Professiona l Pape r 1550– E

USGS
The Loma Prieta, California, Earthquake of 

October 17, 1989—Geologic Setting and

Crustal Structure

RAY E. WELLS, Editor

EARTHQUAKE OCCURRENCE

William H. Bakun and William H. Prescott, Coordinators

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1550–E

Prepared in cooperation with the National Science Foundation
U.S. DEPARTMENT OF THE INTERIOR

GALE A. NORTON, Secretary

U.S. GEOLOGICAL SURVEY

Charles G. Groat, Director

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U.S. Geological Survey, Reston, Virginia: 2004

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Produced in the Western Region, Menlo Park, California

Manuscript approved for publication, May 14, 2003

Text edited by George A. Havach 

Layout and design by Stephen L. Scott 

Library of Congress catalog-card No. 92–32287
CONTENTS

 Pages

Introduction------------------------------------------------------------------------ E1 

 By Ray E. Wells

Stratigraphy and structure across the San Andreas fault zone in the

 Loma Prieta region and deformation during the earthquake----------- 5 

 By Robert J. McLaughlin and Joseph C. Clark

Geophysical and geologic setting of the earthquake, inferred from 

 gravity and magnetic anomalies-------------------------------------------- 49 

 By Robert C. Jachens and Andrew Griscom

Deep geoelectrical structure of the Loma Prieta region---------------------- 81 

 By Brian D. Rodriguez, Victor F. Labson, William D. Stanley, 

 and Jay A. Simpson

Evidence for oceanic crust beneath the continental margin west of 

 the San Andreas fault from onshore-offshore wide-angle 

 seismic recordings----------------------------------------------------------- 107 

 By Thomas M. Brocher, Michael J. Moses, and Stephen D. Lewis

Three-dimensional tomographic analysis of the Loma Prieta region---- -- 127 

 By Eylon Shalev and Jonathan M. Lees

Three-dimensional variations in VP /VS ratio in the Loma Prieta 

 region------------------------------------------------------------------------ -- 143

 By Shashank R. Atre and Clifford H. Thurber

Seismotectonics of the Loma Prieta region determined from three-

 dimensional P-wave velocities, VP /VS ratios, and seismicity----------- 165 

 By Donna M. Eberhart-Phillips and Andrew J. Michael

Teleseismic tomography of the Loma Prieta region: Implications 

 for strain partitioning-------------------------------------------------------- 189

 By Yoko Takauchi and John R. Evans

III
PLATES

[Plates are in pocket]

1. Geologic map of the Loma Prieta region, California 

2. Stratigraphic columns across the San Andreas fault in the Loma Prieta 

 region, California, comparing differences in thickness, lithology, and 

 basement rock types in major fault blocks

3. Maps showing residual-magnetic fields, isostatic-residual-gravity fields, 

 and generalized geology of the southern San Francisco Bay and Loma 

 Prieta regions, California

IV
CONTENTS

————— Page

Background------------------------------------------------------------------- E1

Geologic setting-------------------------------------------------------------- 1

Crustal structure from potential-field and electromagnetic methods-- 2

Seismic velocity structure of the crust and mantle----------------------- 2

Concluding remarks---------------------------------------------------------- 4

—————

BACKGROUND

Although some scientists considered the Ms=7.1 Loma 

Prieta, Calif., earthquake of 1989 to be an anticipated event, 

some aspects of the earthquake were surprising. It occurred 

17 km beneath the Santa Cruz Mountains along a left-stepping restraining bend in the San Andreas fault system. 

Rupture on the southwest-dipping fault plane consisted of 

subequal amounts of right-lateral and reverse motion but 

did not reach the surface. In the area of maximum uplift, 

severe shaking and numerous ground cracks occurred along 

Summit Road and Skyland Ridge, several kilometers south of 

the main trace of the San Andreas fault. The relatively deep 

focus of the earthquake, the distribution of ground failure, the 

absence of throughgoing surface rupture on the San Andreas 

fault, and the large component of uplift raised several questions about the relation of the 1989 Loma Prieta earthquake 

to the San Andreas fault: Did the earthquake actually occur 

on the San Andreas fault? Where exactly is the San Andreas 

fault in the heavily forested Santa Cruz Mountains, and how 

does the fault relate to ground ruptures that occurred there 

in 1989 and 1906? What is the geometry of the San Andreas 

fault system at depth, and how does it relate to the major 

crustal blocks identified by geologic mapping?

Subsequent geophysical and geologic investigations of 

crustal structure in the Loma Prieta region have addressed 

these and other questions about the relation of the earthquake to geologic structures observed in the southern Santa 

Cruz Mountains. The diverse papers in this chapter cover 

several topics: geologic mapping of the region, potentialfield and electromagnetic modeling of crustal structure, and 

the velocity structure of the crust and mantle in and below 

the source region for the earthquake. Although these papers 

were mostly completed between 1992 and 1997, they provide 

critical documentation of the crustal structure of the Loma 

Prieta region. Together, they present a remarkably coherent, 

three-dimensional picture of the earthquake source region—a 

geologically complex volume of crust with a long history 

of both right-lateral faulting and fault-normal compression, 

thrusting, and uplift.

GEOLOGIC SETTING

McLaughlin and Clark (this chapter) place the earthquake 

within the context of the Late Cenozoic geologic history of 

strike-slip faulting, folding, and thrusting in the southern 

Santa Cruz Mountains. They present detailed geologic maps 

and cross sections covering a 700-km2 area of the Loma Prieta 

region and note that the geologic histories on either side of 

the San Andreas fault differ considerably. The San Andreas 

fault is defined by a clear bedrock trace which separates 

terranes that have undergone hundreds of kilometers of 

right-lateral motion.

Late Cenozoic displacements on the San Andreas and 

its subsidiary faults have juxtaposed three major geologic 

terranes in the region. West of the Zayante-Vergeles fault, 

Tertiary marine strata overlie Salinian granodiorite basement 

that has been transported northward from southern California. Between the Zayante and San Andreas faults, the thick, 

folded Cenozoic marine-basin sequence of the La Honda 

block is underlain by the gabbro of Logan, a Jurassic terrane 

of oceanic affinity that is also displaced northward 300 km 

from Eagle Rest Peak near the Tehachapi Mountains. East of 

the San Andreas fault, McLaughlin and Clark map a series of 

fault slivers consisting of Franciscan Complex subterranes 

and overlying Mesozoic Great Valley sequence and Tertiary 

sedimentary strata.

The mapped terrane-boundary faults, with the possible 

exception of the San Andreas fault itself, are largely southwest-dipping thrust or reverse faults that have accommodated significant shortening normal to the plate boundary, 

in addition to strike-slip motion. Some of these faults—for 

example, the Monte Vista, Berrocal, and Shannon faults—are 

associated with shallow seismicity, have youthful geomorphic features, and are interpreted by McLaughlin and Clark 

to be active structures accommodating uplift of the eastern 

Santa Cruz Mountains. During the earthquake, triggered 

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

INTRODUCTION

By Ray E. Wells, U.S. Geological Survey

E1
E2 EARTHQUAKE OCCURRENCE

slip occurred on some of the foothills thrust faults along the 

southern margin of the Santa Clara Valley. McLaughlin and 

Clark view the uplift and shortening associated with the Loma 

Prieta earthquake to be consistent with the history of recent 

thrust faulting in the Santa Cruz Mountains.

Coseismic ground cracking along Summit Road and Skyland 

Ridge followed a zone of previously mapped echelon faults 

along ridgetops south of the San Andreas fault. McLaughlin 

and Clark note that this was also the zone of maximum uplift 

during the earthquake and suggest that lateral and vertical 

inhomogeneities in fault slip and crustal composition near 

the northern tip of the rupture zone could have contributed 

to focusing of tectonic deformation in the Summit Road area. 

They also note, however, that postearthquake investigations 

demonstrated that the displacements on ground cracks tended 

to follow the regional slope and that many of these ground 

cracks were associated with large landslide complexes. In 

the San Andreas rift valley, the mapped bedrock trace of the 

San Andreas fault is sinuous, and the bends correlate with the 

toes of large, coherent-block landslide complexes and zones 

of detachment faulting along the valley walls. For example, at 

the Wrights railroad tunnel, right-lateral displacement of the 

tunnel in 1906 was almost 400 m south of the mapped bedrock trace of the San Andreas fault. McLaughlin and Clark’s 

mapping suggests that a history of gravitational collapse 

of topography during seismic shaking has left a substantial 

imprint on the tectonic geomorphology of the San Andreas 

fault in the Santa Cruz Mountains.

CRUSTAL STRUCTURE FROM 

POTENTIAL-FIELD AND 

ELECTROMAGNETIC METHODS

Guided by geologic constraints, Jachens and Griscom (this 

chapter) interpret gravity and magnetic-field anomalies to 

construct a model of the subsurface crustal structure and composition beneath the Santa Cruz Mountains and southern Santa 

Clara Valley. In their model, the Loma Prieta region consists of 

seven crustal blocks juxtaposed by faults in the San Andreas 

system. Southwest of the San Andreas fault, two nonmagnetic 

basement terranes of Salinian granodiorite are separated by 

a magnetic basement terrane correlated with the Jurassic 

gabbro of Logan. A Tertiary sedimentary-basin sequence, as 

much as 10 km thick, overlies the gabbro, and both blocks are 

interpreted to be thrust beneath the Ben Lomond block along 

a southwest-dipping Zayante-Vergeles fault. All three blocks 

are truncated by the San Andreas fault on the east and by the 

San Gregorio fault on the west.

Northeast of the San Andreas fault, curvilinear magnetic 

anomalies delineate the boundaries of four tectonic blocks 

consisting of Franciscan Complex and overlying Coast Range 

ophiolite, Great Valley sequence, and Tertiary strata. Jachens 

and Griscom interpret these terranes to be the uplifted remnants of a regional Franciscan tectonic wedge emplaced onto 

Sierran basement along the 600-km-long Coast Range-Great 

Valley contact to the east. Its exposure in the Santa Cruz 

Mountains resulted from Late Cenozoic northward transport 

by the San Andreas fault system.

A comparison of the crustal model with the aftershock 

distribution suggests that most events are attributable to the 

San Andreas, Zayante-Vergeles, and Sargeant faults. Epicenters of aftershocks that plot off the main fault traces are 

commonly associated with dipping fault segments, and the 

San Andreas fault is interpreted to dip steeply southwest to 

vertical at depth. Within 1 km of the surface, the fault may 

locally dip gently to the northeast, putting nonmagnetic rocks 

over magnetic Purisima Formation on the northeast side of 

the surface trace. This unusual geometry coincides with 

southward-dipping Tertiary detachment faults northeast of the 

San Andreas fault (see McLaughlin and Clark, this chapter). 

