minus pre-earthquakemay not be measurably different from zero (Burgmann 

and others, 1992). Lacking a better estimate of the relative postseismic deformation, we plot zero displacement 

along the Black Mountain profile to indicate that the relative postseismic displacement is likely to be small. Along 

with the null data, we also plot an estimate of the uncertainty in the displacement accumulated during the first 

1.3 years after the earthquake. We estimate this uncertainty by taking the uncertainty in the velocity measured 

during the post-earthquake interval January 1990 to May 

1992 (Burgmann and Segall, written commun., 1993) and 

multiplying by 1.3 years. The resulting error bars place a 

lower bound on the detection threshold of postseismic 

deformation relative to the pre-earthquake rate. Relative 

postseismic deformation smaller than these error bars 

would go unnoticed in 1.3 years. 

At t=5t krnlh the predicted fault-trace-normal components of displacement are inconsistent with the data along 

the Loma Prieta profile, but along the Black Mountain 

profile they may be small enough to satisfy the data. If we 

were to add about 0.3 ppm or 0.2 ppdyr of regional 

fault-normal compression to our model, then we could 

probably match the data along the Loma Prieta profile but 

then probably violate the Black Mountain data. The predicted strike-slip components of displacement appear to 

be remarkably consistent with the data along the Loma 

Prieta profile but may not be consistent with the observations along the Black Mountain profile, where the computed displacements appear to exceed the observations. 

The predicted vertical displacements along the Loma Prieta 

profile may be sufficiently large to be detected by GPS 

measurements, whose precision in the vertical is about 20 

to 30 mm (Davis and others, 1989; Burgmann and others, 

this chapter). The predicted vertical displacements along 

the Black Mountain profile are extremely small. The predicted tilts along the Loma Prieta profile exceed 1 pradian and so could potentially be detected by leveling 

measurements. 

At t=20tr kdh, the predicted displacements far exceed 

all of the observations, and so we can say with confidence 

that 20t kdh > 1.3 yr or t/h > 0.07 yrlkm. 

If we accept the agreement at t=5t kmlh between predicted and observed strike-slip displacements along the 

Loma Prieta profile while overlooking the disagreement 

with fault-trace-normal motion there as well as possible 

disagreement with strike-slip motion along the Black 

Mountain profile, then we can conclude that 5t kdh = 

1.3 yr or t/h = 0.3 yrlkm. 

If we assume that postseismic slip is occurring within a 

fault zone with h 5 1 km and take p30 GPa, we then 

obtain an estimate of the effective viscosity of the deep 

aseismic portion of the fault zone; q S 3x1017 Pa-s. eff. This value is lower than any obvious interpretation of 

laboratory measurements (Carter and Tsenn, 1987; Kirby 

and Kronenberg, 1987a, 1987b) of solid-state creep of 

crustal rocks at mid-crustal conditions (Lachenbruch and 

Sass, 1973). 

Li and Rice (1987) and Fares and Rice (1988) studied 

two-dimensional earthquake cycle models by comparing 

model calculations to the observed decay in strain rate 

that followed the 1906 San Francisco earthquake (fig. 1). 

Their physical model is equivalent to our Model 3 except 

that they assumed that the aseismic down-dip portion of 

the fault zone is freely slipping. As discussed previously, 

this assumption is equivalent to assuming that t/h in the 

aseismic portion of the fault zone is much much less than 

in the lower crust, that is (t,/h)^ << (t/h)^ It is likely 

that this assumption only affects their solutions at shorttime following the earthquake and does not strongly affect their estimate of (t/h),. They obtained values in the 

range 0.2 yrlkm <: (t/h)lc 5 1.7 yrlkm. The contemporary 

velocity field in the northern San Francisco Bay region 

(Prescott and Yu, 1986) requires values at the low end of 

this range (Fares and Rice, 1988). 

We, therefore, conclude that t/h may be approximately 

uniform in the deep aseismic portion of the fault zone and 

in the lower crust and appears to take on a value of about 

0.3 yrlkm thickness. The apparent spatial uniformity oft/ 

h indicates that the proposed deep postseismic slip that 

occurred in the 1.3 yr following the Loma Prieta earthquake and the basal shear that followed the 1906 earthquake might be controlled by the same physical process. 

We interpret the apparently consistent values of (t& 

and (t/h), and the corresponding low value of the effective viscosity of the material in the deep aseismic portion 

of the fault zone as indications that postseismic relaxation 

may result from frictional sliding assisted by elevated pore 

pressure. This process may be taking place both in the 

deep portion of the fault zone and in horizontal shear 

zones in the lower crust. In the next section of this paper 

we evaluate this proposition by examining models of 

postseismic deformation that incorporate a constitutive law 

for frictional sliding. 

NON-LINEAR VISCOELASTIC MODELS: 

HOT-FRICTION 

We now represent the aseismically shearing regions with 

a nonlinear creep law derived from a steady-state version 

of the laboratory-derived rate- and state-dependent friction law (Stesky, 1975, 1978; Dieterich, 198 1; Ruina, 1983; 

Blanpied and others, 1991). At steady-state the resistance 

to sliding is described approximately by 

where V is the slip rate, c is a constant, a is the effective 

normal stress, and T, is the shear stress when V=Vo. We 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUAKE D267 

interpret T as the pre-earthquake shear stress across the 

aseismic portion of the fault and, in the case of Model 3, 

across the basal shear zone in the lower crust, a as the 

effective normal stress equal to total normal stress minus 

pore pressure, V as the slip rate across the shear zone, and 

V as the pre-earthquake slip rate. In using a steady-state 

form of the friction law, we implicitly assume that 

postseismic slip greatly exceeds the slip weakening distance and so the state variable remains always at its steadystate value, appropriate to the current slip rate. The factor 

c is more commonly denoted as a - b (see, for example 

Blanpied and others (1991)). While we use the friction 

law to describe hot frictional slip occurring deep within 

the crust, the same law has been used in studies of shallow aseismic creep on near-surface portions of a fault 

zone (Wesson, 1988; Marone and others, 1991). 

We make our calculations relative to the pre-earthquake 

state and define new parameters in that context. The friction law becomes 

where ? and f are the shear stress and slip speed measured relative to the pre-earthquake state, t = T - T, and A 

V = V- V.,. In a narrow shear zone, a state of pure shear 

will dominate the deviatoric stress tensor (Rice, 1992). 

Therefore, from the previous section, q= &? and 

Â£c = y / T/3 = VI ah, where q and eCr are now the Mises 

equivalent tensile stress and strain rate measured relative 

to the pre-earthquake state, and h is, as before, the thickness of the shear zone. Inverting the friction law we thus 

obtain 

so that the nonlinear creep law employed in our finite 

element calculations is 

We refer to the above formulation as our "hot-friction" 

model. 

Dimensional analysis tells us that the general form of 

the post-seismic slip rate across the aseismically slipping 

regions, measured relative to the pre-earthquake rate, will 

have the form 

where t is the time since the earthquake, AT is a representative coseismic static-stress change in the creep zone, Vo 

is a representative pre-earthquake slip rate, and x is 

position. The time-scale for postseismic relaxation is 

now governed by the quantities caH/pVO, Adca, and 

XLH. 

In our finite element calculations we assume values of 

H, such as rupture depth in the case of Models 1 and 2, 

and thickness of the elastic portion of the crust in the case 

of Model 3, just as with the previous linear models. We 

additionally assume that p=30 GPa, as in all previous 

analyses. To constrain V we have studied twodimensional earthquake cycle models in the manner of Li 

and Rice (1987) and Fares and Rice (1988) but using the 

steady-state friction law. Preliminary analyses in which 

we compare the decay in strain rate to that observed after 

the 1906 earthquake (fig. 1) indicate that 0.1 5 V/Vpiale 2 

1.0 in the deep aseismic portion of the fault zone within 

the crust, where Vph is the average plate velocity. This 

range is consistent with values obtained by Tse and Rice 

(1986) and by Rice (1993), in which they incorporate the 

full rate- and state-dependent constitutive law throughout 

the crustal fault zone and thereby model both stable and 

unstable fault slip throughout the earthquake cycle. We 

use Vo=VPlate=20 mmlyr, which is representative of estimates of the average slip rate along this segment of the 

San Andreas fault (Working Group, 1990). In our calculations, smaller values of V would yield smaller values of 

postseismic slip rate since there is a direct trade-off between V and our measure of time. Finally, laboratory 

measurements indicate that at mid-crustal conditions of 

elevated temperature with pore fluids present, 0.010 2 c 2 

0.020 (Blanpied and others, 1991). These constraints on 

H, p, V, and c, together with comparisons of model predictions and geodetic data, provide us a means of constraining the effective normal stress, a. 

We analyze Models 1, 2, and 3 (fig. 2) while incorporating the hot-friction constitutive law in the deep aseismic 

portion of the fault zone and lower crust. In figure 7 we 

plot profiles of postseismic displacement analogous to 

those in figure 6. The solution plotted is for Model 1 with 

ca=0.5 MPa. Larger values of CCT yield greatly reduced 

postseismic slip rate and so also reduced displacement at 

the free surface. This observation is consistent with inspection of the creep law and with our dimensional analysis. The former indicates that the initial postseismic slip 

rate will scale exponentially with the ratio Adca. Models 

2 and 3 yield displacement profiles nearly identical to 

those in figure 7 at times less than or equal to 2 years. 

Only at greater times do the surface displacements from 

the three models differ substantially. 

Perhaps the greatest difference between the response of 

the nonlinear hot-friction model and the linear viscoelastic model is their time dependence. In figure 8 we plot the 

strike-slip components of displacement, from nonlinear 
LOMA PRIETA PROFILE BLACK MOUNTAIN PROFILE 

DISTANCE NORTHEAST OF FAULT TRACE (mLOMETERS) 

Figure 7.-Profiles of postseismic displacement analogous to those in figure 6, but for is in the range 0.010 to 0.020 (Blanpied and others, 1991). As in figure 6, the observed 

a version of Model 1 (fig. 2C) in which the viscoelastic material in the deep aseismic displacements correspond to 1,3 years after the earthquake. The relative agreement 

region of the fault zone obeys the nonlinear hot-friction law. The solution plotted is between the model predictions and the data indicates that the effective normal stress in 

for c0=0.5 MPa, where o is the effective normal stress, equal to the total normal stress the deep aseismic portion of the fault zone is on the order of 0.5 MPdO.015 = 30 MPa. 

minus the pore pressure. In laboratory measurements made at mid-crustal conditions, c 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSmR ASSOCIATED WITH THE EARTHQUAKE D269 

and linear versions of Model 1, of a point lying on the 

Loma Prieta profile 13 km southwest of the model fault 

trace (see figs. 2C, 4, 6, and 7). We again use co=0.5 

MPa in the nonlinear model, as in figure 7, and then choose 

the time-scale for the linear model so that the displacement at 2 years equals that of the nonlinear model. The 

nonlinear hot-friction model yields very rapid postseismic 

motion immediately after the earthquake in comparison to 

the linear model. This very rapid short-term motion results principally from rapid postseismic slip that occurs 

near the edge of the coseismic rupture, where AT is large 

compared to co. Given time, this highly stressed region 

of the model relaxes to the degree that ? < co. After that 

point, the nonlinear creep law behaves more or less linearly, since the exponential dependence on shear stress 

becomes approximately linear when ? << co. 

The measured postseismic velocity field appears to be 

time independent during the initial two post-earthquake 

years (Burgmann and Segall, 1991; Lisowski and others, 

1991a; Burgmann and others, 19921, but the uncertainty 

in the data is sufficiently large that rejection of the nonlinear model on the basis of its strong time-dependence 

may not be possible. Future examination of the data is 

required to establish the degree to which time-dependence 

is allowed. Note that in figure 8, the initial 9 mm of 

predicted postseismic displacement occurs within 2 days 

of the earthquake. 

In figure 9 we plot contours of Mises stress in the deep 

aseismic portion of the fault zone. Again, we examine 

Model 1 with co=O.S MPa as in figures 7 and 8. Figure 

9A shows the static coseismic state, and figure 9B shows 

the state after 2 years of relaxation. The static coseismic 

Mises stress is less than 1 MPa beyond a distance of 

about 7 km from the edge of the rupture. Therefore, it is 

only within a narrow region close to the edge of the rupture that the creep law behaves strongly nonlinearly, since 

it is only here that ? is ever greater than co=OS MPa. 

Note the tendency for development of a relaxation front 

that propagates from the edge of the rupture into the creep 

zone as relaxation proceeds. 

If the friction law that we have examined applies to the 

earth, then immediately following an earthquake very rapid 

postseismic slip at depth occurs wherever AT > co. As a 

result, estimates of coseismic deformation based on geodetic data may be contaminated by postseismic deformation whenever the first post-earthquake survey lags too 

far behind the time of the earthquake. In figures 7 and 8, 

for example, over half of the displacement that accumulates in 1.3 years occurs during the first month following 

the earthquake. Note, however, that the displacements at 

issue here are small compared to the coseismic displacement. 