One possible explanation for this relation is suggested by the 

widespread, structurally controlled ground cracking, extension, and downslope ground displacements that have occurred 

along ridgetops in the Loma Prieta region. During major 

earthquake shaking, the detachment faults are appropriately 

oriented to accommodate downslope transport of tectonic 

slivers across the San Andreas fault.

A deep geoelectric survey of the Loma Prieta region by 

Rodriguez and others (this chapter) confirms that lithologically 

distinctive crustal blocks are bounded by the San Andreas and 

Zayante-Vergeles faults at depth. High resistivities are characteristic of the Salinian granodiorite block southwest of the 

Zayante-Vergeles fault and of a large block of Franciscan rocks 

to the northeast of the San Andreas fault. A low-resistivity 

region in the uppermost 7 km of crust between the Zayante and 

San Andreas faults correlates with the thick, folded Tertiary 

marine sedimentary sequence mapped between the two faults. 

At midcrustal depths, a low-resistivity channel that extends 

to 19-km depth beneath the San Andreas fault is interpreted 

to represent fluids in the fault zone. The heterogeneity of 

the crust in the Loma Prieta region contrasts with the more 

uniform geoelectric structure observed in the Gabilan Range 

profile across the creeping section of the San Andreas fault 

to the south.

SEISMIC-VELOCITY STRUCTURE

OF THE CRUST AND MANTLE

Brocher and others (this chapter) have used onshore recordings of airgun profiles offshore of Santa Cruz and Año Nuevo 

to construct a velocity model of the continental margin in the 

Loma Prieta region west of the San Andreas fault. The crust 

consists of two major components: a landward-thickening 

wedge of rocks with 6.0- to 6.5-km/s velocities that is consistent with the granitic or quartz monzonite crust of the Salinia 

terrane, and an underlying 5-km-thick, 6.8- to 7.1-km/s-velocity mafic unit, dipping gently east beneath the continental 

margin and apparently continuous with the offshore Pacific 
E3

plate. Brocher and others consider that this mafic layer may be 

arc-derived lower crust or a magmatically underplated layer, 

but they prefer to interpret this high-velocity layer as a slab of 

oceanic crust, possibly Pacific plate, that has underthrust the 

margin during Pliocene-Quaternary plate reorganization and 

crustal extension in the Basin and Range Province. They note 

that similar relations are found along the coast southward to 

Morro Bay. The hypocenter of the earthquake at about 17-km 

depth occurred in the high-velocity layer in the lower crust, 

consistent with previous correlations of high-velocity regions 

and earthquake rupture zones.

The westward-tapering tectonic wedge west of the San 

Andreas fault zone mirrors the tectonic wedge of Franciscan 

Complex emplaced eastward over Sierran basement on the 

east side of the San Andreas fault, and is thought to accommodate some of the shortening perpendicular to the San 

Andreas fault. The resulting concave-upward Moho may 

help to focus seismic energy from earthquakes along the San 

Andreas fault zone.

In a series of three papers on the seismic tomography of the 

Loma Prieta region, Shalev and Lees (this chapter), Atre and 

Thurber (this chapter), and Eberhart-Phillips and Michael (this 

chapter) investigate the relation between the three-dimensional 

velocity structure of the source region, its geologic structure, 

and the earthquake process. Shalev and Lees used raypaths 

from 844 aftershocks to generate a three-dimensional P-wavevelocity model of the Loma Prieta region. Their model clearly 

images a low-velocity region in the uppermost 10 km of crust 

between the San Andreas and Zayante-Vergeles faults. This 

feature coincides with tightly folded Tertiary strata of the 

La Honda Basin as mapped by McLaughlin and Clark (this 

chapter), a deep sedimentary basin also characterized by its 

low density and low resistivity (see Jachens and Griscom, this 

chapter; Rodriguez and others, this chapter). Between 6-km 

depth and the lower limit of the model at 16-km depth, a highvelocity prism dips steeply westward beneath the San Andreas 

fault and coincides with the zone of aftershocks below 6-km 

depth. Shalev and Lees offer two interpretations of the highvelocity body: a more rigid, high-velocity rock that acts as an 

asperity along the fault, or, as suggested by its low resistivity, 

a weaker, serpentinite-bearing fault zone.

Atre and Thurber (this chapter) recognize significant threedimensional variations in VP /VS ratio in the Loma Prieta 

region. By combining their results with independent VP

data and the gravity and magnetic constraints, they examine 

variations in crustal composition across the major faults. 

Southwest of the Zayante-Vergeles fault, high VP values and 

low VP /VS ratios in the upper 7 km of the crust are consistent 

with the granodioritic composition of the Salinian terrane 

and support the existence of a Zayante-Vergeles fault to at 

least 7-km depth. High VP /VS ratios at 3-km depth northeast 

of the Zayante-Vergeles fault correlate with porous Tertiary 

sedimentary rocks of the La Honda Basin and continue across 

the San Andreas fault into heterogeneous, ophiolite-bearing 

thrust packages of the Franciscan Complex. At 7-km depth, 

high VP /VS ratios of 1.84 to 2.0 suggest gabbroic compositions beneath the sedimentary layers, consistent with geologic 

and potential-field data indicating gabbro of Logan beneath 

the La Honda Basin. A major transition to low VP /VS ratios 

below 11-km depth could reflect a Salinian lower crust or 

may reflect healing of microcracks and a relative increase in 

S-wave velocities.

In a major synthesis, Eberhart-Phillips and Michael (this 

chapter) present three-dimensional VP and VP /VS models of 

the Loma Prieta region derived from inversion of earthquake 

and explosive source traveltimes recorded at 1,700 permanent 

and temporary stations. The P-wave-velocity model and

 the relocated earthquake hypocenters reveal a high-P-wavevelocity body below 6-km depth that extends the length of 

the Loma Prieta rupture and separates it from a near-vertical 

San Andreas fault 5 km to the northeast. Eberhart-Phillips 

and Michael suggest that this body controlled not only the 

extent of the rupture but also the pattern of background 

seismicity on the adjacent San Andreas and Sargent faults, 

thus explaining how the earthquake filled in a gap in the 

regional microseismicity. As in their previous studies, the 

asperities with the greatest moment release correspond to 

regions where higher velocity material is in contact with the 

fault. The upward termination of the Loma Prieta rupture 

correlates with its propagation into a volume of low-density 

Tertiary sedimentary rocks with low P-wave velocities and 

high VP /VS ratios, possibly reflecting high pore pressures and 

different frictional behavior.

Northeast of the San Andreas fault, the P-wave-velocity model images a southwest-dipping thrust zone correlative with the southwestern Santa Clara Valley thrust 

belt, which places Franciscan Complex and other Mesozoic

 terranes over Late Cenozoic and Quaternary valley fill (see 

McLaughlin and Clark, this chapter). Loma Prieta aftershocks tend to be concentrated in the hanging wall of this 

thrust, and most events have thrust mechanisms. Aftershock 

focal mechanisms in the main-shock rupture zone indicate a 

completely heterogeneous stress field after the main shock 

and do not support the uniform fault-normal compressionalstress field expected after total stress drop on the main-shock 

plane. Eberhart-Phillips and Michael consider both the Loma 

Prieta fault surface and the San Andreas fault to be parts of 

a complex, 5-km-wide San Andreas fault zone with several 

potential rupture surfaces. They also conclude that this section 

of the San Andreas fault zone does not have a single type of 

characteristic earthquake.

Takauchi and Evans (this chapter) use teleseismic travel 

time residuals to determine the velocity structure of the lower 

crust and upper mantle in the Loma Prieta region. They find 

horizontal velocity gradients across the San Andreas fault at 

all depths, including the upper mantle, suggesting that the 

fault has been a steep, narrow locus of shear. A low-velocity 

feature near the Moho beneath Loma Prieta peak is interpreted 

as a crustal root, 2 to 5 km thick, at the base of the crust. 

Two upper-mantle low-velocity features between 45- and 

INTRODUCTION
E4 EARTHQUAKE OCCURRENCE

70 -km depth are separated by 45 km right-laterally across the 

San Andreas fault. Takauchi and Evans note that this offset 

is small in comparison with the total offset across the fault 

and is compatible with models in which deep mantle slip is 

now accommodated beneath the Hayward-Calaveras faults 

to the east.

CONCLUDING REMARKS

Geologic and geophysical investigations of the crustal 

structure in the Loma Prieta region indicate that this region 

consists of a complex assemblage of crustal blocks, juxtaposed by hundreds of kilometers of right-lateral displacement and related imbricate thrusting along the San Andreas 

fault system. The structural complexities extend to depth. 

Although the San Andreas fault can be mapped as a throughgoing dextral-slip fault, the southwest-dipping Loma Prieta 

rupture surface appears to be a separate strand and is one of 

several southwest-dipping thrusts along this section of the 

San Andreas fault system. Several fault strands are capable 

of accommodating plate motion, and more than one type of 

characteristic earthquake is probable along this section of the 

San Andreas fault system.

At depth, the composition of geologic materials in the fault 

zone may have controlled the distribution of fault slip during 

the earthquake. High-velocity and high-density rocks correlate with asperities, or zones of high slip, on the fault plane, 

whereas the updip termination of the Loma Prieta rupture 

correlates with weak, low-density sedimentary fill in the La 

Honda Basin.

Near the surface, tectonic mountain-building processes in 

the Santa Cruz Mountains compete with partial gravitational 

collapse of the topography during intense seismic shaking. 