Returning again to figure 7, we compare the model 

predictions to the geodetic data. Along the Loma Prieta 

0.0 0.5 1 .O 1.5 2.0 

TIME SINCE EARTHQUAKE (YEARS) 

Figure 8.-Time histories of displacement of the free surface for nonlinear versus linear versions of 

Model 1 (fig. 2C). The nonlinear hot-friction model used ~~0.5 MPa, as in figure 7. The time-scale for 

the linear model was chosen so that the displacement at two years equals that of the nonlinear model. 

The point for which displacement is plotted lies along the Loma Prieta profile 13 km to the southwest of 

the fault trace. Note that with the nonlinear hot-friction model, the displacement at 1 month after the 

earthquake is nearly one-half of the total that accumulates by the end of 2 years. 
D270 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

profile, the hot-friction model does reasonably well in 

comparison to the observed strike-slip component of displacement but does not generate displacements compatible with the observed fault-trace-normal displacements. 

The addition of a uniform fault-trace-normal component 

of compression would bring the predictions toward agreement with the data along the Loma Prieta profile. Along 

the Black Mountain profile, the computed strike-slip displacements appear to exceed the observed values, but, as 

discussed previously, the error bars do not include the 

uncertainty in the pre-earthquake velocity field, nor do 

they address the unconstrained rigid body motion associated with the pre-earthquake EDM data. All of these comparisons are much like those made with the linear models 

(fig. 6A). 

We can conclude that the effective normal stress in the 

deep aseismic portion of the fault zone is extremely low, 

provided that we accept the agreement between the predicted and observed strike-slip displacements along the 

Loma Prieta profile while overlooking the disagreements 

both with fault-trace-normal motion there as well as with 

strike-slip motion along the Black Mountain profile. If we 

accept that c0=0.5 MPa and take c=0.015 from the laboratory measurements (Blanpied and others, 19911, we then 

obtain an estimate of the effective normal stress in the 

deep aseismic portion of the fault zone; 0=33 MPa, This 

MISES STRESS (MEGAPASCALS) 

Figure 9.-Contours of Mises stress in the deep aseismic portion of the fault zone for the nonlinear hot-friction version of 

Model 1 (fig. 2C) using c0=0.5 MPa. The contour interval is 0.1 MPa and the maximum contour level is 1.2 MPa. A, Static 

coseismic stress state. B, Stress state after 2 years of relaxation. Note the development of a relaxation front that propagates 

from the edge of the rupture into the creep zone as relaxation proceeds. 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUA= D27 1 

value is a factor of 10 smaller than the effective overburden stress that would exist at 18 km depth if pore-pressure was hydrostatic. Therefore, if deep aseismic slip is 

the cause of the observed postseismic strike-slip deformation near Loma Prieta and if hot-frictional shearing is the 

mechanism by which this process takes place, then pore 

pressure in the deep aseismic portion of the fault zone 

must be substantially elevated above the hydrostatic level. 

DISCUSSION AND CONCLUSIONS 

ALTERNATIVE GEOMETRIC MODELS 

We have demonstrated that models of postseismic deformation that include deep sources of viscous relaxation 

can generate displacement fields that agree with some, 

but not all, of the geodetic data. 

Our models appear to be capable of producing strikeslip displacements that are compatible with the observations along the Loma Prieta profile but may exceed those 

made along the Black Mountain profile. We have examined one model of postseismic deformation in which 

coseismic slip is non-uniform. The goal was to establish 

the degree to which postseismic strike-slip displacement 

along the Black Mountain profile could be suppressed by 

moving the centroid of coseismic strike slip away from 

Black Mountain. The non-uniform slip model is motivated 

by analysis of leveling data (Marshall and others, 1991) 

and by seismological estimates of the distribution of 

coseismic slip (Beroza, 1991; Hartzell and others, 199 1 ; 

Steidl and others, 1991; Wald and others, 1991). We have 

examined an extreme model in which all of the strike slip 

is in the southeast half of the rupture and all of the reverse slip is in the northwest half of the rupture. The 

moment and average rake are as in the uniform slip models. The non-uniform slip model employs the linear viscoelastic rheology in the deep aseismic portion of the fault 

zone and so falls into the category of Model 1 (fig. 20. 

As expected, the non-uniform model does suppress 

postseismic strike-slip displacement along the Black Mountain profile in comparison to the corresponding uniformslip model. At t=5tr kdh the strike-slip displacement 

along the Black Mountain profile is about 70 percent of 

that obtained with the uniform slip model and plotted in 

figure 6A. However, this reduction is more than compensated for by an accompanying reduction in strike-slip displacement along the Loma Prieta profile. To match the 

data there we must now choose 1.3 yr = t = 20tr kdh, 

which corresponds to larger strike-slip displacements along 

the Black Mountain profile than those obtained with the 

uniform slip model at t=5t,. kmlh. Thus, the non-uniform 

coseismic slip model does not help to suppress postseismic 

strike-slip displacements along the Black Mountain profile relative to along the Loma Prieta profile. 

Our models also appear to be incompatible with the 

fault-trace-normal contraction observed along the Loma 

Prieta profile. This discrepancy could be suppressed by 

adding a uniform fault-trace-normal compressive strain to 

our models of magnitude about 0.3 ppm, or 0.2 ppdyr. 

This strain would be interpreted as a regional postseismic 

change in fault-trace-normal strain rate, though we are 

not aware of any independent data that reveals such a 

feature. The addition of this regional fault-trace-normal 

compression would, furthermore, make our models less 

consistent with the data from the Black Mountain profile, 

if the change in velocity along the Black Mountain profile 

is indeed not measurably different from zero (Burgmann 

and others, 1992). 

We conclude that alternative models of postseismic deformation should be examined in future work, Two overlapping viewpoints should be considered. First, our 

deep-slip models do a remarkably good job of matching 

the strike-slip component of postseismic displacement 

along the Loma Prieta profile, but they do not appear to 

be capable of generating the observed fault-trace-normal 

contraction there. Therefore, future models should include 

similar deep-slip sources, while additionally incorporating sources that contribute a greater degree of fault-tracenormal contraction. Second, our deep-slip models may 

generate strike-slip displacements along the Black Mountain profile that exceed the observed displacements. Therefore, future work should also examine models that 

incorporate sources of deformation that are shallower than 

our deep-slip sources. By moving the deformation sources 

toward the free surface the width of the postseismic displacement field will be narrowed relative to those generated with our deep sources, This narrowing of the 

displacement field will make it easier to satisfy the data 

from the Black Mountain profile, if indeed the current 

lack of postseismic signal holds up to further data analysis. A planned re-leveling across the Loma Prieta profile 

(W.H. Prescott, written commun., 1993) may also help to 

resolve the depth of the postseismic deformation source. 

Burgmann and others (1992) have performed trial-anderror searches to determine the location of the rectangular 

dislocation with uniform slip that best fits the geodetic 

data. They conclude that the postseismic source probably 

lies at depths above the 18-km-deep Loma Prieta hypocenter. They further conclude that postseismic slip occurred not only in the Loma Prieta aftershock zone but 

also on a shallow thrust fault lying north of Loma Prieta. 

If their model is correct, then perhaps postseismic slip in 

the aftershock zone represents stress relaxation between 

patches that slipped seismically in the earthquake. This 

interpretation is similar to our deep-source models, but 

now creeping patches lie within the seismogenic zone and 

presumably represent aseismic "non-asperity" regions. This 

view is in accord with the interpretation that aftershocks 

copcentrate in regions that have relatively small coseismic 
D272 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

slip (Oppenheimer and others, 1990) and are therefore 

most severely stressed by the earthquake. Analyses like 

ours could be performed by including shallow creeping 

patches within the seismogenic zone, whose location is 

constrained by seismologic estimates of the slip distribution in the earthquake. The shallow thrust fault proposed 

by Burgmann and co-workers may be confined to rather 

shallow depths and consequently may represent stable fault 

creep (Tse and Rice, 1986; Scholz, 1990) of the sort 

examined by Marone and others (1991). Such a model 

could be examined within the framework of our finite 

element analyses or through use of the boundary element 

method. 

COMPARISON TO STRAIN-METER 

OBSERVATIONS AT SAN JUAN BAUTISTA 

The strain rate within a few kilometers of the surface 

trace of the San Andreas fault near San Juan Bautista 

(SJB) has been monitored by a Gladwin borehole tensor 

strain meter since 1983. The observations surrounding the 

time of the Lorna Prieta earthquake are summarized here 

(Gwyther and others, 1992). Approximately one year prior 

to the Loma Prieta earthquake, the right-lateral strain rate 

increased to about 1 pradianlyr. Immediately following 

the earthquake, a 2-month transient occurred in which the 

right-lateral strain rate actually became negative and about 

0.5 pradian of left-lateral strain accumulated. Following 

that short transient, the strain rate increased to about 2 

pradianlyr, right-lateral, and has remained at that level for 

about 2 years. These changes in strain rate presumably 

reflect variations in slip rate on the fault that may be 

associated with the earthquake. 

While our models do not attempt to include the effects 

of the shallow San Andreas creep zone that extends 150 

km to the southeast from SJB, it is still interesting to 

compare our model predictions to the data. In the case of 

our nonlinear hot-friction models, the right-lateral strain 

rate near SJB is very high immediately following the earthquake but rapidly decays to about 0.4 pradianlyr above 

the pre-earthquake rate when averaged over the interval 

0.5 to 2.0 postseismic years. This temporal behavior is 

qualitatively consistent with the time history of displacement plotted in figure 8. None of our models predict an 

immediate postseismic transient in which left-lateral strain 

accumulates in the vicinity of SJB. Furthermore, the predicted rate of postseismic strain measured relative to the 

pre-earthquake rate is only 40 percent of the observed 

value. The former is about 0.4 pradianlyr, while the latter 

is about 1 pradianlyr. We conclude that models that attempt to include the effects of the shallow San Andreas 

creep zone are required to obtain better agreement with 

the data from SJB. 

ALTERNATIVE PHYSICAL PROCESS 

We have proposed that postseismic deformation may 

result from frictional sliding assisted by elevated pore pressure. Alternatively, postseismic transient deformation may 

result from transient creep (Carter and Kirby, 1978; B. 

Evans, oral commun.). The displacement across our deep 

aseismic fault zone, required to match the observed strikeslip displacement along the Loma Prieta profile at 1.3 

years, is about 0.5 m at a distance of 2.5 km from the 

edge of the coseismic rupture. If this deformation occurs 

within a shear zone with thickness h 5 1 km, then the 

corresponding transient strain is ytr 2 0.5 x 10-~and its average rate is y tr 2 2 x lo-'' I s. 

SUMMARY 

We have demonstrated that postseismic relaxation can 

increase the stress change at intermediate distances from 

the edge of the coseismic rupture by an amount well in 

excess of the coseismic change. Postseismic relaxation 

can have this same effect in the far-field, but the magnitude of the total stress change there is very small compared to typical earthquake stress drops. Along the 

Hayward and Calaveras fault trends, northeast to north of 

Loma Prieta, both coseismic slip and postseismic relaxation reduce the right-lateral shear stress and the compressive stress. 

Our models of deep aseismic fault slip occurring beneath the seismogenic zone are capable of producing strikeslip displacements consistent with the initial 1.3 years of 

observed postseismic deformation along the profile of GPS 

stations that crosses through the epicenter (Lisowski and 

others, 1991a). However, these models appear to be incapable of producing fault-trace normal contraction compatible with the observations there, and an additional source 

of deformation may be required. For example, Burgmann 

and others (1992, and this chapter) investigate models 

that include postseismic fault slip on shallow thrust faults 

north of Loma Prieta, while Savage and others (1994) 

consider models that include postseismic collapse of the 

coseismic rupture zone in the direction perpendicular to 

the plane of the rupture. Linker (1993) has shown that the 

magnitude and time-scale of Savage and co-workers' proposed fault-trace-normal collapse may be consistent with 

a relaxation time-scale controlled by postseismic flow of 

pressurized fluid out of the fault zone and into the country rock. 

Along the Black Mountain GPS profile, northwest of 

Loma Prieta, postseismic displacements predicted by these 

same deep-slip models may exceed the observations, 

though a careful treatment of all of the pre-earthquake 

and post-earthquake geodetic data is required to better 
MODELS OF POSTSEISMIC DEFORMATION AND STRESS TRANSFER ASSOCIATED WITH THE EARTHQUAKE D273 

establish a measure of the change in the velocity field 

associated with the earthquake (see Biirgmann and others, 

this chapter). 

The time-scale for postseismic relaxation appears to be 

sufficiently long that the average background tectonic 

stress rate is likely to exceed the stress rate resulting from 

relaxation, except perhaps close to the edge of the 

coseismic rupture. If the time-scale for relaxation of the 

lower crustal detachment zone turns out to be on the order 

of 16 years or less, then its contribution to the stress rate 

within about 20 krn of the end of the Loma Prieta rupture 

will be at least comparable to the average background 

stress rate. 

In our linear viscoelastic models, the parameter controlling the relaxation time is the ratio of material relaxation time to the thickness of the deforming region, t/h. 

This parameter appears to have a value of approximately 

0.3 yrlkm thickness, at least in the deep aseismic portion 

of the fault zone. If relaxation takes place in a shear zone 

that is less than 1 km thick and if the shear modulus is 30 

GPa, then the effective viscosity in the deforming region 

must be less than about 3x10~~ Pa-s. This value is at least 

one order of magnitude lower than any obvious interpretation of laboratory measurements of the steady-state creep 

of solid crustal rocks at mid-crustal conditions. 

The laboratory-derived hot-friction model can yield 

postseismic deformation of magnitude comparable to that 

observed, but only if ddd(1n V) is on the order of 0.5 

MPa, where T is the resistance to sliding and V is the slip 

speed. Laboratory measurements indicate that ddd(1n V) 

can be written approximately as c<7 (Stesky, 1975, 1978; 

Dieterich, 1981; Ruina, 1983), where <7 is the effective 

normal stress equal to total normal stress minus pore pressure and c is on the order of 0.015 for mid-crustal conditions (Blanpied and others, 1991). Therefore, to be 

consistent with our models, the effective normal stress in 

the deep aseismic portion of the fault zone must be extremely low, perhaps on the order of 30 MPa. 

In the context of the laboratory-derived hot-friction 

model, the occurrence of postseismic deformation may be 

evidence that pore pressure in the aseismic portions of the 

fault zone and perhaps the lower crust is near lithostatic. 

ACKNOWLEDGMENTS 

We wish to thank R. Biirgmann, M. Lisowski, J.C. Savage, and P. Segall for early access to the GPS data and for 

helpful discussions. We additionally thank B. Evans and 

R.W. Simpson for helpful discussions. P. Segall and R.W. 

Simpson provided careful reviews that led to correction 

of a sign error. This work was supported by the United 

States Geological Survey, National Earthquake Hazards 

Reduction Program, under grants G1844 and G1788 to 

Harvard University. Supercomputer services were provided 

by grants to Harvard University from the NSF National 

Center for Supercomputing Applications, Urbana Illinois. 

Some of the study was done through access to a network 

of workstations supported primarily by J. Bloxham, A.M. 

Dziewonski, G.A. Ekstrom, and R.J. O'Connell at Harvard. 

Hibbitt, Karlsson, and Sorensen, Inc. of Pawtucket, Rhode 

Island, provided ABAQUS through an academic license 

agreement. 

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THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKE OCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

A SHEAR STRAIN ANOMALY FOLLOWING THE LOMA PRIETA 

EARTHQUAKE 

By M.T. Gladwin, R.L. Gwyther, and R.H.G. Hart, 

University of Queensland, Australia 

CONTENTS INTRODUCTION 

Page 

D277 

277 

278 

28 1 

283 

283 

ABSTRACT 

Borehole tensor strain instruments deployed along the 

San Andreas fault for the past 10 years have provided 

sufficient resolution and stability to sample regional tectonic processes. Data obtained from an instrument at San 

Juan Bautista in the near field region of the Loma Prieta 

earthquake provide the first high-resolution continuous 

shear strain observations associated with a large earthquake. A change in fault-parallel shear-strain rate of approximately 1 pe per year occurred about a year prior to 

the earthquake and persisted to the time of the event. The 

strain rate decreased immediately after the earthquake, 

but following the Chittenden sequence of earthquakes in 

April 1989, a new and higher rate of fault-parallel shear 

accumulation (0.84 pe per year relative to the 1989 rate) 

was established. This strain rate has continued through 

1993. Associated creep-rate changes are apparent at a number of sites on the surface trace of the fault within 30 km, 

indicating that the measured change of strain rate at the 

time of the earthquake has regional significance. We propose that the observed strain accumulation results from 

increased slip around a nearby locked section of the fault 

arising from loading by the failed Loma Prieta source 

region to the north. This model is consistent with suggestions of an increased probability of a moderate earthquake 

near San Juan Bautista and with evidence that interactions 

between fault regions are important in earthquake processes. 

A Gladwin borehole tensor strainmeter (Gladwin, 1984) 

installed near the San Andreas fault at San Juan ~autista 

in late 1983 has provided continuous areal and shear strain 

data with sub-nanostrain resolution and long-term stability better than 100 ne per year (Gladwin and others, 1987). 

Raw data from the instrument consist of diameter changes 

in three directions at 120' to each other in the horizontal 

plane. These are reduced to areal strain E, and engineering shear strains fi and y2 (approximately parallel to and 

at 45' to the fault respectively), which are defined in terms 

of strain tensor components (cXx, E , and ex by 

where, as subscripts to e, the x direction is east and y is 

north. The strainmeter is grouted into the surrounding rock, 

and this instrument inclusion is softer to shear than to 

compression. Observed strain components are thus scaled 

by hole-coupling parameters (Gladwin and Hart, 1985) 

determined by tidal calibration. 

Data from borehole inclusions are initially dominated 

by grout compression of the instrument, by thermally 

controlled decay as the instrument site re-establishes 

equilibrium with its surroundings, and by an exponential 

recovery of the virgin stress field relieved at the borehole 

during the drilling process (Berry and Fairhurst, 1966; 

Berry, 1967). The exponential signals are characteristic of 

viscoelastic rheology as typified by Burghers solids (Jaeger and Cook, 1976). They have no relevance to the monitoring of regional strain changes and were removed by an 

exponential least-squares fitting procedure over the interval January 1984 to February 1988 (Gladwin and others, 

1991). 
D278 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Raw data from the three instrument gauges commencing in 1984 (3 months after installation) are shown in 

figure 1A. In determination of the exponentials to be removed from the raw gauge data streams, all data known 

to be contaminated were eliminated from the fit (Gladwin 

and others, 1991; Gwyther and others, 1992). The same 

intervals of data are used for determination of exponentials 

for all three gauges, and no linear trend has been removed. 

The resulting exponentials, determined from data during 

early 1984, and from May 1986 to April 1988, are shown 

also in figure 1A (offset for clarity). The data following 

the Morgan Hill earthquake in April 1984 and anomalous 

data associated with field experiments at the site in 1986 

were excluded. The instrument was off line for 6 months 

during 1987. All data after April 1988 were also excluded 

and provide no constraint on the least-squares fitting. Figure 1B gives the residuals from the determined 

exponentials. The residuals indicate stable gauge behavior from mid-1986, an emerging anomaly on all components beginning in late 1988, the strain offsets for the 

Loma Prieta earthquake in October 1989 and significantly 

differing behavior on all gauges since that time. The 

postseismic data indicate immediate postseismic recovery 

for about 3 months, and following the Chittenden earthquakes in April 1990 the establishment of new, relatively 

linear trends. 

These residuals were then reduced to the strains ca, yl, 

and y, shown in figure 2A. The data differ slightly from 

those presented in Gladwin and others (1991), owing to a 

refinement of the selection of data windows since that 

time. The dominant signals present are the coseismic strain 

offsets of the Loma Prieta earthquake. These have been 

documented elsewhere (Gladwin and others, 1991); the 

present discussion is confined to the observed strain-rate 

changes. In figure 2B the coseismic offsets have been 

removed from the data to make long term trends more 

apparent. 

It is important to investigate how choices made in estimating and removing the exponential borehole response 

can influence interpretation of the strain data determined 

from the gauge residuals. These effects are examined in 

figure 3. The representative strain data shown are produced from residuals obtained for three different data windows marked a, b and c used in the fitting procedure. We 

are documenting an apparent change of strain rate at the 

time of the Loma Prieta earthquake. The critical issue is 

whether the choice of data interval significantly affects 

determination of the strain-rate change. It is evident from 

figure 2 that any reasonable choice of fitted window demonstrates that a gradient change occurred in the raw 

datasets. For the extremes of fitting intervals shown, the 

effect on the observed change of strain rate before and 

after the Loma Prieta earthquake are shown in table 1. 

Interval c is inappropriate because it extends into the 

data which is to be used to determine the pre-earthquake 

gradient. In the following discussion, interval b has been 

used because it gave the best variance over the available 

1987 dataset. 

RESULTS 

An anomalous change in y1 is apparent by late 1988, 

showing a remarkably linear strain accumulation of 1 p~ 

per year relative to the pre-1988 rate (Gladwin and others, 

1991). The azimuth of maximum shear for the accumulating shear strain is approximately parallel to the local San 

Andreas fault strike. The long-term stability of the measurements is particularly evident in the areal strain data. 

Areal strain is estimated from the sum of the three components and is seen to be constant (with the exception of 

the coseismic offset at the Loma Prieta event) at better 

than 1 microstrain over the whole 10-year period. 

Immediately after the earthquake the fault-parallel shearstrain rate decreased for about 2 months and then gradually returned to the pre-anomaly value. These data are 

shown in figure 4. In May 1990, following the Chittenden 

aftershock sequence (four magnitude 4 to 5 earthquakes 

on April 18, 1990, centered on an area approximately 15 

km north west of SJT), the present linear yl shear accumulation rate of approximately 2 ,LK per year with the 

original sense had been established. 

The absolute strain rates may of course differ from those 

indicated on these figures (approximately 1 ,u& per year 

from late 1988 to the Loma Prieta event, and approximately 2 p~ per year relative to the pre-1988 rate from 

May 1990 through December 1993) because the data have 

been effectively detrended by the exponential removal 

procedure. The point at issue is that there are significant 

changes of strain rate documented in the data, one preceding the Loma Prieta earthquake and another following it. 

Both appear to be linear with time, and their relative magnitudes are not an artifact of the exponential removal procedure. 

An alternative and useful means for determining the 

physical significance of these data is to plot yl against y2 

(see fig. 5). The shear state at a particular time is represented by a point in this shear space, its history is represented by the locus of these points, and the shear required 

to move from one shear state to another is the vector 

Figure 1.-A, Raw gauge data for the SJT site beginning 3 months 

after installation. Fitted exponential curves are shown offset for clarity. 

The recording system was nonoperational for 5 months during 1987. B, 

Residual gauge data produced by removal of the fitted exponentials. 

Units are nominal microstrain only. 
A SHEAR STRAIN ANOMALY FOLLOWING THE LOMA PRIETA EARTHQUAKE 

e ........................................ -u.. -. ,,.. 

Gauge 3 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

!=! 

. IÃ‘ 

03 

1-1 GO 

0 4 

8 

z 

y< Areal ' 

\ /------ Gamma 2 

*^I B LEI LE2 
A SHEAR STRAIN ANOMALY FOLLOWING THE LOMA PRIETA EARTHQUAKE D28 1 

connecting these points. With the local trace of the San 

Andreas approximately at N48OW, the y; axis also represents approximately fault-parallel shear, and the fi axis 

approximate extension normal to the fault. Time is marked 

at 3-month intervals, and direction of the coseismic offset 

of the Loma Prieta event is indicated in the upper and 

lower parts of the figure by the arrow. The Chittenden 

aftershock sequence is marked as CH on the lower part of 

the figure. 

The shear-strain step C-C', coseismic with the Loma 

Prieta event, is seen to increase both right-lateral shear 

across the fault (arising mainly from the y, component, 

+I840 ne) and normal compression (arising mainly from 

the y, component, -3790 ne). This fault-normal compression from shear alone must be combined with the contribution due to the areal strain step (+2140 n~) to find the 

effective change in fault normal strain in the vicinity of 

the instrument of approximately 1600 ne. Apart from a 5- 

month period immediately prior to the Loma Prieta event 

(see B-C in the figure) during which the maximum shear 

accumulation was oriented at N56OW, the predominant 

trend of both anomalies (see sections A-B and D-E in the 

figure) is a shear vector at angle E6OS corresponding to a 

maximum shear accumulation at N42OW, close to the local San Andreas fault strike of N48OW. 

To verify instrument and coupling stability, response to 

earth strain tides was examined using the dominant, thermally uncontaminated tidal components 0 and Mr.. Sixtyday data windows of 90-minute data were used to provide 

normalised tidal component amplitude every 30 days. The 

strain step of the Loma Prieta event and other easily identifiable strain steps associated with earthquakes or creep 

events were removed from the record before the tidal 

analysis. Results are shown in figure 6 for the e and y, 

data sets. Error bars indicate the precision of determination for each point, assuming equal partition of noise over 

all tidal components. It is clear that there have been no 

significant or systematic changes of tidal admittance on 

this instrument over the whole period under discussion. 

4 Figure 2.-A, Areal strain and shear strains from the SJT borehole 

tensor strainmeter at San Juan Bautista near the San Andreas fault in 

northern California. Exponential trends have been removed from this 

data. A dominant feature is the coseisrnic strain step from the Loma 

Prieta earthquake in October 1989. B, Removal of this step reveals the 

details of the strain records, in particular the striking anomaly in the y, 

component, the relative constancy of the other two strain components, 

and trend reversal on y, for 3 months following the Loma Prieta event. 

All steps in the data can be associated with seismic events or nearby 

creep events. For example, the times of the two Lake Elsman earthquakes are indicated (LEI, LE2), as are the times of the Loma Prieta 

earthquake (LP), the Chittenden earthquake sequence (Ch), and creep 

events also monitored on a nearby creepmeter (0) and documented in 

table 2. 

This result has been confirmed elsewhere (Linde and others, 1991). 