The abundant coseismic extensional ground cracking along 

ridgetops, a correlation between the trace sinuosity of the 

San Andreas fault and large slope failures along the riftvalley wall, and geophysical evidence for shallow thrusting 

in the bottom of the rift-valley all suggest that large-scale 

downslope movement along preexisting geologic structures 

may characteristically overprint tectonic deformation in the 

Loma Prieta region.
E5

CONTENTS

————— Page

Abstract----------------------------------------------------------------------- E5 

Introduction------------------------------------------------------------------- 6 

Stratigraphy southwest of the San Andreas fault------------------------ 7 

 Basement rocks of the Salinia terrane southwest of the 

 Zayante fault (Cretaceous and older)--------------------------- 7 

 Mafic basement rocks of the Salinia terrane between the 

 Zayante and San Andreas fault zones (Jurassic)-------------- 7 

 Tertiary section southwest of the San Andreas fault-------------- 13 

 Possible buried Paleocene sequence------------------------- 13 

 Eocene to lower Miocene sequence------------------------- 13 

 Lower to middle Miocene sequence------------------------- 16 

 Upper Miocene to Pliocene sequence----------------------- 16 

Stratigraphy northeast of the San Andreas fault------------------------- 16 

Franciscan Complex (Jurassic and Cretaceous)------------------------- 17 

 Slate and phyllite of Loma Prieta peak (Jurassic?)--------------- 17 

 Coast Range ophiolite (Jurassic)------------------------------------ 18 

 Great Valley sequence (Upper Jurassic to Upper Cretaceous)-- 18 

 Upper Jurassic and Lower Cretaceous rocks--------------- 19 

 Upper Cretaceous rocks--------------------------------------- 19 

 Eocene sequence------------------------------------------------------ 19 

 Lower Eocene to lower Miocene sequence------------------------ 20 

 Pre-Pliocene sequence northeast of the El Sombroso and 

 Sierra Azul Blocks------------------------------------------- 20 

 Temblor Sandstone (Oligocene to middle Miocene)------ 20 

 Monterey Shale (middle Miocene)-------------------------- 20 

 Quaternary and Tertiary nonmarine deposits northeast 

 of the San Andreas fault------------------------------------- 20 

 Santa Clara Formation (Pliocene and Pleistocene) 20 

 Alluvial and landslide deposits (Pleistocene and 

 Holocene)----------------------------------------------------- 21

Structure southwest of the San Andreas fault---------------------------- 21 

 Zayante-Vergeles and related buried faults------------------------ 22 

 Butano fault------------------------------------------------------------ 23 

 San Andreas fault zone----------------------------------------------- 23 

 San Andreas fault in the Loma Prieta region---------------------- 26 

 Faults of the Summit Road-Skyland Ridge area------------------ 26 

 Earthquake deformation in the Summit Road-Skyland 

 Ridge fault zone---------------------------------------------- 27

Structure northeast of the San Andreas fault----------------------------- 28 

 Sargent fault----------------------------------------------------------- 29 

 Hooker Gulch-Sierra Azul fault zone------------------------------ 30 

 Aldercroft fault-------------------------------------------------------- 31

 Faults of the southwestern Santa Clara Valley thrust belt------- 31 

 Age of thrust-belt initiation and amount of displacement------- 33 

 Lexington fault zone-------------------------------------------------- 36 

Modeling the structural evolution of the Loma Prieta region---------- 38 

 Structural and stratigraphic constraints---------------------------- 38 

 Deep-crustal constraints---------------------------------------------- 38 

 Wedge tectonic model------------------------------------------------ 40 

 Page 

1989 Loma Prieta earthquake---------------------------------------------- 41 

Conclusions------------------------------------------------------------------- 41 

Acknowledgments----------------------------------------------------------- 42 

References cited-------------------------------------------------------------- 42

 —————

ABSTRACT

The Loma Prieta region is divided by the San Andreas 

and Zayante faults into three structural blocks with different 

basement terranes. Tertiary strata that overlie the basements 

of these structural blocks record separate but overlapping 

histories of uplift, shortening, and strike-slip events, the most 

recent of which was the 1989 Loma Prieta earthquake.

Cretaceous (≥91 Ma) granitic rocks intrusive into older 

high-grade metasedimentary rocks of the Salinia terrane are 

exposed locally in the Ben Lomond block southwest of the 

Zayante fault. Northeast of the Zayante fault, in the La Honda 

block, the basement is covered by a Tertiary sedimentary 

section, as much as 6 km thick. High magnetization and high 

density suggest that this buried basement probably represents 

the northwestward continuation of the bodies of Jurassic 

(165–161 Ma) gabbro and undated anorthositic gabbro near 

Logan and San Juan Bautista.

The juxtaposition of continental Salinian basement with 

oceanic Logan basement along the Zayante fault apparently 

required significant latest Cretaceous to early Eocene dipslip displacement. As much as 3 to 4 km of additional dip 

slip along the Zayante fault during Oligocene (Zemorrian) 

to early Miocene (Relizian) time resulted in unroofing of the 

Salinia terrane to the south and thick sediment accumulation 

and volcanism in the La Honda block to the north. Regional 

transpression during the Miocene and Pliocene (when it was 

most pronounced) reactivated this fault. Aftershocks of the 

1989 Loma Prieta earthquake associated with the Zayante fault 

at 7- to 10-km depth suggest secondary fault reactivation, but 

that slip did not propagate to the surface.

The San Francisco Bay block northeast of the San Andreas 

fault and southwest of the Silver Creek fault consists of a 

composite basement of Franciscan Complex, Coast Range 

ophiolite, and Great Valley sequence. Seismic data suggest 

that these geologic units occur from the surface to 8- to 10-

km depth beneath the Loma Prieta region. Early Jurassic to 

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE IN THE 

LOMA PRIETA REGION AND DEFORMATION DURING THE EARTHQUAKE 

By Robert J. McLaughlin, U.S. Geological Survey; 

and

 Joseph C. Clark, Indiana University of Pennsylvania

E5
E6 EARTHQUAKE OCCURRENCE

Late Cretaceous metaclastic, metaigneous, and pelagic rocks 

of the Bald Mountain-El Sombroso and Permanente terranes 

of the Franciscan Central belt, and Jurassic (181?–156 Ma) 

ophiolitic rocks of the Coast Range ophiolite, compose most 

of the exposed basement. Beneath Loma Prieta peak, a thin 

fault sliver of slate and phyllite resembling Jurassic rocks of 

the Sierra Nevada foothills in metamorphic grade, composition, and texture is interpreted to have been thrust upward and 

northeastward from a separate metamorphic basement that 

may underlie Coast Range ophiolite and Franciscan rocks.

The San Francisco Bay block is broken into five subsidiary 

fault blocks by a system of dextral reverse-slip faults that 

include the Sargent, Hooker Gulch, Soda Springs, Sierra 

Azul, Berrocal, and Shannon faults. Most of these boundary faults are delineated by discontinuous lenticular bodies 

of mafic to ultramafic igneous rocks derived from the Coast 

Range ophiolite.

A major change in Oligocene to Miocene stratigraphy 

occurs across the Berrocal and Sargent faults, which partly 

form the boundary between the El Sombroso and Sierra Azul 

subsidiary blocks to the southwest and the Uvas, Los Capitancillos, and Santa Teresa subsidiary blocks to the northeast. 

Along the Sargent fault, as much as 1.4 km of uplift, 1.5 to 

1.8 km of reverse slip, and 0.5 to 1.1 km of shortening probably began in the early Miocene (18 Ma) and may have been 

largely completed by the early late Miocene (10 Ma).

The southwestern Santa Clara Valley thrust belt contains 

several anastomosing faults northeast of the Sargent fault, 

including the Lime Kiln, Berrocal, Shannon, and Monte 

Vista faults. As much as 3.0 km of shortening and uplift and 

4.2 km of reverse slip have occurred across this thrust belt 

since the early late Miocene (10 Ma). Faults of the thrust belt 

locally exhibit evidence for surface displacement within the 

past 15–18 k.y. Compressional deformation along the Monte 

Vista and Berrocal-Shannon fault zones during the earthquake 

suggests that these thrusts interact with the San Andreas fault 

and (or) that uplift and shortening across the San Andreas fault 

zone is accommodated by these thrust faults.

Tectonic evolution of the Loma Prieta region has included 

13 to 15 km of unroofing of the Coast Ranges, which occurred 

largely from 95 to 70 Ma southwest of the San Andreas fault 

and from about 65 to 56 Ma northeast of the fault. Southwest 

of the San Andreas fault, as much as 4 km of uplift and basaltic 

volcanism occurred during the Oligocene to early Miocene 

across the Zayante fault. Northeast of the San Andreas fault, 

Oligocene to early Miocene uplift and erosion was associated 

with low-angle attenuation faulting. Oblique transpressional 

deformation began in the early Miocene southwest and northeast of the San Andreas fault and became most pronounced 

during the late Pliocene (3 Ma), producing cumulative uplift 

of about 6 km and shortening of about 5 km. This transpressional deformation was largely concurrent with 480 km 

of Miocene and later aggregate dextral slip across the San 

Andreas, Pilarcitos, and San Gregorio faults and faults of the 

eastern San Francisco Bay region. About 136 km of this slip 

is attributable to displacement along the Pilarcitos fault and 

the Peninsular segment of the San Andreas fault.

A model involving tectonic wedging is here invoked to 

explain the relation between uplift, blind thrusting, and strikeslip faulting during the earthquake. The 17-km depth of the 

main shock approximates the postulated depth to the floorthrust zone of a tectonic wedge complex. The Loma Prieta 

rupture, therefore, may have propagated upward to 8- to 10-

km depth along a steeply dipping dextral reverse fault from 

a reactivated segment of the flat floor of the tectonic wedge 

complex. Secondary extensional surface deformation was 

focused along the Summit Road-Skyland Ridge fault zone, 

southwest of the principal trace of the San Andreas fault, and 

less intense compressional deformation was accommodated 

along eastward-vergent thrust faults of the southwestern Santa 

Clara Valley thrust belt.

Data collected by other workers indicate that the Summit 

Road-Skyland Ridge fault zone was activated in the 1906 San 

Francisco earthquake. Other evidence indicating the partitioning of secondary slip to preexisting Quaternary faults of the 

southwestern Santa Clara Valley thrust belt during the 1989 

Loma Prieta earthquake further suggests that Loma Prietatype events have occurred in the past and should be expected 

to recur in this region in the future.

INTRODUCTION

The epicenter of the 1989 Loma Prieta earthquake was 

located in the Santa Cruz Mountains (fig. 1), a structurally 

complex region of heavy vegetation and limited rock exposures. The San Andreas fault bisects the region, displacing 

rocks to the southwest hundreds of kilometers northwestward 

relative to rocks on the northeast side of the fault. These 

juxtaposed rocks are further divided into subsidiary fault 

blocks with different basements and different stratigraphic 

and structural histories.

The purpose of this paper is to describe the tectonic histories 

and geometries of the major structural blocks of the Loma 

Prieta region, as recorded by the stratigraphy and structure. 

This information places constraints on the timing of past 

movements and rates of slip on major faults in the region and 

suggests their relations to the San Andreas fault system. An 

understanding of these relations is important in evaluating the 

significance of the earthquake and of the potential for future 

damaging earthquakes in the San Francisco Bay region.

The essential features of the geology of the Loma Prieta 

region (fig. 1), including the existence of major reverse-slip 

faults linked with the San Andreas fault system, were well 

known by at least 1980. It was pointed out in 1974 that thrust 

faults along the southwest side of the Santa Clara Valley are 

linked at depth to the San Andreas fault and that these thrust 

faults, if linked to each other, would be capable of an M≥7.0 

event, on the basis of the half-length of the thrust system 

(McLaughlin, 1974; Sorg and McLaughlin, 1975). However, 
E7

the 70° SW.-dipping fault on which the 1989 Loma Prieta 

earthquake occurred was completely unknown until that 

event.

The seismic potential of thrust faults in northern California was unclear until the 1983 Coalinga, Calif., earthquake, 

which occurred on an eastward-vergent blind thrust beneath 

the Coast Ranges northeast of the heavily monitored Parkfield 

segment of the San Andreas fault (Wentworth and Zoback, 

1990). The possible significance of that earthquake to the 

potential for a major thrusting event in the San Francisco Bay 

region was still not fully appreciated until the 1989 Loma 

Prieta earthquake.

Since then, geologists have begun to view faulting in the 

northern Coast Ranges, including the San Francisco Bay 

region (fig. 1), much differently, in light of the enormous 

number of new data that have been accumulated regarding 

deep crustal structure. Much has been also learned about the 

distribution of blind thrusts. Previously unrecognized buried 

thrusts have been found to underlie areas of recently folded 

rocks in several areas. Buried thrusts could be more youthful 

and could pose a greater seismic hazard than thrusts that are 

mapped at the surface; however, the details of their earthquake recurrence and of how these faults interact with the 

major strike-slip faults of the North American plate transform 

boundary remain unclear and controversial. The 1989 Loma 

Prieta earthquake pointed out to us that a more thorough 

knowledge of the dynamics of these interactions is essential 

to predicting the timing or locations of future earthquakes in 

the San Francisco Bay region.

Just 2 years before the earthquake, we began a detailed mapping of four 7A-minute quadrangles astride the San Andreas 

fault in the southern Santa Cruz Mountains. The structural 

and stratigraphic relations discussed herein are based on this 

timely recent work, together with many years of previous U.S. 

Geological Survey investigations.

STRATIGRAPHY SOUTHWEST OF

THE SAN ANDREAS FAULT

The San Andreas fault in the Santa Cruz Mountains (figs. 

1, 2) juxtaposes a Mesozoic plutonic basement southwest 

of the fault with a different Mesozoic composite basement 

northeast of the fault, composed of the accretionary Franciscan Complex, ophiolitic rocks, and forearc-basin deposits 

of the Great Valley sequence. Since the early Miocene, these 

Mesozoic basement terranes have been displaced hundreds of 

kilometers across the San Andreas fault (fig. 1). Southwest of 

the San Andreas fault, the area that includes the Los Gatos, 

Laurel, and Loma Prieta quadrangles is divided by the Zayante and Vergeles faults into two structural blocks: the Ben 

Lomond block to the southwest and the La Honda block to the 

northeast (figs. 1, 2). In the Ben Lomond block, crystalline 

rocks of the Salinia terrane are locally exposed and occur at 

relatively shallow depths beneath Tertiary strata that are tilted 

and deformed into broad open folds. Except for a thin sliver 

of gabbro along the San Andreas fault between Logan and 

San Juan Bautista (fig. 1A), the basement northeast of the 

Zayante-Vergeles fault is completely buried by a Tertiary section, as much as 6 km thick, that is tightly folded and locally 

overturned (see pl. 1).

BASEMENT ROCKS OF

THE SALINIA TERRANE

SOUTHWEST OF THE ZAYANTE FAULT

(CRETACEOUS AND OLDER)

In the Loma Prieta region, granitic intrusive and highgrade metasedimentary rocks of the Salinia terrane form 

the basement southwest of the Zayante fault. Small isolated 

exposures in the Laurel quadrangle indicate that the granitic 

rocks range in composition from quartz diorite on the west 

to granodiorite and quartz monzonite on the east (see pl. 1). 

Pelitic schist crops out locally in the southwest corner of 

the quadrangle. These crystalline rocks have been mapped 

westward into the adjoining Felton quadrangle, where quartz 

diorite predominates in the more extensive exposures of Ben 

Lomond Mountain (fig. 1B; Clark, 1981, pl. 2). Exploratory 

wells south of the Zayante fault and its southeastward extension, the Vergeles fault, confirm the subsurface continuity of 

basement rocks of the Salinia terrane with granitic rocks in 

the northern Gabilan Range, 30 km to the southeast (Clark 

and Rietman, 1973, pl. 1; Ross and Brabb, 1973, fig. 2).

Radiometric (Rb-Sr and U-Pb) ages and structural data 

(Compton, 1966; Ross, 1978; Mattinson and James, 1985; 

James, 1992) suggest that granitic rocks of the Salinia terrane 

in the Santa Cruz Mountains were emplaced 103–91 Ma 

(mid-Cretaceous) at ≥18-km depths (Hansen and others, 

1991) and unroofed to less than 3- to 5-km depth from late 

Maastrichtian to early Paleocene (67–61 Ma) time (fig. 3; 

Arrowsmith and others, 1992; Bürgmann and others, 1994). 

Similarities in lithology and metamorphic grade suggest 

that the metasedimentary rocks are correlative with the Sur 

Series of Trask (1926). Although the depositional age of 

the Sur Series is uncertain, analysis of detrital zircons from 

schist and quartzite from Ben Lomond Mountain northwest 

of the city of Santa Cruz (fig. 1) indicates a mid-Cretaceous 

metamorphism for these metasedimentary rocks. The detrital 

zircons originally crystallized from an igneous melt as early 

as the Precambrian (1.7 Ga; James, 1984, 1992).

MAFIC BASEMENT ROCKS OF THE SALINIA 

TERRANE BETWEEN THE ZAYANTE AND 

SAN ANDREAS FAULT ZONES (JURASSIC)

Northeast of the Zayante fault and southwest of the San 

Andreas fault (fig. 1), the crystalline basement is buried 

beneath as much as 6 km of Tertiary rocks. A greater depth 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E8 EARTHQUAKE OCCURRENCE

Figure 1.—Loma Prieta region, Calif., showing major fault blocks and fault zones. A, Regional setting. BSF, Bartlett Springs fault; CA, 

Calaveras fault; CSZ, Cascadia subduction zone; FF, Franklin fault; GF, Garberville fault; GLF, Garlock fault; HAY, Hayward fault; HF, 

Hosgri fault; MF, Maacama fault; MFZ, Mendocino Fracture Zone; NAD, Navarro discontinuity; NSAF, northern section of the San Andreas 

fault (north of the San Francisco peninsula); PF, Pilarcitos fault; PFZ, Pioneer Fracture Zone; PLT, Pleito thrust; PRT, Pastoria-Rand thrust 

zone; RCF, Rodgers Creek fault; SAF, San Andreas fault, including Peninsular segment; SGF, San Gregorio fault; SNF, Sur-Nacimiento 

fault; TBF, Tolay-Bloomfield fault; ZVF, Zayante-Vergeles fault. B, San Francisco Bay block, showing locations of plate 1 and figure 2A. 

Star, epicenter of October 18, 1989, main shock.

Great Valley

KLAMATH

 BLOCK

GORDA

PLATE

VIZCAINO

BLOCK

PACIFIC

PLATE

PACIFIC

PLATE

SIERRA

NEVADA

SALINIAN BLOCK

SALINIAN BLOCK

CSZ

MFZ

GF

MF

BSF

NAD

NSAF

TBF

RCF

PFZ

HAY 

SF BAY

PF

BLOCK

ZVF

SGF

SAF

SNF HF

PLT PRT

GLF Neenach

Volcanics

Offset Permanente

terrane

NEW IDRIA

PIERCEMENT

MT. DIABLO

PIERCEMENT

Sutter Buttes

FF

CA

Pinnacles

Volcanics

Lake Berryessa

Clear Lake

Eureka

Point Arena

San

Francisco

Loma

Prieta

Peak

Logan

San Juan

Bautista

Paicines

Coalinga

Parkfield

Gold

Hill

Cholame

San Luis Obispo

Eagle Rest

Peak

Santa Barbara

41°

40°

39°

38°

37°

36°

35°

125° 124° 123° 122° 121° 120° 119° 118°

Area of

Figure 1b

PACIFIC OCEAN

?

Santa Cruz

Monterey

Bay

Watsonville

Hollister

Strata associated with Jurassic 

through Eocene forearc deposits and 

underlying Jurassic Coast Range 

ophiolite

Contact

Fault—Dashed where approximately located, showing relative 

movement

Thrust faultóSawteeth on upper plate

Former western edge of North American plate

Granitic basement rocks

Neogene volcanic rocks

Franciscan Complex

EXPLANATION

SAF

50 0 50 100 MILES

50 0 50 100 KILOMETERS

A
LOSGATOS

SATEH

LAUREL LO

PR

AREA OF FIGURE 2A

AREA OF 

121 

° 52'30

122° 122° 07'30''

Fault—Dashed where approximately located, showing relative movementThrust Fault—Sawteeth on upper plate

EXPLANATION

Shoreline

0 20 KILOMETERS

0

5

5 5 10 15 MILES

5 10 15

S A N 

F R A N C I S C O 

 

 B A Y 

B L O C K

E A S T 

B A Y 

 

B L O C K

CALAVERAS FAULT

CALA VER AS

FAULT

HAYWARD

FAULT

HAYWARD

FAUL T

S I LVER CRE EK FAULT

SA N

ANDREAS

FAUL T

S AN

ANDREA S

FA ULT

PILA RCI TO S

FAULT

SAN GR E G O R IO FAULT

L A

H O N D A

B L O C K

( G a b b r o i c b a s e m e n t )

( G a b b r o i c b a s e m e n t )

( G r a n i t i c - m e t a m o r ph i c b a s e m e n t)

B E N L O M O N D B L O C K

( G r a n i t i c - m e t a mo r p h i c b a s e m e n t)

MON TE VISTA FAUL T

B E R R O C A LF A U L T

L IMEKI LN F

BUTANO

FAUL T

(Composite ophiolite-Franciscan basement)

SHANNON FAULT

BERROCAL FAULT

SARGENT

FAULT

ZAYANTE-VERGELES FAULT

?

?

PIGEON

?

POINTBLOCK(Unexposedophiolitic-Franciscan(?)basement)

37° 15'

122° 15'

Corte Madera lithofaciesof Santa Clara Formation

LOMA PRIETA EARTHQUAK

MAINSHOCK EPICENTER

? ?

?

?

?

?

PACIFIC

OCEAN

San Francisco

Woodside Palo Alto

Oakland

Los Gatos

Loma 

Prieta Peak

Ben Lomond

Mountain

Mount Madonna

B

Figure 1.—Continued
E10 EARTHQUAKE OCCURRENCE

Figure 2.—Epicentral area in the Loma Prieta region, Calif. (fig. 1). Star, epicenter of October 17, 1989, main shock. A, Part of the San Francisco Bay block, showing locations of major and subsidiary fault blocks, fold axes, lines of cross sections S–2 through S–7 (see pls. 1, 2), 

exploratory wells discussed in text (numbered crossed circles), and distribution of focused surface deformation (stippled areas) from earthquake. 