DISCUSSION 

Apart from coseismic steps, the strain data show several readily identifiable steps which correlate closely with 

the main creep events on the closest creepmeter XSJ, 2 

km to the east of our instrument. Events from late 1988 to 

1992 are tabulated in table 2, and the detailed correspondence of the events with the main creep events gives further confidence that the data represent regional tectonic 

activity. Further, the events show (Gladwin and others, 

1993) remarkable similarity suggestive of a small source 

(less than 0.5 kilometers in depth and at most a few kilometers in extent) directly under San Juan Bautista. The 

shear strain resulting from such a source cannot account 

for the size of the post Loma Prieta strain anomaly. Other 

explanations for this anomaly need to be examined. 

A non-tectonic source from the instrument or its 

immediate vicinity is unlikely because of the stability of 

the tidal response, consistency of our internal instrument 

checks, the long-term stability of the areal strain record 

(better than 1 p~ over 10 years), and the detailed correspondence in time of all observed strain steps with either 

earthquakes or creep events on nearby creep meters. 

Though the strain signals could arise from small-scale 

processes in a nearby section of the fault, observations of 

anomalous creep events at three sites up to 30 km away 

indicate a more extended source. Figure 7 shows longterm creepmeter data for 5 sites covering 40 km of the 

fault south of San Juan Bautista with long-term trends 

(Burford, 1988). A distinct increase in creep rate following the Loma Prieta earthquake is evident on sites XSJ, 

XHR and CWC, spanning 16 km of the fault. The 

XFL site (29 krn from XSJ) shows only a marginal increase, and the more remote site XMR (40 km) shows no 

effect. 

The creep anomalies in figure 6 are unusual, especially 

for CWC and XHR, and begin at the time of the Loma 

Prieta earthquake. The creep anomaly at the nearby XSJ 

begins about the time of the establishment of the new 

shear strain anomaly at SJT, which, given the causal time 

correspondence, suggests that these signals are not just 

the consequence of normal interactions between fault sections in this creeping section, but are linked to the earthquake. We conclude that the failure of the Loma Prieta 

source region transferred load to the San Juan Bautista 

region just to its south, resulting in increased creep rate. 

The simplest explanation of an increased creep rate is 

frictional response to the increased fault-parallel shear 

loading indicated by the coseismic y, step at San Juan 

Bautista. 
D282 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

However, slip via creep does not itself result in linearly Prieta source region particularly after the Chittenden afincreasing elastic strain. We suggest that our linear shear- tershock sequence (April 1990). 

strain anomaly is best explained by continued aseismic The response at the site immediately after the Lorna 

slip around a nearby locked section of the fault, the slip Prieta earthquake and prior to the Chittenden events seems 

being associated with loading transferred from the Loma to indicate that this load transfer towards the SJT site was 

Figure 3.-Representative 7, shear-strain data illustrating the effect of the choice of data window used for the 

exponential fitting. The intervals used in three separate fit sets are marked as a, b, and c. Estimates of the change of 

strain rate from before to after the Loma Prieta event are only marginally affected for the three extreme choices 

shown. The choices have some effect on the timing of the pre-event anomaly only. Note that for choice c the 

change of gradient is well defined before the end of the fitting interval, indicating that the end point of the fit 

interval is not producing the effects discussed. The changes of gradient that determine the anomalies under discussion are clear on all residual sets. 
A SHEAR STRAIN ANOMALY FOLLOWING THE LOMA PRIETA EARTHQUAKE D283 

Table 1 .-Average gradient of shear strain y, determined over three periods 

[Values are residual gauge data after removal of exponential functions fitted over the 

three selection intervals 1985102 to 1988104 (a), 1986106 to 1988104 (b), and 1985105 

to 1989107 (c). The lowest row of data indicates that the change in strain rate after the 

Loma Prieta earthquake is very similar for the three fitting intervals shown.] 

Selection Interval 

1989 to L.P. 1.56 k .08 1.16 k .09 0.75 k .09 

Difference 0.84 + .13 0.85 & .14 0.79 k .15 

initially prevented by the strength of an asperity region 

associated with the Chittenden aftershock sequence. When 

this region ultimately failed in 1990, the SJT site immediately began to respond to the fault parallel shear expected 

from the Loma Prieta event in the sense of a continued 

afterslip at the Loma Prieta region. 

A range of simple dislocation models of slip around a 

locked region were considered. The model which resulted 

in a fault-parallel shear-strain rate most comparable to 

that observed was a locked region extending along the 

fault strike at depth of 1.5 to 5 km, with the surrounding 

fault surface slipping at the regional average of 14 mm 

per year. While this model indicated a locked region shallower than the study mentioned above, uncertainty in the 

absolute value of the observed y'y strain rate indicates that 

our data are broadly consistent with the presence of such 

a locked region at moderate depth on the fault. The data 

thus suggest an increased probability of a moderate earthquake in the San Juan Bautista region. 

We have previously suggested (Gladwin and others, 

1991) that the pre-Loma Prieta strain changes were related to a broad regional effect, based on the timing of the 

Lake Elsman foreshock and a marginal geodetic anomaly 

(Lisowski and others, 1990) near the Loma Prieta source. 

The effects we are reporting also appear to have some 

regional expression in increased creep activity since the 

Loma Prieta earthquake. The anomaly in yl is now so 

large that it would be expected to be detectable by geodetic observations in the area. Unfortunately, the major 

geodetic network is centered well to the north of San Juan 

Bautista, making even inversion for the determination of 

the southern extent of the Loma Prieta rupture zone itself 

difficult (Johnston and others, 1990). Recent data from 

the southern end of this network is not yet available, but a 

change of fault-parallel shear-strain rate of l/.~ per year 

coincident with the Loma Prieta event should easily be 

identified in the observational period to date if it extended 

over the total geodetic network involved. These data are 

critical in providing constraints on the scaling of the 

anomaly in the northerly direction. Some data has been 

taken in a small HP3808 network extending from San 

Juan Bautista north to Pajaro Gap, particularly following 

the earthquake (Johnston, oral comm., 1993), but the results from this network are not yet published. Data from 

this network was not taken for many years prior to the 

Loma Prieta earthquake, so there is no likely contribution 

on the issue of whether or not a change of gradient has 

occurred following the event. There is no associated areal 

strain anomaly indicated in the data, and so no corroboration of the anomaly by reference to the dilatometer at 

Searle Road, approximately 5 km away, is available. The 

dilatometer shows no significant areal strain change. 

Scaling of the anomaly cannot be determined from a 

single site, and the significance of these data will remain 

unknown until it is confirmed or denied from measurements at other nearby instrument sites. There is no evidence that the observed anomaly extends sufficiently far 

north to be measured in the Loma Prieta geodetic network. Our modelling suggests the presence of a source 5 

km long. This source would produce minimal deformation in the geodetic network. Hence, although we are confident that the observed anomaly is not an instrumental 

artifact or of very localized origin, the only currently available supportive evidence that this anomaly is of regional 

significance is the increased creep activity in stations south 

of SJT. 

ACKNOWLEDGMENTS 

This work was performed under a grant from the U. S. 

Geological Survey, Department of the Interior. Contents 

do not represent policy of that agency, and no endorsement of the agency is to be assumed. The instruments 

were previously developed under awards of the Australian Research Grants Scheme and fabricated by R. Willoby 

and staff in-house. We thank R. Liechti for maintenance 

support, K. Breckenridge for data retrieval and assistance 

with creepmeter data, and Drs. A. Linde and M. Johnston 

for general support in the program. 

REFERENCES CITED 

Berry, D.S., 1967, Deformation of a circular hole driven through a 

stressed viscoelastic material: International Journal of Rock Mechanics and Mining Science, V. 4, p. 18 1 - 187. 

Berry, D.S., and Fairhurst, C., 1966, Influence of rock anisotropy and 

time-dependent deformation on the stress-relief and high-modulus 

inclusion techniques of in-situ stress determination: American Society for Testing Materials, STP 402, p. 190-206. 

Burford, R.O., 1988, Retardations in fault creep rates before local 

moderate earthquakes along the San Andreas fault system, central 
D284 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

California: Pure and Applied Geophysics, v. 126, p. 499-529. 

Gladwin, M.T., 1984, High precision multi component borehole deformation monitoring: Reviews of Scientific Instruments, v. 55, p. 

201 1-2016. 

Gladwin M.T., Breckenridge, K.S., Hart, R. and Gwyther, R.L., 1994, 

Measurements of the strain field associated with episodic creep 

events on the San Andreas Fault at San Juan Bautista, California: 

Journal of Geophysical Research, v. 99, no. B3, p. 4559-4565. 

Gladwin, M.T., Gwyther, R.L., Hart, R. and Francis, M., 1987, Borehole tensor strain measurements in California: Journal of Geophysical Research, v. 92 , no. B8, p. 7981-7988. 

Gladwin, M.T., Gwyther, R.L., Higbie, J.W. and Hart, R., 1991, A 

medium term precursor to the Loma Prieta earthquake?: Geophysical Research Letters, v. 18, no. 8, p. 1377-1380. 

Gladwin, M.T., and Hart, R., 1985, Design parameters for borehole 

strain instrumentation. Pure and Applied Geophysics, v. 123, p. 59- 

88. 

Gwyther, R.L., Gladwin, M.T., and Hart, R.H.G., 1992, A shear strain 

anomaly following the Lorna Prieta earthquake: Nature, v. 356, 

p,142-144. 

Jaeger, J.C. and Cook, N.G.W., 1967, Fundamentals of rock mechanics: London, Chapman and Hall. 

Figure 4.-Detail of the strains immediately following the Lorna Prieta earthquake. The y, record shows that the 

post-earthquake strain rate shown in figure 2 did not begin until after the Chittenden sequence in April 1990. 
A SHEAR STRAIN ANOMALY FOLLOWING THE LOMA PRIETA EARTHQUAKE D285 

Johnston, M.J.S., Linde, A.T., and Gladwin, M.T., 1990, Near-field 

high resolution strain measurements prior to the October 18, 1989, 

Loma Prieta Ms 7.1 earthquake: Geophysical Research Letters, v. 

17, no. 10, p. 1777 - 1780. 

Linde, A.T., Gladwin, M.T., and Johnston, M.J.S., 1991, The Loma 

Prieta earthquake, 1989 and earth strain tidal amplitudes: an unsuccessful search for associated changes: Geophysical Research 

Letters , v. 19 no. 3, p 317-320. 

Lisowski, M., Prescott, W.H., Savage, J.C., and Svarc, J.L., 1990, A 

possible geodetic anomaly observed prior to the Loma Prieta, 

California, earthquake: Geophysical Research Letters, v. 17, no. 8, 

p. 121 1-1214. 

pre Lorna Prieta 

................ .................................................... 

y, Microstrain 

-..... LeP.......;....................:....................;.,.,....*..,.....,.. ^ : post Lorna prieta 

yi Microstrain 

Figure 5.-Pre- and post-Loma Prieta plots of tensor strain components The initiation and completion of the step associated with the Loma 

y, against yl in pe for the period July 1986 to April 1991. The shear Prieta earthquake is shown by L.P. The size of this step was y,=+1840 

strain y, is a measure of shear strain approximately parallel to the San ne and y,=-3790 ne. The Chittenden aftershock sequence strain step is 

Andreas fault at this locality, while y' is a measure of shear strain in indicated by Ch. For any e (Ay,, Ay2), the maximum 

approximately a north-south or east-WeSt direction, and thus extension shear strain change is and 6fc=l/2tan-l[~y2/~yl] is 

normal to the fault. Three monthly intervals are indicated on the figure. the angk of for the axis of maximum extension. 