B, Summit Road-Skyland Ridge area, showing locations of Pliocene and younger fold axes, areas of the Purisima Formation (dark shaded), 

faults in Summit Road-Skyland Ridge fault zone (red lines), principal trace of the San Andreas fault, the abandoned Wrights Railroad Tunnel 

(which was damaged in 1906 San Francisco earthquake), and distribution of landslides along the San Andreas fault. Circled numbers 1–4, 

exposures of principal trace of San Andreas fault discussed in text: 1, Spanish Ranch Road; 2, beneath landslide in tributary to Los Gatos 

Creek; 3, tributary to Los Gatos Creek adjacent to Holy City No.1 exploratory well, which penetrated the San Andreas fault at 1,108-m depth; 

4, Moody Gulch. Circled numbers 5–11, other localities discussed in text: 5, north portal of the Wrights Railroad Tunnel; 6, Moody Gulch oil 

field (cross hachured area); 7, Lexington fault zone; 8, trenched 1989 coseismic cracks and faults in the Summit Road-Skyland Ridge fault 

zone on California Highway 17 at Summit Road; 9, trenched 1989 and older coseismic-deformation features associated with sag locality in the 

Summit Road-Skyland Ridge fault zone south of Summit Road; 10, coseismic cracks along Morrell Road that exhibited sinistral displacement 

in October, 1989 and April 1906; 11, 1989 coseismic cracks exhibiting dextral extension along the Sargent fault at Lake Elsman.

Fault—Dashed where approximately located, 

showing relative movement

Thrust Fault—Sawteeth on upper plate

Anticline—Showing crest line and direction of 

plunge

Syncline—Showing crest line and direction of 

plunge

Bodies of water

Areas of focused deformation during Loma Prieta earthquake. 

Includes deformation associated with the Summit Road-Skyland 

Ridge fault zone and compressional deformation associated with faults 

of the southwestern Santa Clara Valley thrust system

2.

S-3 Locations of structure sections shown in Plate 1

Exploratory well discussed in text: 1. Union Holy City No.1; 2. Union 

Teresa Hihn No.1; 3. Union Teresa Hihn No. 2; 4. Bean Hill Hihn No.1.

EXPLANATION

0

0 5

5 10 15 KILOMETERS

10 MILES

MONTE

VISTA

FAULT

BERROCAL FAULT

SAN

ANDREAS FAULT

BLACK

MOUNTAIN

BLOCK

LA

HONDA

BLOCK

BUTANO FAULT

LA

HONDA

BLOCK (GABBROIC)

Glenwood

Syncline

ZAYANTE - VERGELES

FAULT

BEN LOMOND

BLOCK

 D

SANTA

TERESA BLOCK

BERROCAL

FA ULT

LIME KILN FAULT

SOMBROSO

BLOCK

EL

? S-1

ALDERCROFT FAULT

S-2

SIERRA AZUL FAULT LOMA

PRIETA

THRUST

SIERRA

AZUL

BLOCK

SARGENT FAULT

U D

SARGENT FAULT

Laurel Anticline

SHANNON FAULT

S-5

BEN TROVATO MELANGE

BERROCA L FAUL T

S-3

S-4

LOMITA FAULT

S-6

SAN

ANDREAS FAULT

S-7

UVAS

 BLOCK

122°07'30'' 122°00'

121°52'30'' 121°45'00'' 121°37'30''

37°

15'

00''

37°

07'

30''

37°

00'

2.

3. 4.

1.

LEXINGTON

FAULT

LOS

CAPITANCILLOS

BLOCK

SILVER CREEK FAULT

COYOTE PEAK FAUL T

SODA SPRINGS FAULT

CASTRO FAULT

CARNADERO F

LITTLE ARTHUR CK FAULT

HOOKER GULCH FAULT

LOMA PRIETA

EARTHQUAKE

MAINSHOCK EPICENTER

OUTLINE OF AREA

OF FIGURE 2B

Lexington

Reservoir

Photo, figure 7C

Photo, figure 7E

Cupertino

Stevens

Creek Reservoir

Saratoga

New

Almaden

Summit Syncline

Uvas

 Reservoir

Calero

 Reservoir

U

Los

Gatos

A

is suggested by recent velocity and resistivity models that 

include a low-velocity, low-resistivity unit between these 

faults that extends to more than 8-km depth. This unit is 

interpreted to be largely Tertiary marine sedimentary rocks 

(Eberhart-Phillips and others, 1990), although highly fractured basement may produce similar low resistivity.

Deep exploratory wells between the San Andreas and Zayante-Vergeles faults (fig. 1) have failed to reach basement. 

One well (Union Teresa Hihn No. 2), located about 600 m 

northeast of the mapped surface trace of the Zayante fault, 

deviated southward across a buried strand of the Zayante 

fault to reach quartz monzonite at 1,047-m depth (Ross 
E11

Fault—Dashed where approximately located, 

showing relative movement

Anticline—Showing crest line and direction of 

plunge

Syncline—Showing crest line and direction of 

plunge

Landslide—Showing direction of movement

Wrights Railroad tunnel (abandoned)

Stream—Arrow showing direction of flow

EXPLANATION

D

U

SARGENT FAULT

SARGENT FAULT

Laurel

Anticline

Glenwood Syncline

ZAYANTE FAULT

ZAYANTE FAULT

Uvas Creek Soquel Creek

Corralitos C reek

Lau re l Creek

Laurel Creek

Burns Creek

Summit Road

Summit Road

122°00'00''

121°52'30''

37°07'30''

 D

U SAN

ANDREAS

FAULT

Lexington

Reservoir

7

4

6

3

8

9

2 5

10 11

1

 Lake Elsman

Mount

Umunhum

Loma Prieta Peak

 D

U

 D

U

SAN

ANDREAS

FAULT

HIGHWAY

17

HIGHWA Y17

Los

Gatos

Cre ek

Old San Jo se Rd.

Highland Way

Brow n's Creek

Mount

Thayer

Skyland Ridge

Santa Rosalia

Mountain

121°45'00''

0 4 KILOMETERS

0

1

1 1 2 3 MILES

1 2 3

0.5

0.5

B

Figure 2.—Continued

and Brabb, 1973; McLaughlin and others, 1988b, sheet 2). 

Union Hihn No. 1 well (cross sec. S–6, pl. 1) to the east 

failed to reach basement at a total depth of 2,361 m, where 

it bottomed in lower Miocene rocks (Ross and Brabb, 1973, 

p. 277; McLaughlin and others, 1988b). If the subsurface 

thickness of the pre-lower Miocene sedimentary section here 

approximates the 3,260 m of composite Eocene and Oligocene section exposed in the central Santa Cruz Mountains 

to the northwest, then the top of basement in this area would 

be at a depth of as much as 5.4 km. If Paleocene strata are 

also preserved here beneath Eocene rocks, the thickness of 

this Tertiary section could be almost 6 km.

An elongate body of distinctive gabbro crops out southwest 

of the San Andreas fault at Logan (fig. 2). Surface exposures 

of these crystalline basement rocks extend southeastward for 

10 km to the vicinity of San Juan Bautista (fig. 1A), where 

the exposures are largely of undated, banded anorthositic 

gabbro containing conspicuous secondary fibrous amphibole 

(Ross, 1970; Dibblee and others, 1979). On the southwest, 

the anorthositic gabbro is locally faulted against green marine 

mudstone beds that yield foraminifers and nannofossils diagnostic of lower-bathyal to abyssal depths and an early Eocene 

age (Bukry and others, 1975; Kristin McDougall, written 

commun., 1991). At Logan, the gabbroic unit includes hornblende-quartz gabbro with a granular texture and a Pb-U age of 

161 to 165 m.y. (Johnson and O’Neil, 1988; James and others, 

1994). An initial 87Sr/86Sr ratio of 0.7039 (table 1) suggests that 

the hornblende-quartz gabbro has oceanic affinities.

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E12 EARTHQUAKE OCCURRENCE

Figure 3.—Summary of tectonic history across the San Andreas fault in the Loma Prieta region, Calif. (fig. 1), comparing differences in 

timing and amount and style of tectonism in major fault blocks. CRO, Coast Range ophiolite.

Past geophysical and geologic studies have postulated that 

the basement between the Zayante-Vergeles and San Andreas 

faults consists of granitic and metamorphic rocks (Clark and 

Rietman, 1973), Jurassic gabbro (Hanna and others, 1972), 

Franciscan rocks (Ross and Brabb, 1973; Coppersmith, 

1990), or mafic oceanic rocks (Ross, 1984; Clark and others, 

1991). Recent gravity and magnetic models (see Jachens 

and Griscom, this chapter), however, indicate that these 

basement rocks have a high magnetization and high density, 

consistent with northwestward subsurface continuation of 

the Jurassic hornblende-quartz gabbro of Logan and the 

undated anorthositic gabbro near San Juan Bautista.

The northern part of the La Honda block is underlain by 

rocks that are less magnetic and less dense than the gabbro 

of Logan (see Jachens and Griscom, this chapter) and that 

presumably represent a southeastward extension of the granitic rocks exposed to the northwest at Montara Mountain 

southwest of the Pilarcitos fault (fig. 1B). The north magnetic 

boundary of the Logan gabbroic basement with this granitic 

basement (fig. 1B) is about 5 km north and subparallel to the 

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E13

surface trace of the Butano fault (see Jachens and Griscom, 

this chapter). Thus, the gabbro of Logan and the anorthositic 

gabbro near San Juan Bautista (fig. 1A) appear to compose 

an elongate block of mafic oceanic rocks within the Salinia 

terrane that is juxtaposed with granitic and metamorphic 

rocks of the terrane. This block of mafic rocks of the Salinia 

terrane is bounded on the northeast by the San Andreas fault, 

on the southwest by the Zayante-Vergeles fault, on the west 

by the San Gregorio fault, and on the north by a presumed 

buried fault (fig. 1B). On the basis of gravity and magnetic 

models (see Jachens and Griscom, this chapter), the gabbroic 

basement appears to be truncated at depth by a flat boundary 

interpreted to be a low-angle fault. The rocks beneath this 

low-angle fault are less magnetic than the gabbro and are 

also believed to be crystalline basement (see Jachens and 

Griscom, this chapter).

TERTIARY SECTION SOUTHWEST OF

THE SAN ANDREAS FAULT

Cretaceous strata are unknown in the Santa Cruz Mountains 

between the San Andreas fault and the San Gregorio fault to the 

southwest (fig. 1). The Tertiary section exposed there consists 

largely of marine clastic sedimentary rocks ranging in age from 

Paleocene to Pliocene, with a composite thickness of more than 

10 km (Clark and Brabb, 1978). This section is divisible into 

four major sequences separated by regional unconformities 

(see pl. 1). The exposed basal beds of each sequence locally 

rest on crystalline rocks of the Salinia terrane.

POSSIBLE BURIED PALEOCENE SEQUENCE

Paleocene rocks are exposed west of the study area (fig. 1B) 

near Ben Lomond Mountain, where erosional remnants of 

the Locatelli Formation that rest on the crystalline basement 

consist of basal, shallow-marine arkosic sandstone overlain 

by bathyal to abyssal siltstone. As much as 275 m thick, the 

sequence is of late Paleocene (Ynezian) age (Clark, 1981). 

Paleocene strata may occur in the subsurface above the basement of the La Honda block (see pl. 1), but coeval rocks have 

not been penetrated by exploratory wells in this area, and the 

oldest sedimentary rocks exposed at the surface northeast of 

the Zayante fault are Eocene.

EOCENE TO LOWER MIOCENE SEQUENCE

The Eocene to lower Miocene sequence consists of the 

Butano Sandstone, the Twobar Shale and Rices Mudstone 

Age, method of Source 87Sr/

86

87 Sr Sr/86Sr Lithologic unit dating, and mineral of data ratio dated, where known age date sources

(Hornblende-quartz) gabbro of

Logan.

161–165 Ma, Pb-U,

zircon.

1 0.7039 (initial) 1, 2

Hornblende-quartz gabbro of

Eagle Rest Peak.

159–177 Ma, Pb-U,

zircon.

1 .7034–0.7039

(initial)

1, 2

Hornblende-quartz gabbro of

Gold Hill.

144 Ma, Pb-U, zircon 1 .7038 (initial) 1, 2, 5

Coast Range ophiolite of Sierra

Azul block (hornblende-albite

pegmatite in dikes and sills).

144–168 Ma, Pb-U,

zircon.

4 .70242–0.70404

(initial)

3

Franciscan Complex, Permanente

terrane (basalt and tuff).

136–120 Ma, 87Sr/

86Sr

initial, whole rock.

7 .70394–0.70998

(0.70365 initial)

7

Diabase of Laurel Creek and

Morrell Cutoff Road.

Unknown, possibly

Oligocene to Miocene.

This

report

.70439–0.70596

(0.70437 initial)

6

Tertiary basalt flows of the

Felton quadrangle.

23.7 Ma, K-Ar, plagioclase.

8 .70360 (initial) 6

Diabase and gabbro of Highland

Way and Diablo Gulch.

Unknown, possibly

Mesozoic.

This

report

.70403–0.70692

(initial)

3

Table 1.—Summary of 87Sr/ 86Sr and radiometric age data for igneous rocks of the Loma Prieta region, Calif., and related 

areas along the San Andreas fault

[Data sources: 1, James and others (1994), with age of 161 Ma assumed in calculating 87Sr/86Sr (initial) ratio; 2, Kistler and others (1973); 3, R.W. Kistler 

(written commun., 1990), with age of 160 Ma assumed in calculating 87Sr/86Sr (initial) ratio; 4, J.L. Wooden (written commun., 1992), with true age interpreted 

as ≥168 Ma; 5, R.W. Kistler (written commun., 1990), with age of 143 Ma assumed in calculating 87Sr/86Sr (initial) ratio; 6, R.W. Kistler (written commun., 

1990), with age of 20 Ma assumed in calculating 87Sr/86Sr (initial) ratio; 7, R.W. Kistler(written commun., 1990), with age derived from Rb/Sr isochron; 8, 

Turner (1970)] 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E14 EARTHQUAKE OCCURRENCE

Members of the San Lorenzo Formation, the Zayante Sandstone, the Vaqueros Sandstone, and the Lambert Shale. 

Basaltic flows locally are included in the Vaqueros Sandstone. 

Only the lower part of this sedimentary sequence is exposed 

in the Ben Lomond block south of the Zayante fault (fig. 1), 

where as much as 1.1 km of conglomeratic beds of the Butano 

Sandstone rests nonconformably on basement rocks of the 

Salinia terrane (see pl. 1).

Northeast of the Zayante fault, in the La Honda block, the 

base of the sequence is not exposed, but as much as 3.1 km 

of the section is preserved along the Glenwood syncline and 

Laurel anticline south of the Butano fault (see pl. 1). These 

rocks also crop out along the Summit syncline north of the 

Butano fault and in fault blocks southwest of and near the 

San Andreas fault.

The Butano Sandstone is interpreted to have been deposited 

at lower-bathyal to abyssal depths by northward-flowing currents as the midfan and interchannel facies of a submarine fan 

(Nilsen, 1984). Its arkosic composition suggests derivation 

from a granitic source terrane that lay to the south, probably 

in the vicinity of present-day Monterey Bay (fig. 1A).

The overlying mudstone beds of the Twobar Shale Member 

of the San Lorenzo Formation record undisturbed marine 

deposition at lower-bathyal depths in an anoxic basin. Benthic foraminifers are assignable to the Narizian Stage of the 

middle to late Eocene (Smith, 1971), and planktic foraminifers 

indicate open-ocean connections.

The Rices Mudstone Member of the San Lorenzo Formation 

locally records shallower marine conditions. Along Soquel 

Creek in the Laurel quadrangle and to the east in the Loma 

Prieta quadrangle, the lower part of this unit is massive, finegrained glauconitic sandstone, containing locally abundant 

Pitar and other mollusks characteristic of inner-neritic depths 

and a Refugian (late Eocene) age (Smith, 1971; Clark and 

others, 1989). The upper part of the Rices Mudstone Member 

is nodular mudstone that yields benthic foraminifers diagnostic of upper-middle-bathyal (500–1,500 m) depth and an early 

Zemorrian (Oligocene) age (McDougall, 1989, 1991).

The Vaqueros Sandstone is restricted to the La Honda block, 

where it was laid down in environments ranging from shallow-marine shelf to deep-sea fan (Stanley, 1990). In this area, 

benthic foraminifers from mudstone interbeds in the lower 

part of the formation are diagnostic of bathyal depths and an 

early Zemorrian (Oligocene) age (Clark and others, 1989). To 

the south, along Hinckley Creek in the Laurel quadrangle, the 

upper 92 m of the Vaqueros Sandstone contains thick Dosinia

and Ostrea biostromes indicative of shallow-marine conditions (Clark and others, 1989). Locally, along the Glenwood 

syncline, the Vaqueros Sandstone encloses conglomeratic 

interbeds of the Zayante Sandstone, a nonmarine facies that 

was deposited in alluvial fans along the Zayante fault at the 

southern margin of the La Honda Basin (see pls. 1, 2).

The limited areal extent of the Zayante Sandstone, together 

with the absence of the marine San Lorenzo Formation and 

Vaqueros Sandstone southwest of the Zayante fault, was 

interpreted by Clark and Rietman (1973) to be the result 

of regional uplift of the Ben Lomond block along this fault 

during Oligocene time (fig. 3). Arkosic debris shed from the 

emergent Salinia terrane to the south resulted in nonmarine 

deposition of the Zayante Sandstone along the fault and a rapid 

northeastward facies change within the Vaqueros Sandstone 

from marine shelf to deep-sea fan.

Mafic igneous rocks that occur in isolated exposures and in 

the subsurface of this part of the La Honda block are probably 

correlative with the more voluminous Oligocene and (or) 

Miocene Mindego Basalt of the central La Honda block and 

with nearby basalt flows in the Felton quadrangle to the west. 

These basalt flows, as much as 60 m thick (Clark, 1981), have 

been dated by K-Ar methods at 23.7±0.7 Ma (late Oligocene 

or early Miocene; Turner, 1970; age recalculated by Fox and 

others, 1985, using currently accepted International Union 

of Geological Sciences constants). Their low 87Sr/86Sr ratio 

of 0.70360 (R.W. Kistler, written commun., 1991) suggests 

an oceanic magmatic source (table 1). The Mindego Basalt 

was believed by Stanley (1990) to have been erupted in an 

extensional tectonic setting, concurrent with displacement 

along the Zayante fault, which served to localize the less 

voluminous volcanic activity in this area.

In the study area (fig. 1), basaltic flows and (or) sills are 

locally exposed along the Glenwood syncline, along California 

Highway 17 north of the Zayante fault in the Laurel quadrangle, and near the west edge of the Los Gatos quadrangle 

(see pl. 1). Basaltic rocks that are also reported as intrusive into 

the San Lorenzo Formation both in surface exposures and in 

the subsurface of the Moody Gulch oil field in the Los Gatos 

quadrangle (Krueger, 1943) were penetrated by Union Hihn 

No. 1 well (cross sec. S–6, pl. 1; fig. 2A) in the Loma Prieta 

quadrangle (McLaughlin and others, 1988b).

Altered fine- to coarse-grained diabase and gabbro also 

occur in fault slivers along Laurel Creek in the Laurel quadrangle and along Morrell Cutoff Road in the Los Gatos quadrangle. The diabase and gabbro are of oceanic composition 

(

87Sr/86Sr ratio, 0.70437–0.70596) and geochemically (TiO2

content versus FeO/MgO ratio, and Ti versus Cr contents) are 

nearly identical to the Mindego Basalt.

In the Loma Prieta quadrangle, fine- to medium-grained, 

locally titaniferous diabase, basalt, and gabbro (cross sec. S–7, 

pl. 1) occur in fault-bounded slivers along the northeast side 

of the San Andreas fault. The geochemistry of these rocks and 

a 87Sr/86Sr ratio of 0.707 (table 1) suggest an oceanic-ridge or 

offridge formational setting. Their position just to the northeast of the principal trace of the San Andreas fault, however, 

suggests that these exposures may be more closely associated 

with the Mesozoic basement northeast of the fault rather than 

with the Mindego Basalt to the southwest.

The Lambert Shale is conformable above the Vaqueros 

Sandstone along the Glenwood syncline, where from 460 to as 

much as 550 m of the shale crops out in the Laurel quadrangle 

(Clark and others, 1989). Throughout this area, the upper part 

of the Lambert Shale has been removed by erosion.
E15

A

B

? ?

?

?

?

? ? ?

?

?

?

?

? ? ?

?

? ? ? ? ?

? ?

?

?

?

? ?