A SHEAR STRAIN ANOMALY FOLLOWING THE LOMA PRIETA EARTHQUAKE 

Table 2.-Surface creep events recorded at XSJ creepmeter at San Juan Bautista in 1988- 

1992 period, and corresponding strain offsets observed at SJT 

[Each strain event preceded the surface creep event by less than 1 hour] 

No. Date Strain event Creep event 

7i 72 &U Size 

ne n& ne mm 

1 21.11.88 -100 -40 -80 1.7 

2 22.12.89 -77 -38 -60 0.5 

3 27.01.90 -69 -4 1 -53 0.6 

4 07.04.90 -90 -55 -75 1.9 

5 09.06.90 -100 -50 -90 3.9 

6 04.07.90 None 0.3 

7 30.08.90 -120 -52 -9 1 3.2 

8 20.1 1.90 -50 -45 -70 2.1 

9 08.03.91 -95 -50 -80 3.6 

10 06.08.91 -120 -50 -90 5.0 

11 27.12.91 -90 -50 -80 0.3 

12 07.06.92 -1 10 -45 -90 5.3 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Figure 7.-Surface-creep data (provided by K. Breckenridge, USGS) from creepmeters along the fault south of San Juan Bautista, with positions of 

creepmeters XSJ, XHR, CWC, XFL and XMR, tensor strainmeter SJT, and Chittenden aftershock Ch indicated in the inset map. The trend lines 

indicated are from Burford (1988). Note that for the XMR data, the scale at the right axis should be used. 
THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17,1989: 

EARTHQUAKEOCCURRENCE 

AFTERSHOCKS AND POSTSEISMIC EFFECTS 

A MAGNETOTELLURIC SURVEY OF THE LOMA PRIETA EARTHQUAKE 

AREA: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND LOWER 

CRUSTAL CONDUCTIVITY 

By Randall L. Mackie and Theodore R. Madden, Massachusetts Institute of Technology; and 

Edward A. Nichols, University of California, Berkeley 

CONTENTS 

Page 

D289 

289 

290 

29 1 

292 

293 

303 

305 

307 

309 

309 

ABSTRACT 

We conducted a magnetotelluric survey around the Loma 

Prieta earthquake zone one year after the earthquake to 

look for possible electrical anomalies in the lower crustal 

part of the earthquake zone. We collected data at stations 

on each side of the San Andreas fault from Hollister to 

near Palo Alto. We used a two-dimensional transverse 

magnetic (TM) mode inversion algorithm to interpret the 

observed TM mode data since it is the TM mode that is 

sensitive to lower crustal resistivity variations near oceancontinent boundaries, whereas the transverse electric (TE) 

mode is not. Inversions of the data suggested a vertical 

zone of increased lower crustal conductance in the vicinity of the San Andreas fault near the region that experienced the main shock and aftershocks. The inferred zone 

of increased lower crustal conductance may be a permanent feature of this part of the San Andreas fault zone, or 

it may be a temporal feature related to the Loma Prieta 

earthquake. Because healing and ductile deformation processes in the lower crust should decrease its conductance 

with time, we speculate that this anomalous zone may be 

related to the Loma Prieta earthquake. With only data 

collected after the earthquake, we cannot confirm this 

speculation. However, because of the short time period 

between the Lorna Prieta earthquake and our measurements, we think that fluids were already present there 

before the earthquake, but that the rupturing process of 

the earthquake increased the lower-crustal fluid connectivity. Regardless, the conductance values determined by 

our inversions correspond to lower-crustal connected fluid 

porosities of about 0.06 percent for a 5-km-wide leakage 

zone. Higher fluid porosities would be needed if the 

fluids were concentrated in narrower zones, but we cannot determine the exact geometry of these zones. 

INTRODUCTION 

More than 5,000 aftershocks occurred over a distance 

of 60 km along the San Andreas fault zone and up to 20 

km in depth following the Loma Prieta earthquake 

(Oppenheimer, 1990). The main shock and aftershocks 

generally fell along a fault plane that dips 70' to the west 

and strikes approximately 130' from north. The earthquake was unusual because the epicenter was at a depth 

of almost 18 km (Dietz and Ellsworth, 1990), which is 

deeper than expected for this region of the San Andreas 

fault (Furlong and Langston, 1990), and because it is debatable whether or not the earthquake actually occurred 

on the San Andreas fault (Segall and Lisowski, 1990; Dietz 

and Ellsworth, 1990). 

Large-scale tectonic processes, such as earthquakes, may 

have a dramatic influence on the electrical conductivity of 

the crust. This is because these processes can affect the 

fluid regime in the crust (Torgensen, 1990), where the 

primary control of electrical conductivity is connected saline fluids. The crust, in general, becomes more resistive 

with depth, although some regions of the world appear to 

have conductive zones in the middle or lower crust (for 

example, Jones, 1987). At shallow depths, the increased 

resistivity is due to the narrowing of cracks (Brace and 

others, 1965). Deeper in the crust, temperatures are high 
D290 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

enough so that rocks deform ductilely rather than brittlely 

(Heard, 1976), and this should eliminate much of the connected porosity (Brace and Kohlstedt, 1980). It is more 

likely, however, that the middle-to-lower crust actually 

deforms semi-brittlely (Carter and Tsenn, 1987; Evans 

and others, 1990), which may be partly responsible for 

the fact that finite resistivities are measured for the lower 

crust. 

Because the main shock of the Loma Prieta earthquake 

was at a depth of 18 km, brittle deformation occurred 

down to at least that depth in the Loma Prieta region, and 

possibly deeper. Furthermore, it is possible that semibrittle deformation, where the main deformation mechanisms are crystal plasticity and microcracking (Evans 

and others, 1990), would exist to even greater depths. 

Such deformation might have caused an increase in the 

fracture and microcrack connected porosity in the lower 

crustal part of the fault zone and hence an increase in 

conductivity in that zone. (An alternate scenario is that 

fluids were brought into this zone after the earthquake, 

but we believe this is less likely because of the time scales 

required for fluids to percolate through a fairly impermeable crust.) 

Fluids at high pore pressures in the lower crustal part 

of a fault zone have been proposed to play an important 

role in the earthquake cycle (Rice, 1992). The existence 

of elevated pore pressures in a very narrow permeable 

slip zone having the usual laboratory frictional values 

may explain why the San Andreas fault does not show a 

pronounced heat flow peak and why the principal stress 

directions are at large angles relative to the trace of 

the San Andreas (Rice, 1992). Fluids at high pore 

pressures in the fault slip zone would give the lower 

crust a special role in the earthquake cycle as the pore 

pressure there is probably already close to the overburden 

pressure. 

Magnetotelluric (MT) measurements, under certain geographic circumstances, can provide information about vertical zones in the lower crust that have anomalous electrical 

properties. This occurs when currents are forced through 

these zones, which is the case near ocean-continental 

boundaries because of the mismatch of the current systems as a function of depth. Because the San Andreas 

fault system is parallel and close to the coast, we can take 

advantage of the currrent system set up by the oceancontinent adjustments to study the electrical properties of 

the lower crustal part of the fault zone. Therefore, we 

collected MT data in the Loma Prieta earthquake region 

to investigate the lower crustal properties of that area. 

Unfortunately, since all of our MT data were collected 

after the earthquake, we will not be able to make any 

statements concerning time changes in conductivity across 

the fault zone that might have been due to the earthquake 

itself. 

THE MAGNETOTELLURIC TM-MODE 

COAST EFFECT 

Electromagnetic fields are affected by an ocean-continent boundary in two very distinct ways. One type of 

coast effect is due to the concentration of near-surface 

electrical currents running parallel to the coast in both the 

ocean and the oceanic upper mantle, which causes anomalies in the vertical magnetic field recorded near an ocean 

coast. In the work herein, we are concerned with another 

coast effect-one that perturbs the electric fields running 

perpendicular to the coast. 

Because the ocean has such a high conductance, most 

of the induced electrical current in oceanic regimes is 

concentrated in the ocean itself up to periods of 500 s, 

whereas in a continental regime the current is mostly in 

the mantle. Because electrical currents are divergencefree, there is continuity in the normal component of current at the ocean-continent boundary. This results in extra 

current perpendicular to the coast (the transverse magnetic or TM mode) in the continental upper crust compared to what would be induced were the ocean absent. 

Consequently, the currents must readjust in some broad 

zone around the ocean-continent boundary. This readjustment takes place by leaking currents out of the upper 

crust, across the resistive lower crust, and into the mantle 

(fig. 1). In this readjustment zone, the current system is 

still dominated by the oceanic current system. 

The width of this readjustment zone depends on the 

conductance of the upper crust and the resistance of the 

lower crust (Ranganayaki and Madden, 1980). The distance at which the current has been reduced by a factor of 

e by leaking into the mantle is termed the adjustment 

distance, and is equal to the square root of the product of 

the integrated upper crustal conductance and the integrated 

lower crustal resistance (Ranganayaki, 1978). The adjustment distance can be 100 km or more for continental 

regimes, and even at these distances the upper crustal 

current system is still contaminated by excess ocean 

currents. It actually takes several adjustment distances 

for the current system to return to normal continental 

values. 

Vertical zones in the lower crust that have an anomalously high conductance can dramatically alter the TM 

response across that zone (Park and others, 1991; Mackie, 

1991). This is because more current is attracted out of the 

ocean by this leakage zone, but less current remains in the 

upper crust inland of this zone. This results in a dramatic 

shift in the TM-mode response across the zone. The transverse electric (TE) mode response would not be sensitive 

to this zone. This is because the component of the electric 

field parallel to conductivity boundaries is continuous 

across those boundaries, and therefore, the TE-mode responses are not anomalously perturbed from their normal 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D29 1 

values. By carefully following the readjustment of the 

TM current system perpendicular to the ocean-continent 

boundary, one can map out anomalous vertical features in 

the lower crustal resistivity structure. (However, we would 

not be sensitive to anomalously conductive horizontal layers in the lower crust since we can only determine the 

integrated vertical resistivity properties.) 

Because most of the geologic features in California are 

oriented parallel to the ocean-continent boundary, we are 

in a good position to look for anomalous zones in the 

lower crust in the region associated with the Loma Prieta 

earthquake. By mapping the TM-mode response across 

the fault at several different locations, we can look for 

anomalous vertical conductivity channels in the lower 

crustal part of the fault zone. 

GEOLOGIC FRAMEWORK 

California has been subjected to many tectonic events 

throughout its history, including arc-related compressional 

orogenies, arc volcanism and plutonism, and Cenozoic 

transform motion (Burchfiel and Davis, 1972, 1975; 

Dickinson, 1981). The geology in the area of the Coast 

Ranges of central California, as shown in figure 2, is characterized by three main lithologic sequences (Bailey and 

others, 1970; Page, 198 1): (1) the Salinian formation (plutonic and metamorphic rocks from a magmatic arc environment), (2) the Franciscan formation (a subduction zone 

complex), and (3) the Great Valley sequence (forearc basin sediments). Interspersed among the outcrops of these 

formations are surface deposits that are predominately 

Cenozoic and Mesozoic marine and non-marine sediments 

and alluvium. 

The Salinian block, bounded on the northeast by the 

San Andreas fault, is composed of metasedimentary rocks 

that have been invaded by granitic plutons. It is believed 

to have been translated some 600 km northward from its 

original position as part of the Klamath-Sierra Nevada 

terrane. The resistivities of the granites of the Salinian 

formation are much more resistive than the sedimentary 

and metamorphic rocks to the east, and are typically around 

500 Q-m (Mazzella and Morrison, 1974). 

The Franciscan formation is an assemblage of both oceanic and terriginous materials. It consists of graywacke, 

shale, mafic volcanic rock, chert, limestone, and metamorphic rocks of zeolite and blueschist facies. These components commonly occur in a pervasively sheared 

argillaceous matrix. Much of the Franciscan formation 

has undergone blueschist metamorphism at temperatures 

of around 200Â° and pressures of around 6-8 kbar (Page, 

1981). It is believed that the Franciscan formation originated in a subduction zone along the western margin of 

the North American plate in the late Mesozoic era and 

that these rocks underwent metamorphism as the oceanic 

plate was subducted under the continental plate. Surface 

conductive mantle 

Figure 1.-Simple sketch showing the magnetotelluric TM-mode coast ally leaks off into the mantle as one moves away from the coast. This 

effect. The ocean induces extra current that gets trapped in the continen- leakage effect is what makes one sensitive to the electrical properties of 

tal upper crust. This extra current in the continental upper crust gradu- the lower crust. 
D292 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

electrical measurements show typical resistivity values of 

5-50 Q-m for the sedimentary and metamorphic rocks of 

the Franciscan formation (Mazzella and Morrison, 1974). 

Finally, the Great Valley sequence is composed of stratified terriginous clastic sediments, primarily gray-wackes 

and shales, derived from the Klamath-Sierra Nevada terrane. These sediments are believed to rest on an ophiolite 

basement, beneath which the Franciscan formation has 

been thrust, and are even more conductive than the Franciscan rocks, typically being about 1-10 Q-m. 

THE MAGNETOTELLURIC DATA 

A magnetotelluric survey was conducted in conjunction 

with Electromagnetic Instruments, Inc. (EMI) from October 20-30, 1990, one year after the Loma Prieta earthquake. EM1 collected a total of ten sites in pairs of two 

simultaneous recordings, one on each side of the fault 

(except for site 10, which was on the fault because of 

logistical problems). Fig. 2 shows the locations of these 

stations. EM1 collected Ex, E, Hx, H, and Hz data at 

each pair of stations using synchronized clock control. 

They recorded data in three bands, with sample frequencies of 500 Hz, 20 Hz, and 2 Hz, thus yielding coverage 

over a period range from approximately 0.005-1000 s. 

The data were processed to yield a 10 by 10 cross-power 

matrix per frequency (approximately 6 frequencies per 

decade) for each pair of stations. From this cross-power 

matrix, MT parameters could be computed using the 

horizontal magnetic fields at the second site as a remote 

reference. 

Figure 2.Ã‘Simplifie geologic map (after Bailey, 1966) showing the locations of the Franciscan Formation, the Salinian Block, and major faults. In 

between the Franciscan and Salinian outcrops are surface deposits of marine and non-marine sedimentary rocks and alluvial deposits. Also shown 

are the location of the MT stations and the profiles used for the inversion analysis. 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D293 

Remote referencing was fairly successful for canceling 

out the effects of random noise disturbances in the data at 

most sites, but there were additional noise problems that 

the remote referencing was not able to deal with. Robust 

processing methods also were not able to deal with those 

noise problems (Egbert, oral commun., 1992). In general, 

the data were noisier the closer they were to the industrialized San Francisco Bay area to the north. Sites 5 and 6, 

the furthest north, were the noisiest data of all and were 

basically unusable. Site 8 also was noisy, which we later 

discovered was because that site was located close to an 

electrical power sub-station. 

At many of the sites (1,2,3,4,5,6) the data in the band 

from 1 to 100 s seemed to be contaminated by some manmade noise source. There are two lines of evidence for 

this. First, the magnetotelluric responses at these sites in 

this frequency band are characterized by slopes close to 1 

in log of apparent resistivity versus log of period, and a 

sudden decrease of the phase to near 0Â° which is characteristic of local noise sources (Qian and Pedersen, 1991). 

This is clearly evident in the MT data from site 1, Calero, 

shown in figure 3. Second, and probably more important, 

there is a large anomaly in the tipper magnitude from 1 to 

100 s at these same sites. (Figure 4 shows an example of 

the anomaly from site 1, Calero.) The tipper is a measure 

of the ratio of the vertical magnetic field to the horizontal 

magnetic field. Local industrial noise sources, which act 

like vertical magnetic dipoles and horizontal electric dipoles, can create large tipper magnitudes which are clearly 

non-MT. These sources generate coherent electromagnetic 

energy, but their impedances are different than those of 

natural MT signals at periods longer than 1 s. At periods 

greater than 100 s, however, the natural MT signals were 

stonger than the man-made noise. It is possible that this 

noise problem is due to the BART transportation system 

in the San Francisco Bay area, which acts like a large 

vertical dipole (Fraser-Smith and Coates, 1978; Ho and 

others, 1979), although any local industrial source could 

cause similar features in the data. 

At the same time that EM1 was collecting data, we 

(Madden and Mackie) collected single-site, longer-period 

(10-1000 s) MT data at several additional stations, mostly 

along the San Andreas fault zone. We attempted to gather 

data at the same time EM1 was collecting long-period 

data so we could use the EM1 data for remote referencing 

(we synchronized our clocks with EMI's clocks every 

couple of days). However, because of logistical problems 

in setting up the stations, we were only successful in coordinating this at one site (site 11, shown on the map in 

fig. 2). At the other stations, the daytime man-made noise 

problems prevented us from obtaining valid impedance 

estimates. The data from the EM1 group were generally 

less contaminated with the man-made noise because they 

typically collected their long-period data (greater than 1 

s) during the night, when such noise is at a minimum, 

whereas we collected all our data during the day when 

these noise problems were highest. Unfortunately, the natural MT signals are also usually stronger during the day 

than at night, which reduces the advantage of night-time 

recording. 

TWO-DIMENSIONAL INVERSION OF 

THE DATA 

An eigen-state analysis (LaTorraca and others, 1986) 

was used to decompose the impedance tensors for each 

site. At all stations except site 3, the major electric eigenvector direction, which corresponds to the maximum amplitude eigen value of the impedance tensor, was 

approximately perpendicular to the coastline, and this is 

the mode we identify as the transverse magnetic (TM) 

mode. The TM mode near ocean-continent boundaries has 

lower phases at longer periods as compared with the 

TE mode. This is due to ocean electrical current perpendicular to the coastline being trapped in the continental 

upper crust (Mackie, 1991). At site 3, the maximum amplitude eigen value corresponded to the TE mode, and the 

TM mode corresponded to the minimum eigen value. We 

suspect this is due to local inhomogeneities that caused a 

static shift in the electric field perpendicular to the ocean 

and probably gave us a mixed MT impedance. 

Data from sites 5, 6, and 8 were unusable because they 

were too contaminated by man-made noise. Data from the 

remaining sites were mostly usable, although we edited 

out the data at some of the sites (1,2,3,4) from 1 to 100 s 

that appeared contaminated by man-made noise sources. 

Finally, selective editing of outliers and obviously inconsistent data was done for all stations prior to running twodimensional inversions of the data. The remaining data 

were input to the two-dimensional TM-mode inversion 

algorithm described in Mackie and others (1988). We used 

only TM-mode inversions of the data for two reasons: (1) 

the lower crustal resistivity anomalies discussed earlier 

affect only the TM mode and (2) two-dimensional TMmode interpretation is more accurate for profiles centrally 

located across elongate three-dimensional bodies 

(Wannamaker and others, 1984). We do, however, show 

the TE-mode data and model predictions even though the 

TE data were not used in the inversions. 

The inversion procedure we used to interpret the data is 

the one we term the maximum likelihood inverse (Mackie 

and others, 1988; Madden, 1990). The maximum likelihood inverse gives the solution that maximizes the joint 

probability of fitting the observed data, subject to data 

covariance constraints, and adhering to an a priori model, 

subject to model covariance constraints (Tarantola, 1987). 

Our formulation of the inverse problem holds the block 

geometry of the system fixed-only resistivities are al- 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

EM1 site 1, Calero 

A A AAAA A 

AAAAA A i 

i t 

- 10-3 I 0-2 10-1 10% 101 1 02 103 I o4 

period (seconds) 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D295 

lowed to vary during the inversion. Additionally, in all 

Figure 3,-Eigen-analysis results of the MT data at site 1, Calero. our inversions, we gave equal weight to fitting the phases Shown are the maximum and minimum apparent resistivities, associated 

EM phases, and the electric field and magnetic field eigen-directions. 

and In (amplitudes). 

Note the steep slope in the apparent resistivities and phases near 0" in The profiles which we did inversions are shown 

the period range from 1 to 100 s. We believe these are not natural MT in figure 2. Each profile used data from only two or three 

responses but rather are due to Bay Area Rapid Transit noise contami- stations. For each profile, we ran many inversions of the 

nating the signals. This behavior was at most sites in our survey area. data using different constraints and a priori models. The a priori models were all generally constrained 

using available geologic and geophysical data. First, we 

EM1 site 1, Calero, Tipper results 

Figure 4.-Plot of the tipper amplitude at 

site 1, Calero. Shown are the amplitudes 

in the direction where Hz is most correlated with H. Note the large anomaly in 

the period range from 1 to 100 s. We believe this is evidence of Bay Area Rapid " 

1 0-3 1 0-2 10-I 1 O0 10' 1 02 1 O3 104 Transit interference in the natural signals. 

Similar anomalies were observed at most 

period (seconds) sites in our survey area. 
D296 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

used long-period (greater than 1 hour) MT data from the 

Hollister area (Bennett, 1985; Mackie and others, 1988), 

shorter-period (10-10,000 s) MT data from central California (Park and others, 1991), and the results of a controlled-source EM measurement on the Pacific Ocean floor 

(Cox and others, 1986) to constrain the lower crust and 

mantle a priori resistivity values. Second, we used the 

seismic tomographic inversions and MT analysis of 

Eberhart-Phillips and others (1990) to help constrain the a 

priori upper crustal resistivity variations. Third, we used 

bathymetric contour maps for approximating the ocean 

floor topography. Finally, we used surface geologic information (Clark and Rietman, 1973) for setting the very 

near surface resistivity information. We show one a priori 

model for profile A in figure 5. Other a priori models 

were tested that included putting vertical conductive zones 

in the lower crust and making the entire lower crust much 

more conductive. The a priori models for the other profiles are essentially the same except for minor upper crustal 

differences. The inversion procedure weights fitting the 

data much more strongly than adhering to the a priori 

model (by a factor of lo5 in the ratio of the data to model 

covariances), although it never completely looses sight of 

the a priori model. The inversion procedure itself is not 

biased, although there certainly are biases introduced by 

our choice of an a priori model and by the a priori model 

constraints. 

Profile A uses data from sites 1 and 2 and goes over the 

epicenter of the main earthquake shock. We feel that these 

data sites represent our best data, so we will focus on the 

results for this profile. In the following discussion, it should 

be understood that a lightly weighted smoothing operator 

and a small diagonal term were always included in the 

model covariance operator used in the inversion, and that 

an "unconstrained inversion" means no additional constraints were used beyond those, whereas a "constrained 

inversion" means additional constraints were used. A grayscale plot of the results of an unconstrained two-dimensional TM-mode inversion of the data from these sites is 

shown in figure 6. The horizontal scale of the plot is the 

square root of distance from the ocean-continent boundary. The vertical scale is the square root of depth, but we 

PROFILE A (aptos-calero), a priori model 

100 30 0 10 I 30 100 

resistivity in ohm-m 

Figure 5.-Grayscale plot of the a priori model used for inversions sections-one for 0-45 km and one for 45-600 km. On this and all other 

along profile A (fig. 2). Resistivities are given in a-m. The horizontal grayscale plots, the station locations are marked with asterisks and the 

scale goes as the square root of distance from the coastline. The vertical San Andreas fault is marked with an arrow. See the text for a discussion 

scale goes as the square root of distance and is broken up into two of how the a priori values were set. 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D297 

have broken it up into two different sections, the first one 

for 0-45 krn depth and the second for 45-600 km depth. 

The models actually extend much further out on either 

side, but we show only the central portions of each one as 

there were only minor variations beyond the section we 

show. The station locations are marked with an asterisk. 

The inversion resulted in a wide area of increased conductance in the lower crust around the fault zone, with the 

largest change (10 times more conductance) concentrated 

in the immediate vicinity of the fault zone. Additionally, 

the lower crust was made more resistive away from the 

fault zone both under the ocean and under the continent. 

These changes were made by the inversion routine in order to, first, attract enough current onto the continental 

upper crust, and second, to lose enough of that current 

between the two data sites to fit the data at those sites. 

We show the predictions from this model in figure 7 as 

solid lines. (We show the predictions for all periods between 0.01 s and 2000 s even though the data from 1 to 

100 s were not used in the inversion.) The TE-mode predictions are also shown, although we did not use them in 

the inversion analysis. (At some of the data sites there 

was a static shift between the observed TM- and TEmode apparent resistivities. Therefore, in plotting out the 

predicted TE-mode data, we applied, where necessary, a 

static shift correction to the predicted TE data so that they 

agreed with the observed TE data at the higher frequencies.) As seen on this and following plots, the TE-mode 

predictions are actually quite good at most sites except 

site 7, Hollister, where they are the worst. This gives us 

increased confidence as to the validity of our two-dimensional interpretation. 

We ran a second inversion of the data from profile A 

starting with the same a priori model except that we constrained the lower crust and upper mantle to remain at 

their a priori values. We show the resulting model in 

figure 8. For this inversion, we required all resistivity 

changes to be put into the upper crust. There are mostly 

minor changes in the upper crustal resistivity values as 

opposed to the results from the previous inversion, except 

just above the lower crust near the fault zone and under 

the Sierra Nevada. In both those places, the conductance 

PROFILE A (aptos - calero), unconstrained inversion 

100 30 0 10 1 30 100 

resistivity in ohm-m 

Figure 6.Ã‘Grayscal plot of the result of a smoothed but otherwise unconstrained two-dimensional TM-mode inversion of the data along profile A 

(fig. 2). Note the broad area of leakage put into the lower crust in the vicinity of the San Andreas fault zone. The data predicted by this model are 

shown in figures 7 and 9. 
Site 2, Aptos Site 1, Calero 

TM mode 

TE mode 

period (seconds) 

Figure 7.Ã‘Observe and predicted data at (left) site 2, Aptos (west of the San Andreas fault), and 

(right) site 1, Calero (east of the San Andreas fault). The solid lines are the predicted data from the 

model in figure 6, and the dashed lines are the predicted data from the model in figure 8. We show 

TM mode 

period (seconds) 

both the predicted TM- and TE-mode data, even though only the TM-mode data were used in the 

analysis. The predicted TE-mode data were static shifted so that they agreed with the observed TEmode data at the highest frequencies. 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D299 

was increased as compared to the previous inversion. This 

was most likely necessary to attract the additional current 

onto the continental upper crust and to pull it down from 

the surface to the bottom of the upper crust across the 

fault zone. However, the increased conductance under the 

Sierra Nevada is probably unrealistic (Park and others, 

1991). We show the TM- and TE-mode predictions from 

this model in figure 7 as dashed lines. 

Looking at figure 7, it seems as if there are not many 

differences in the predictions between the two models. 

This is because it is difficult to see subtle changes in the 

slopes of the apparent resistivities, but the changes in the 

phase are more obvious since they show integrated effects 

of changes in the apparent resistivity curves. In figure 9, 

we plot the observed and predicted TM phases for the two 

models just discussed. In this figure, the circles correspond to the TM phase data from site 2, Aptos (nearest 

the ocean), and the triangles correspond to the TM phase 

data from site 1, Calero. The predicted data from each site 

are shown with solid and dashed lines respectively. The 

TM-mode data show an increase of about 5-10' in the 

magnitude of the phase from site 2, Aptos, to site 1, Calero, 

and with our choice of a priori model, this is fit only 

when the inversion puts a conductive window in the lower 

crust. 