L ua C l er

keer

GsoL sot a keer C

SEA LEVEL

5

1015

W

E

N

S

W E

NS

W E

OBSERVED

EXPECTED(from focalmechanism)

DEXTRAL SL PI (EXPECTED)

O( BSERVED)

S NI SI TRAL SL PI

OBSERVED

EXTENSION

DIRECTIONS

N30 ̊ E

N10 ̊ E

N18 ̊ W

N60 ̊ W

N18 ̊ W

N10 ̊ W

N60 ̊ W

RANGE IN

CRACK

ORIENTATIONS

INFERRED

ORIENTATION

OF

COMPRESSION

VECTORS BEN LOMOND BLOCK LA HONDA BLOCK

S NI SI TRAL SL PI

O( BSERVED)

DEXTRAL SL PI

E( XPECTED)

NET MAXIMUM

NET MAXIMUM EXTENSION

COMPRESSION

RANGE IN

CRACK

ORIENTAT OI NS

DEPTH (KM) TRANSITIONAL BOUNDARY

CRYSTALLINE ROCKS

(GRANITIC TO MAFIC PLUTONS,

+ METAMORPHIC WALL ROCKS,

± UNDERTHRUST FRANCISCAN

ROCKS

"OCEANIC"

BASEMENT

(7.0 TO 7.1 KM/SEC)

(~6.0 TO 6.4 KM/SEC)

(~6.4 TO 7.0 KM/SEC)

CRYSTALLINE

ROCKS

MAINSHOCK (7.0 TO 7.1 KM/SEC)

)

GABBRO OF LOGAN

FRANCISCAN

COMPLEX

(~5.0 TO 6.0

KM/SEC)

COAST

RANGE

OPHIOLITE

OPHIOLITIC ±

FRANCISCAN ROCKS

(~6.0 TO 6.4 KM/SEC)

SANANDREAS

FAULT

1.8

2 Y=1.

X=0.5

1.3

2.3

70 ̊ 

RELATIVE MOTION VECTORS,

OCT. 17, 1989 FAULT RUPTURE

LET UPLIFT = Y

Y = 1.3 (SIN 70 ̊)

Y = 1.2M

LET SHORTENING = X

X = 1.3 (COSIN 70 ̊)

X = 0.5M

NET SLIP TO N10 ̊ 

W

RAKE = 140 ̊

FAULT FROM OLSENAND HILL, 1993

CRYSTALLINE METAMORPHIC PLUTONIC BASEMENT

(~6.4 TO 7.0KM/SEC)

SEA LEVEL

5

1015

SEA LEVEL5

1015

S80 ̊ W

N

S

ZAYANTE

FAULT

BUTANO

FAULT

SUMMIT ROADSKYLAND RIDGE

FAULT ZONE

ALDERCROFTFAULT

Figure 4.—Surface deformation and crustal structure in the Summit Road-Skyland Ridge area (fig. 2

B). 

A, Rose diagrams comparing observed and expected horizontal surface-deformation 

fields during 1989 Loma Prieta earthquake. 

B, Block diagram showing inferred crustal structure across the San Andreas fault and possible relation to primary and secondary slip during 

1989 Loma Prieta earthquake. Red echelon faults at surface and shallow subsurface are fissures in the Summit Road-Skyland Ridge fault zone. Loma Prieta rupture is shown in red at depth, extending upward from main shock to base of the gabbro of Logan. Deep configuration of the San Andreas fault is partly inferred from Olson and Hill (1993). Crustal structure to about 10-km depth is partly inferred from Jachens and Griscom (this chapter), and below about 10-km depth is highly speculative and inferred from indicated seismic velocities (Fuis and Mooney, 1990; Rufus Catchings, oral commun., 1993; see Brocher and others, this chapter).

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E16 EARTHQUAKE OCCURRENCE

A decrease in the influx of coarse detritus, together with 

subsidence of the La Honda Basin, resulted in deposition of 

organic and phosphatic mudstone beds of the Lambert Shale 

at upper-middle-bathyal (500–1,500 m) depths (see pl. 1), 

largely under anoxic conditions (Clark, 1981). Arkosic sandstone interbeds in the upper part of the Lambert Shale were 

probably derived from the Salinia terrane to the south and laid 

down by occasional turbidity flows. Benthic foraminifers in 

the Lambert Shale in the Laurel quadrangle are diagnostic of 

the Saucesian to early Relizian Stages of the early Miocene 

(Clark, 1981; McDougall, 1991).

LOWER TO MIDDLE MIOCENE SEQUENCE

The lower to middle Miocene sequence consists of a widely 

transgressive basal sandstone unit, the Lompico Sandstone, 

and an overlying organic mudstone unit, the Monterey Formation. This marine sedimentary sequence is best preserved in 

the Ben Lomond block and locally occurs in the subsurface 

of the La Honda block (see pls. 1, 2; fig. 3).

Surface exposures of this sequence in the study area (see 

pls. 1, 2) are restricted to the western part of the Laurel 

quadrangle, where 192 m of section is locally exposed south 

of the Zayante fault. There, the Lompico Sandstone rests 

unconformably on the Butano Sandstone and includes a 

coquina bed containing mollusks, barnacles, and foraminifers 

diagnostic of near-shore, shallow-marine conditions and an 

early to middle Miocene (Relizian to Luisian) age, according to Clark and others (1989). The conformably overlying 

micaceous siltstone and subsiliceous organic mudstone beds 

of the Monterey Formation yield pelecypods and benthic 

foraminifers characteristic of neritic depths and a Luisian 

(middle Miocene) age.

Northeast of the Zayante fault, where the sequence is not 

exposed, 17 m of siliceous strata of the Monterey Formation 

and 70 m of the Lompico Sandstone was penetrated by Bean 

Hill Ltd. Hihn No. 1 well (cross sec. S–7, pl. 1) in the Loma 

Prieta quadrangle (McLaughlin and others, 1988b). Thus, the 

lower to middle Miocene sequence probably extended over 

all of this area but is now restricted to erosional remnants 

beneath the upper Miocene-Pliocene unconformity.

UPPER MIOCENE TO PLIOCENE SEQUENCE

The upper Miocene to Pliocene sequence consists of a 

transgressive, shallow-marine sandstone unit, the Santa Margarita Sandstone; a deeper-marine siliceous organic mudstone 

unit, the Santa Cruz Mudstone; and a shallow-marine unit, 

the Purisima Formation (see pl. 1). The Purisima Formation records late Miocene and Pliocene continuation of the 

marine transgression that began with deposition of the Santa 

Margarita Sandstone during earlier late Miocene time. The 

transgression spread more widely during the early Pliocene, 

probably recording the major global sea-level rise at about 

5–4.2 Ma (Haq and others, 1987).

The Purisima Formation includes distinctive lithic sandstone 

containing abundant andesitic detritus and a heavy-mineral 

assemblage characterized by hypersthene, augite, basaltic 

hornblende, and green and brown hornblende (Clark, 1981). 

Locally, it includes rhyolitic tuff interbeds. One tuff bed that 

crops out in Forest of Nisene Marks State Park in the eastern 

part of the Laurel quadrangle is interpreted to be of air-fall 

origin and is tentatively correlated with the Nomlaki Tuff 

Member (of the Tehama, Tuscan, and Laguna Formations), 

which was erupted from the southern Cascade Range and 

K-Ar dated at 3.4 Ma (late Pliocene; A.M. Sarna-Wojcicki, 

written commun., 1978).

The sequence rests unconformably on all older units and 

is widely distributed in both the Ben Lomond and La Honda 

blocks. In the Laurel quadrangle, as much as 1,000 m of this 

sequence is preserved along the Glenwood syncline. In this 

quadrangle, the Santa Margarita Sandstone and Santa Cruz 

Mudstone pinch out to the east, and the Purisima Formation 

is more widespread (see pl. 1). The Purisima Formation also 

is preserved in fault slivers along the San Andreas fault zone. 

Throughout the study area (see pl. 1), the upper part of the 

Purisima Formation has been removed by erosion, and locally 

this formation is overlain unconformably by Quaternary 

deposits.

STRATIGRAPHY NORTHEAST OF

THE SAN ANDREAS FAULT

The area northeast of the San Andreas fault is underlain by a 

composite Mesozoic basement consisting of Franciscan rocks, 

Coast Range ophiolite, and Great Valley sequence (see pl. 

1). Seismic reflection and refraction data for the Loma Prieta 

region have been interpreted as indicating that these formations occur from the surface to 8- to 12-km depth (Fuis and 

Mooney, 1990). Newer data to the north in the San Francisco 

Bay region suggest that rocks with seismic velocities compatible with this those of this composite basement may extend 

to 16- to 18-km depth (Hole and others, 1993). To the east 

beneath the Diablo Range, the composite Franciscan, Great 

Valley, and ophiolitic basement is interpreted as extending to 

about 12-km depth, gradually shallowing and thinning farther 

to the east beneath the western San Joaquin Valley (Wentworth and others, 1984; Fuis and Mooney, 1990; Wentworth 

and Zoback, 1990). Below this eastward-thinning basement, 

rocks with seismic velocities like those of the metamorphic 

and plutonic rocks exposed in the Sierra Nevada foothills 

slope southwest to the San Andreas fault (Fuis and Mooney, 

1990).

The composite basement consisting of Franciscan Complex, 

Coast Range ophiolite, and Great Valley sequence is broken at 

the surface into several structural blocks by major late Cenozoic faults. Displacements across these faults have resulted in 
E17

different stratigraphic sections within each block (see pl. 1). 

A major change in Oligocene to Miocene stratigraphy occurs 

across the Berrocal and Sargent faults, which define part of 

the northeast boundaries of the El Sombroso block of Bailey 

and Everhart (1964) and the Sierra Azul block of this paper 

(see pl. 2; fig. 2). The Sierra Azul block is bounded along its 

northwest side by parts of the Aldercroft, Soda Spring, and 

Sierra Azul faults (fig. 2). The study area northeast of the San 

Andreas fault is divided into several additional fault blocks, 

modified after the study by Bailey and Everhart.

FRANCISCAN COMPLEX

(JURASSIC AND CRETACEOUS)

Northeast of the San Andreas fault and southwest of the 

Silver Creek fault (see pl. 1), Franciscan rocks compose most 

of the exposed structural basement. The Franciscan Complex 

consists predominantly of what is termed its Central belt, 

which is largely a melange of igneous and pelagic oceanic 

rocks and clastic sedimentary rocks derived from two highly 

disrupted tectonostratigraphic terranes. Pelagic rocks of the 

Bald Mountain-El Sombroso terrane (McLaughlin and others, 

1991a, b) include radiolarian chert of Early Jurassic to Early 

Cretaceous age that formed at equatorial latitudes (Hagstrum 

and Murchey, 1991). This chert was deposited on an Early 

Jurassic or older basaltic basement formed in near-ridgeseamount to midocean-ridge-basalt (MORB)-like settings 

(McLaughlin and others, 1991a, c). The younger Permanente 

terrane includes both radiolarian chert and pelagic limestone 

of mid-Cretaceous (Aptian) to Late Cretaceous (Turonian) age 

formed at equatorial latitudes and deposited on basaltic rocks 

whose geochemistry and 87Sr/86Sr ratios (table 1) suggest an 

intraplate, seamount or oceanic-plateau formational setting 

(Sliter and others, 1991; McLaughlin and others, 1991a, c).

The depositional age of the sandstone component of the Bald 

Mountain-El Sombroso terrane is Early to mid-Cretaceous, 

on the basis of the ages of radiolarians in clasts of chert in 

conglomerate and on the inferred 90-Ma age of high-pressure 

metamorphism of some weakly foliated metaclastic rocks of 

the terrane (McLaughlin and others, 1991a). The depositional 

age of terrigenous rocks associated with the Permanente 

terrane is poorly constrained, except that locally they depositionally overlie and so are considered to postdate the pelagic 

limestone. Thus, the metaclastic rocks of the Permanente 

terrane are Late Cretaceous (Turonian) or younger.