Unfortunately, the interpretation of these data is nonunique. With the resistive lower crust we put into the a 

priori models, the observed data cannot be fit without 

changes in the lower crustal resistivity between the two 

stations. However, there are other models that can fit the 

observed data without any lower crustal leakage paths 

and using only upper crustal resistivity variations 

(Eberhart-Phillips and others, 1990; Stanley, oral commun., 

1992). These models, however, must have lower crusts 

with much lower resistivities than we have put into our 

models (1,000 0-m vs. 30,000 Q-m). Although these models will fit the observed data at the shorter periods, they 

fail to predict the longer period data that we have previously analyzed at Hollister (Mackie and others, 1988). 

We have, however, used the results of that study to help 

constrain our interpretation in this study. This is because 

at the longer periods involved in the Mackie and others 

(1988) study, the electric fields perpendicular to the 

coastline at longer periods must be similar up and down 

along the coast since the curl of the electric field must be 

small at these periods and any significant changes in the 

PROFILE A (aptos - calero), ID lower crust and upper mantle 

300 km 100 30 0 10 I 30 100 

0.3 1.0 20.0 50.0 200.0 500.0 2000,O 5000.0 20000.0 50000.0 200000.0 

resistivity in ohm-m 

Figure 8.-Grayscale plot of the result of an inversion that kept the lower crust and upper mantle at its a priori values, thus forcing all changes to be 

put into the upper crust. The data predicted by this model are shown in figures 7 and 9. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Model 1, conductive window in lower crust 

10-3 1 0-2 10-1 1 00 1 o1 1 02 103 1 o4 

period (seconds) 

Model 2, uniform lower crust and mantle 

aptos TM phase 

/- 

/- 

/ 

/ 

""calero TM phase 

10-1 1 00 101 1 02 

period (seconds) 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D301 

electric field would involve huge vertical magnetic fields. 

Therefore, any model we produce to fit the shorter period 

data of this study should also do a good job at predicting 

the longer period data of the Mackie and others (1988) 

study. The longer period data are especially sensitive to 

the lower crustal and mantle resistivity properties, and 

therefore, even though we did not explicitly include them 

in the inversion scheme, they played an important role in 

constraining the range of possible models that fit the data. 

The unconstrained inversion increased the conductance 

in the lower crust over a broad zone, but the data are not 

sensitive to the details of the geometry of the lower crustal 

leakage path between the two stations. Physically, it is 

probably more realistic that any anomalies in the lower 

crustal resistivity around the Loma Prieta earthquake area 

Figure 9.-Plot of the observed and predicted TM-mode phases for 

the model of figure 6 and the model of figure 8. Note that the data are fit 

only when the inversion is allowed to put some leakage into the lower 

crust. If the lower crust and upper mantle are uniform, then the observed 

data cannot be fit. 

would be concentrated in a narrower zone. The broad zone 

of lower crustal leakage shown in figure 6 has a horizontal dimension of approximately 50 km. If that lower crustal 

leakage zone was made 10 times narrower (5 km), then it 

should require 10 times more conductance to get the same 

amount of current leakage as the broader zone. For the 

model in figure 6, this corresponds to a lower crustal 

resistivity value somewhere between 100 and 1,000 Q-m 

for a zone 5 km wide versus 1,000-10,000 Q-m for a 50- 

km-wide zone. 

Therefore, we assigned a narrow conductive leakage 

path in the lower crust for the a priori model and ran an 

inversion. This inversion reduced the resistivity in the 

lower crustal fault zone from its a priori value of 1,000 

Q-m to approximately 300 Q-m, and slightly increased 

the resistivity under the western edge of the Sierra Nevada (fig. 10). This was necessary to encourage leakage 

through the narrow zone. The predicted and observed 

TM-mode phases are shown in the top panel of figure 11. 

This model resulted in a little more separation in the phases 

than the previous unconstrained model in figure 6 because the narrow zone loses more current across a short 

PROFILE A (aptos - calero), apriori narrow conductive fault zone 

300 km 100 30 0 10 I 30 100 

0.3 1.0 20.0 50,O 200.0 500.0 2000.0 5000.0 20000.0 50000.0 200000.0 

resistivity in ohm-111 

Figure 10.-Grayscale plot of the result of an unconstrained inversion of the data along profile A where the a priori model had a narrow conductive 

(1,000 Q-m) fault zone in the approximate position of the San Andreas. The inversion made the fault zone even more conductive, forcing it down to 

about 300 Q-m. The data predicted by this model are shown in the top panel of figure 11. 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

apriori narrow conductive fault zone 

Or 

aptos TM phase 

calero TM phase 

period (seconds) 

apriori Great Valley leakage 

oraptos TM phase *--^Â¥& 

a'* 0- 

~.O"~alero TM phase 

period (seconds) 

apriori narrow conductive fault zone and GV leakage 

"'"I- aptos TM phase 

Figure 11 .-Plots of the observed and 

predicted TM-mode phases for three 

different inversion results. In the top 

panel are the predictions for the model 

in figure 10. The middle panel shows 

the results of an inversion with a priori 

Great Valley leakage, and the bottom 

panel shows the results of an inversion with a priori Great Valley leakage and a narrow conductive fault 

zone. 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY 

distance as opposed to the broad zone. (It should be noted 

that due to the smoothing constraints of our inversion 

scheme, it is difficult for the inversion to put in abrupt 

conductivity boundaries unless they were assigned a priori, 

which is why the unconstrained inversion put in a broad 

zone of increased conductance.) 

Although we did not have data further to the east in the 

Great Valley and Sierra Nevada, we felt it was important 

to see how the inferred leakage zones of Park and others 

(1991) on both sides of the southern Great Valley might 

affect the interpretation of our data. Park and others (1991) 

collected data within 100 km of our test area, but he observed longitudinal uniformity of the Great Valley resistivity structure over a distance of approximately 150 km, 

so it is likely that the inferred Great Valley leakage paths 

of Park and others (1991) would exist further to the north 

near our study area. Therefore, we ran an inversion where 

we a priori assigned the Great Valley leakage zones using 

the results of Park and others (1991) as a guide in setting 

the resistivities and widths of those zones. The Great Valley leakage has the effect of attracting more ocean current 

onto the continental upper crust. Consequently, the inversion did not need to make the oceanic lower crust as 

resistive as did the unconstrained inversion without Great 

Valley leakage (fig. 6). As before, however, it put in a 

broad area of increased conductance around the Loma 

Prieta earthquake zone, with electrical properties not much 

different than before. The observed and predicted TMmode phases are shown in the middle panel of figure 11. 

This model does not fit the observed data as well as the 

previous model with an a priori narrow fault leakage zone. 

If, however, we run an inversion where we a priori assign 

the Great Valley leakage and a narrow conductive fault 

zone, we find that the resulting model predicts the observed data very well, and this is shown in the bottom 

panel of figure 11. The lower crustal resistivities of the 

narrow fault zone were again reduced to around 300 0,-m 

from the a priori values of 1,000 0,-m, but the inversion 

also put in increased conductivities around the narrow 

fault zone to account for the extra current from the Great 

Valley leakage. 

SUMMARY OF MODELING RESULTS 

Although the interpretation of our data is non-unique, 

we believe it is likely that there is a vertical lower crustal 

conductive zone in the vicinity of the San Andreas fault. 

In a smoothed but otherwise unconstrained inversion, a 

broad area of increased conductance in the lower crust 

was put in by the inversion. Another inversion of the data 

that fixed the apriori lower crust and upper mantle (all 

changes forced into the upper crust) was not able to fit the 

data as well and required some unrealistic resistivities 

under the Sierra Nevada. It is possible, however, to fit the 

data without any lower crustal leakage paths (EberhartPhillips and others, 1990; Stanley, oral commun., 1992), 

but these models require much more conductive lower 

crusts than we put into our a priori models, and they do 

not give model responses that agree with the longer period data we have analyzed at Hollister (Mackie and others, 1988). The inferred Great Valley leakage of Park and 

others (1991) had only a minor effect on the results of the 

inversions, and that was mainly in the resistance of the 

lower crust away from the fault zone. Although we cannot determine the exact geometry of the lower crustal 

leakage, we suspect a narrow conductive zone is geologically more realistic. 

For the remaining profiles, we tested the various hypotheses for lower crustal leakage just described for profile A. The results for profiles B and C were similar to 

those obtained for profile A in that unconstrained inversions resulted in broad zones of increased conductance in 

the lower crust around the San Andreas fault zone. As for 

profile A, these data could also be fit with models that 

resulted from inversions with a priori narrow lower crustal 

fault zones. For the sake of brevity, however, we will 

only show the model predictions from the results of two 

different inversions of the data for each of the remaining 

profiles. For the first inversion, we started with an a priori 

model that had a narrow conductive fault zone and Great 

Valley leakage. The upper mantle and lower crusts in these 

models were similar to those that resulted from inversions 

of the data along profile A, and for this inversion we 

forced all changes to be made only in the upper crust and 

in the lower crust in the vicinity of the San Andreas fault. 

This is a reasonable approach since all the profiles should 

see approximately the same lower crust and upper 

mantle away from the fault zone. For the second inversion, we took out the narrow fault zone and Great Valley 

leakage and forced all changes to be put only into the 

upper crust. 

The results of the inversions along profile B (figs. 12, 

13) again suggest the existence of a conductive zone in 

the lower crust in the vicinity of the San Andreas fault 

zone similar to that shown in figure 10 for profile A. The 

statements made earlier about non-uniqueness and alternative models apply along this profile as well. There is a 

change of almost 20' in the observed TM-mode phase 

between the two sites along this profile, although they are 

only separated by a distance of 17 km. (The amplitude of 

the phases at site 4, however, seem to be somewhat 

too low in comparison with the slope of the apparent 

resistivity curve, so the true change in phase across the 

zone may be a little less than 20Â°. The only way the 

inversion can even come close to fitting the data is by 

making the narrow fault zone very conductive (approximately 50 0,-m over the 5 km width as compared to 300 Q.- 

m for profile A). The results from the inversion that 

constrains the lower crust and upper mantle to remain 

fixed at their a priori values show that the current levels 
Site 4, Loch Lomond Site 3, Reynolds Kancn 

TE mode 

period (seconds) 

Figure 12.Ã‘Observe and predicted data at (left) site 4, Loch Lomond (west of the San Andreas 

fault), and (right) site 3, Reynolds Ranch (east of the San Andreas fault). The solid lines are the 

predicted data from an inversion with an a priori narrow fault zone, and the dashed lines are the 

predicted data from an inversion where no fault zone was put into the a priori model and all 

TE mode 

period (seconds) 

changes were forced into the upper crust. We show both the predicted TM- and TE-mode data, 

even though only the TM-mode data were used in the analysis. The predicted TE-mode data were 

static shifted so that they agreed with the observed TE-mode data at the highest frequencies. 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D305 

and changes in those levels across the fault zone are incorrectly predicted. 

On the other hand, the data from the sites along profile 

C show very little systematic separation in phases across 

the fault zone (figs. 14, 15). The data at the stations along 

this profile did not have anomalous tippers in the band 

from 1 to 100 s, but the apparent resistivities and phases 

were not quite consistent with each other either (the data 

from 0.1 to 1 s were not used because of noise). For 

inversions where the data from 1 to 100 s were not used, 

as with profiles A and B, results were obtained that indicated that some lower crustal leakage was preferred by 

the inversions, but not nearly as much as indicated for 

profiles A and B. When we included the data from 1 to 

100 s in the inversion, then lower crustal leakage similar 

to that in profile A was preferred. The model predictions 

for the inversion without the conductive fault zone are not 

too different than those with the narrow fault zone, except 

that there is more separation between the curves when the 

fault zone is included. The results along this profile are 

too ambiguous to make any definitive statements about 

lower crustal leakage here, but the inversions clearly favored some lower crustal leakage when no constraints 

were placed on the inversions. 

Finally, we look even further south to profile D, located near Hollister. The town of Hollister is actually in 

the creeping section of the Calaveras fault. The transition 

from creeping to locked behavior along the San Andreas 

fault occurs near the town of San Juan Bautista just west 

of Hollister. Unfortunately, the data from EM1 site 8 was 

unusable, so we only had data from site 7 to use for the 

inversion. The data for this site is a combination of the 

original EM1 data and MIT data from our second trip out 

to the field. The data for periods greater than 10 seconds 

is the MIT data (this was more coherent and less noisy 

than the original EM1 data), and the higher frequency data 

is the EM1 data. With just one station, we cannot say 

anything definitive about the resistivity structure along 

this profile except that the general trend of the data is in 

keeping with the data further to the north. These data can 

be fit equally well with or without a conductive zone in 

the lower crust (fig. 16). 

Before we end this section, let us return briefly to the 

issue of the TE-mode data and model predictions. In all 

cases except site 7, Hollister, the predicted TE-mode data 

were actually fairly close to the observed data even though 

they were not used in the inversion analysis. It is well 

known that the TE-mode response can be dramatically 

altered because of finite strike lengths of bodies or other 

three-dimensional inhomogeneities, whereas the TM mode 

is much less affected (Wannamaker and others, 1984). 

Near the interior edge of a conductive three-dimensional 

body, the fields will be depressed because not enough 

current is induced in the body to bring the levels up to 

their two-dimensional values, nor is enough current gathered into the body to do likewise. This depresses the MT 

response there as compared to what would be observed 

for a two-dimensional model with infinite strike. This flattens out the TE response, which is what we observe at the 

Hollister site. The fact that Hollister sits in a conductive 

valley and thus may be seeing the effects of finite strike 

lengths may explain the disparity at this site between its 

observed and predicted TE-mode data. 

GEOPHYSICAL IMPLICATIONS FOR 

THE SAN ANDREAS FAULT SYSTEM 

In general, the electrical conductivity and distribution 

of fluids in the lower crust has always been somewhat of 

an enigma. The fact that finite resistivities are measured 

for the lower crust most likely implies that it has a small, 

connected, and fluid-filled porosity (Brace and others, 

1965; Lee and others, 1983; Shankland and Ander, 1983), 

although graphite has also been proposed as a possible 

conduction mechanism in the lower crust (Duba and 

Shankland, 1982; Frost and others, 1989). There are many 

lines of evidence suggesting that free water is present to 

at least moderate depths within the crust. These include 

electromagnetic field studies (Shankland and Ander, 1983), 

isotopic studies of batholithic rocks (for example, Taylor, 

1977), geochemical studies (for example, Kerrich and others, 1984; Kerrich, 1986), analyses of metamorphic rocks 

(for example, Etheridge and others, 1984), and seismic 

studies (for example, Nur and Simmons, 1969; Jones and 

Nur, 1982). At greater depths within the crust, the issue is 

far more controversial since granulite facies rocks and 

anhydrous mafic silicates (such as garnets and pyroxenes) 

commonly associated with the lower crust indicate that it 

should be fluid-free (Newton, 1990). It is argued (Yardley, 

1986) that any water introduced into a granulite facies 

lower crust would be consumed by rehydration reactions 

forming lower-grade minerals. However, it is also possible that a small amount of water may be present in the 

lower crust in equilibrium with a hydrated mineralogy 

that is not present at the Earth's surface (Gough, 1986). 

Even if a very small amount of free water does exist in 

the lower crust, it may not necessarily exist in a connected form; rather, because lower crustal rocks are believed to deform ductilely, it is conceivable that the water 

may exist only in isolated pockets or in absorbed hydrous 

phases unless the wetting angles are low enough (Watson 

and Brenan, 1987) or unless fluid pressures can be maintained at close to the rock pressure for geologically significant times (Walder and Nur, 1984). Another problem 

with maintaining water in a permeable crust is that over a 

period of time, the water should migrate upwards and out 

of the lower crust, unless it is trapped by some imperrneable layer, because of buoyancy forces (Bailey, 1990). 
AFTERSHOCKS AND POSTSEISMIC EFFECTS 

apriori narrow conductive fault zone 

-90 I I 11 111111 I 1 1 11 1111 I I I I 1111 I 1 1 1 11111 I I I llll~~ I I I I I Ill1 I I lllll 

I o-~ 10-2 10-1 I o" I o1 I o2 I o3 I o4 

period (seconds) 

no fault zone allowed 

period (seconds) 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D307 

This may not be an issue if the hydraulics of the Earth's 

crust are time-dependent (Nur and Walder, 1990), as is 

quite likely. 

Our inversions of the Loma Prieta magnetotelluric data 

suggest that there may be vertical zones of increased electrical conductance in the lower crust that broadly correlate with the San Andreas fault zone in the area that was 

involved with the Lorna Prieta earthquake. Our findings 

are contrary to those of Park and others (1991) and 

Eberhart-Phillips and others (1990) who found no evidence for an anomalously conductive San Andreas fault 

zone. However, in a study of the telluric field variations 

around the Palmdale section of the San Andreas fault 

(Madden and others, 1993), systematic changes in the telluric field relationships were attributed to small increases 

in the lower crustal conductance in a vertical zone associated with the San Andreas fault (60 milli-Seimens over a 

5-year period, which is too small to be seen by MT measurements). The strain necessary to cause this increase in 

lower crustal conductivity would be less than could be 

measured by the usual surface-strain measurements. 

We believe that connected fluids in the lower crustal 

part of the fault zone are responsible for its high conductance relative to the lower crust away from the fault zone. 

The question, then, is about what amount of fluid are we 

talking? Typically, porosity estimates for the lower crust 

are derived by using Archies Law with an exponent of 2 

(Brace and others, 1965) and the laboratory values for 

saline solution resistivities at high pressures (Quist and 

Marshall, 1968), which probably results in overestimates 

of the fluid porosities. This is because, first of all, an 

exponent of 1 in Archies Law is probably more appropriate for lower crustal conditions. Looking at the data in 

Brace and others (1965), we see that the resistivity of 

crystalline rocks at pressures from 0 to 10 kbar is linearly 

dependent upon the crack volume of the rock, thus implying that the resistivity is linearly related to the crack porosity. In lower crustal crystalline rocks, conduction is 

almost certainly through microcracks and along grain-edge 

tubules as opposed to pores in sedimentary rocks. The 

much more efficient connectivity of this pore geometry 

greatly reduces the amount of fluid needed to explain the 

resistivity of lower crustal rocks compared to Archies Law 

for sedimentary rocks. Secondly, the data from Quist and 

Marshall (1968) were for saline solutions only up to 0.1 

molar, whereas deep crustal fluids are almost certainly 

much more saline (Orville, 1963; Roedder, 1972). Recent 

laboratory work on more concentrated saline solutions 

(Ucok and others, 1980) show fluid resistivities about an 

order of magnitude lower than the results of Quist and 

Marshall (1968) at lower crustal conditions. 

The results of our MT inversions suggest that resistivities 

in the lower crustal fault zone averaged over a 5-km-wide 

zone may be as low as 50 Q-m. Concentrated saline solutions (20 wt percent NaCl) at temperatures of around 300Â° 

for the lower crust have resistivities of about 0.01 Q-m 

(Ucok and others, 1980). Using these values and an exponent of 1 in Archies Law yields a connected porosity of 

approximately 0.02 percent for a 50 Q-m, 5-km-wide fault 

zone if the connections are efficient. We should really 

multiply this by three due to inefficiencies in tubule connection, thus yielding a porosity of 0.06 percent. (Compare this with a 0.0001 percent porosity estimate for the 

lower crust away from the fault zone with a resistivity of 

30,000 Q-m.) If the same amount of conductance were 

put into a fault zone 1 km wide, it would need to have a 

resistivity of 10 Q-m to give the same leakage effect. 

This translates to a porosity of 0.3 percent. Similarly, a 

100-m-wide fault zone zone would require a connected 

porosity of 3.0 percent. 

If the lower crustal part of the San Andreas fault zone 

is as conductive as our inversions suggest, then this will 

have important implications for the fluid regime of the 

lower crustal part of fault zones. The question we cannot 

answer is whether the inferred high conductance of the 

lower crustal part of the fault zone is an inherent feature 

of fault systems, or whether it results from the earthquake 

itself. We do not believe that it is a permanent feature of 

the fault zone because non-elastic healing processes over 

time would close up much of the porosity connectivity, 

thus decreasing the conductivity. We believe a more likely 

scenario is that the hydraulics and conductivity of fault 

zones are time-dependent. While brittle failure occurred 

down to at least 18 km depth, any pore-pressure buildup 

before the earthquake associated with the closing of permeable pathways would increase the likelihood of semibrittle failure down to depths below the main shock of the 

earthquake. This may have increased the efficiency of 

fluid connectivity in and around the lower crustal part of 

the San Andreas fault zone. This hypothesis is based on 

the difficulty of bringing pore fluids into the lower crustal 

zone in such a small time scale (one year). However, it 

also requires that these fluids were inefficiently connected 

before the earthquake, which would probably lead to high 

fluid pressures (lithostatic) and a weakening of the mechanical strength of the lower crustal part of the fault 

zone. This is a very important issue for understanding the 

earthquake cycle (Rice, 1992), but without having analyzed data from before the earthquake, we cannot be any 

more than speculative about these interpretations. 

CONCLUSIONS 

Figure 13.-Plot of the observed and predicted TM-mode phases 

for the inversion results with and without an a priori narrow conductive 

fault zone. Note that the data are fit only when there is lower crustal 

leakage around the San Andreas fault zone. 

We conducted a wide-band MT survey around the Loma 

Prieta region one year after the Loma Prieta earthquake. 

The survey was designed to look for lower crustal electri- 
period (seconds) 

MIT Site 1 1, Watsonville 

10Ã 

Site 10, Mt Madonna 

10Ã 

104 

E 

u 

c 

period (seconds) 

: 

- 

- 

Site 9, Uvas 

.- Sf 10'; 

5 

? 

- 10' 

E TM mode 

TE mode 

21 , , ,,,,," , , ,,,,,,, , , ,,,,,,, , , , ,,,,,, , , ,,,,,,, , , ,,,,,,, , , ,J -60 10-3 lo-a 10-1 ino 101 id* in* in* 

period (seconds) 

Figure 14.Ã‘Observe and predicted data at (left) site 11, Watsonville (west of the San Andreas the a priori model and all changes were forced into the upper crust. We show both the predicted TMfault), (middle) site 10, Mt. Maddona (near the fault), and (right) site 9, Uvas (east of the San and TE-mode data, even though only the TM-mode data were used in the analysis. The predicted TEAndreas fault). The solid lines are the predicted data from an inversion with an a priori narrow fault mode data were static shifted so that they agreed with the observed TE-mode data at the highest 

zone, and the dashed lines are the predicted data from an inversion where no fault zone was put into frequencies. 
A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D309 

cal resistivity anomalies both across the fault zone and 

along it (from north to south). Local industrial noise 

seemed to contaminate the data at some sites in the band 

from 1 to 100 s, and at those sites, we did not use the data 

in that band in the analyses. 

The results of our two-dimensional TM-mode inversions of the data suggested that vertical zones of enhanced 

lower crustal conductivity exist in the vicinity of the San 

Andreas fault. If we forced the changes to be put into a 5- 

km-wide zone, then we obtained lower crustal resistivity 

values as low as 50 a-m in that zone. To explain such 

dramatic conductivities relative to the lower crust away 

from the fault zone, we believe that fluids were present in 

the lower crustal part of the fault zone before the earthquake, but were inefficiently connected, and that the earthquake and its associated aftershocks increased the 

efficiency of fluid connectivity in that zone. An alternate 

interpretation is that the conductance of the fault zone is 

an inherent feature of that zone, but laboratory work suggests that non-elastic healing processes would eliminate 

much of the porosity connectivity, thus reducing the conductivity of that zone. Although our data alone cannot 

distinguish between the two interpretations, we believe 

the time-dependent interpretation is more consistent with 

the mechanics of the lower crust. Without having analyzed data from before the earthquake, it is impossible to 

test this hypothesis. 

ACKNOWLEDGMENTS 

We would like to thank Frank Morrison, Electromagnetic Instruments Inc., and the students from U.C. Berkeley for the tremendous job they did in collecting the MT 

field data and putting up with poison oak, thick brush, 

and long hours in the field. Useful discussions were held 

with Dal Stanley, Frank Morrison, Brian Evans, Gary 

Egbert, Steve Park, Jim Larsen, and Jim Rice. Reviews by 

Dal Stanley, Paul Reasenberg, and an anonymous reviewer 

greatly improved the manuscript. The CRAY X-MP EAI 

464 at the MIT supercomputer facility was used for running all the inversions. This work was supported in full 

by NSF grant 901 19 19-EAR. 

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D310 AFTERSHOCKS AND POSTSEISMIC EFFECTS 

Model 1, conductive window in lower crust 

-10 - watsonville TM phase 

-20 - 

-50 - 

-60 [- uvas TM phase 

period (seconds) 

Model 2, uniform lower crust and mantle 

watsonville TM phas 

. mt madonna TM phase 

uvas TM phase 

- 

- 

- 

- 

- 

- 

I I I I I I Ill I I I IIIIII I I I Illlll I I I 1 I Ill1 I I Ill1111 I I I I IIIII I I I Illlll 

I o-~ I o-~ 10-1 I o" lo1 I 02 I o3 I 04 

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A MAGNETOTELLURIC SURVEY: IMPLICATIONS FOR EARTHQUAKE PROCESSES AND CONDUCTIVITY D311 

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Figure 15.-Plot of the observed and predicted TM-mode phases for the 

inversion results with and without an a priori narrow conductive fault 

zone. It is difficult to tell much difference between the two models 

within the spread of the data, but we believe that data are fit slightly 

better when there is a somewhat conductive fault zone. 

Figure 16.-Observed and predicted data at site 7, Hollister (east of the 

San Andreas fault). The solid lines are the predicted data from an inversion with an a priori narrow fault zone, and the dashed lines are the 

predicted data from an inversion where no fault zone was put into the a 

priori model and all changes were forced into the upper crust. We show 

both the predicted TM- and TE-mode data, even though only the TMmode data were used in the analysis. Both models fit the data equally 

well. 
f AAA^' 1 &fi~* 

A TE mode 

period (seconds) 

* U.S. GOVERNMENT PRINTING OFFICE: 1997 - 573-046 / 20077 REGION NO. 8