Tectonic disruption of the Bald Mountain-El Sombroso 

and Permanente terranes and their mixing into melange of 

the Central belt along the continental margin followed earlier 

metamorphism and uplift of older Franciscan rocks in the 

melange. The later disruption and mixing of the Bald Mountain-El Sombroso and Permanente terranes at the California 

margin occurred sometime between the Late Cretaceous 

(Campanian) and early Eocene, when the rocks were accreted, 

unroofed, and overlain by early Tertiary deposits (fig. 3).

The Franciscan Complex also includes minor blocks of 

garnet-bearing glaucophane schist, eclogite, amphibolite, 

and other sedimentary and volcanic rocks metamorphosed 

to high-pressure-metamorphic-mineral assemblages, as well 

as local thin, lenticular bodies of serpentinite. These rocks 

generally are associated with argillic zones considered to be 

the matrix of the melange. The serpentinite is at least partly 

derived from nearby rocks of the Coast Range ophiolite, 

and some unmetamorphosed terrigenous sedimentary rocks 

sheared into the melange are derived from Upper Jurassic to 

Lower Cretaceous parts of the Great Valley sequence.

Rocks of these Franciscan terranes are imbricated by several 

faults at low angles both to bedding and to lithologic contacts. 

These faults appear to repeat and thicken the section, and are 

folded. Thus, they are probably deformed thrusts, although no 

detailed kinematic data are available. Locally, outcrop-scale 

asymmetric isoclinal and recumbent folds and axial-planar 

cleavage are associated with the thickened section. These 

structures probably reflect contraction at the California margin 

during the Late Cretaceous and early Tertiary accretion of the 

Franciscan Complex (fig. 3).

SLATE AND PHYLLITE OF LOMA PRIETA PEAK

(JURASSIC?)

A thin unit of slate and phyllite (Jlpp, pl. 1) is exposed at 

one locality beneath Loma Prieta peak (cross sec. S–6, pl. 

1; figs. 1B, 2A) in the sole of the Loma Prieta thrust. The 

coarse-grained and pelitic clastic components of this unit 

are flattened, and the rocks exhibit a strong penetrative, slaty 

cleavage. The metapelite also is highly siliceous and locally 

tuffaceous. Flattened lenses of metaconglomerate include 

mostly clasts of metachert, fine-grained porphyritic andesitic 

to rhyolitic volcanic rocks, and granitic rocks. The unit is 

distinguished from rocks of the Franciscan Complex by its 

lithology and geochemistry, as well as by the absence of any 

high-pressure/temperature-metamorphic-mineral phases, 

which almost always occur in Franciscan rocks that have 

undergone equivalent flattening and cleavage development 

(McLaughlin and others, 1991b). On the basis of vitrinite 

reflectance and petrography, the slate and phyllite of Loma 

Prieta peak has, instead, undergone much higher temperature 

and lower pressure metamorphism to greenschist facies, at 

temperatures of ≥310°C (Underwood and others, 1991). 

These rocks are similar in composition and metamorphism 

to the Jurassic Mariposa Slate of the western Sierra Nevada 

metamorphic belt, the Galice Formation of the western 

Klamath Mountains, and the Santa Monica Slate of the Santa 

Monica Mountains in the Transverse Ranges. The slate and 

phyllite of Loma Prieta peak also is similar to rocks that 

occur along the Coast Range fault in the northern Sacramento 

Valley that have been correlated with the Galice Formation 

(Jayko and Blake, 1986). We interpret these rocks to have 

been faulted upward from a deep (≥10 km) sub-Franciscan 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E18 EARTHQUAKE OCCURRENCE

crystalline basement that may underlie structurally shallower 

Franciscan rocks and Coast Range ophiolite (figs. 3, 4), as 

implied in interpretations of deep crustal structure by Fuis 

and Mooney (1990).

COAST RANGE OPHIOLITE (JURASSIC)

Rocks of the Jurassic Coast Range ophiolite, the oceanic 

basement on which Jurassic and Cretaceous clastic marine 

strata of the Santa Teresa and Sierra Azul blocks were deposited, structurally overlie the Franciscan Complex and the 

slate and phyllite of Loma Prieta peak. Gravity and magnetic 

modeling and seismic reflection and refraction data further 

suggest that similar ophiolitic rocks may underlie the Franciscan Complex below about 8-km depth. Field exposures 

in the Loma Prieta region suggest that the ophiolite section 

ranges in thickness from about 400 to 1,700 m (see pl. 1). 

A composite section, on the basis of maximum thicknesses 

for each ophiolite unit (see pl. 1), permits a thickness of as 

much as 2,135 m, although nowhere is this much continuous 

section exposed at the surface. Interpreted aeromagnetic data 

(see Jachens and Griscom, this chapter) suggest that buried 

parts of the ophiolite section could be as much as 3 km thick 

in the eastern part of the study area (fig. 1), and more than 2 

km thick at depth elsewhere.

Although we recognize that magnetic models imply an 

ophiolite section as much as 2.0 to 3.0 km thick at depth, the 

thickness depicted in our structure sections is based on surface 

exposures (see pl. 1). Aeromagnetic modeling, however, also 

indicates the ophiolite section to be complexly segmented by 

wedge-shaped bodies of nonmagnetic rocks interpreted to be 

part of the Franciscan Complex (see Jachens and Griscom, 

this chapter).

From bottom to top, the Coast Range ophiolite is composed of varying thicknesses of cumulate ultramafic rocks, 

cumulate gabbro, diabase to diorite dikes and sills, basaltic 

to andesitic flows, andesitic to quartz-keratophyric tuff with 

local thin intercalations of recrystallized radiolarian chert, 

and coarse-grained quartz-keratophyre pyroclastic breccia 

(McLaughlin and others, 1991a). Initial 87Sr/86Sr ratios for 

the igneous section of the ophiolite range from a primitive 

value of 0.70242 in the cumulate ultramafic rocks, upward to a 

higher value of 0.70404 at the top of the more highly evolved 

volcanic section (table 1).

A zircon separate from dikelets of hornblende-albite pegmatite in the noncumulate intrusive rocks yielded Pb-U ages 

that lie on concordia between 144 and 168 Ma (J.L. Wooden, 

written commun., 1992). Hornblende from the same sample 

yielded a discordant K-Ar age of 138.5 Ma (Nakata and others, 

1993). Two alteration events, one in the Cretaceous and one 

in the Miocene, appear to have completely or partially reset 

the K-Ar and Pb-U ages. Therefore, the upper concordia Pb-U 

age of 168 Ma is considered to be closest to the true cooling age for the intrusive section of the ophiolite. Elsewhere 

in California, Pb-U ages of zircons from plagiogranite in 

cumulate and intrusive sections of the Coast Range ophiolite 

range from 163 to 173 Ma (Hopson and others, 1981, 1991; 

Mattinson and Hopson, 1992).

On the basis of radiolarians contained in interbedded chert 

and tuff in the upper part of the volcanic sequence, the volcanic 

pelagic rocks are as young as Oxfordian (160–156 Ma) and 

possibly as old as Bajocian (181–175 Ma) in their lower part 

(Bonnie Murchey, written commun., 1990). Radiolarians from 

the volcanic pelagic section of the Coast Range ophiolite at 

other localities in the Coast Ranges are dated at early Oxfordian to late Tithonian (163–145 Ma; Hull and others, 1992).

The exposed ophiolitic sequence of the El Sombroso and 

Sierra Azul blocks has been attenuated along high- to lowangle detachment faults that displace structurally higher parts 

of the ophiolite downward onto structurally lower parts of 

the section (see pl. 1). These attenuation faults, which may 

be related to Mesozoic and Tertiary unroofing of the Coast 

Ranges, have had a complex history of reactivation.

GREAT VALLEY SEQUENCE

(UPPER JURASSIC TO UPPER CRETACEOUS)

The Great Valley sequence was defined by Bailey and others 

(1964) as the Upper Jurassic through Cretaceous sedimentary 

strata and associated outlier strata originally deposited in the 

marine depositional basin of the Great Valley of California. In 

the northern and central Coast Ranges, most upper Mesozoic 

marine clastic sequences overlying ophiolitic rocks, which 

also structurally overlie the Franciscan Complex, have been 

assigned to the Great Valley sequence. Some outlier sections 

have no established depositional connection to the Great 

Valley, including the section described herein, and are not 

similar lithologically to any known coeval marine sections 

in the Sacramento or San Joaquin parts of the Great Valley 

sequence or to any coeval rocks of the Great Valley sequence in 

the Diablo Range. For example, the framework-grain compositions of sandstone and the clast compositions of conglomerate 

in Upper Cretaceous rocks of the Sierra Azul block reflect 

a mafic to intermediate-composition volcanic and plutonic 

provenance devoid of chert, with only minor quartzite. Coeval 

rocks along the southwest side of the Great Valley have a more 

quartzose provenance, with conglomerate that contains fewer 

mafic igneous clasts and significantly more chert and quartzite (Seiders and Blome, 1988; Seiders and Cox, 1991). For 

this reason, the term “Great Valley sequence” is used herein 

to indicate only that rocks of this age and structural setting 

occur in the Sierra Azul, Santa Teresa, and Los Capitancillos 

blocks. We do not imply original structural or depositional 

continuity with rocks directly to the east in the Diablo Range 

or Great Valley. Any original connection with rocks to the east 

has been disrupted by Miocene to Holocene large strike-slip 

displacement across the Hayward-Calaveras and San Andreas 

faults, and by additional pre-San Andreas (Campanian to early 
E19

Eocene) strike-slip displacement along the Sur-Nacimiento 

and ancestral Calaveras fault systems (McLaughlin and others, 

1990). In addition, the upper Mesozoic clastic section of 

the Sierra Azul block, as with the underlying ophiolite and 

overlying lower Tertiary rocks, is highly attenuated along 

faults at low angles to bedding. Most of the Lower and Upper 

Cretaceous section of the Great Valley sequence of the Sacramento-San Joaquin Valley and Diablo Range is missing in 

the Loma Prieta region (see pls. 1, 2; fig. 1).

UPPER JURASSIC AND LOWER CRETACEOUS ROCKS

A thin unit of dark-green to black argillite overlies the Coast 

Range ophiolite discontinuously along attenuation faults in the 

Loma Prieta region (see pls. 1, 2; fig. 1). The rocks commonly 

are penetratively deformed, but numerous megafossils indicate 

a Late Jurassic (Tithonian) and Early Cretaceous (Berriasian? 

to Valanginian) age (McLaughlin and others, 1988b, 1991a, 

b). The rocks contain local, lenticular to nodular, phosphatized 

or silicified carbonate concretions. The basal beds of this unit 

are tuffaceous. In one place north of the Calero Reservoir, a 

thin lens of fossiliferous mudstone and sandstone composed of 

serpentinite detritus overlies a slice of Coast Range ophiolite 

(see pl. 1). This argillite unit in the lower part of the Great 

Valley sequence is generally no more than 100 to 200 m thick 

and is bounded above by another low-angle attenuation fault 

and locally by younger southwest-side-up reverse faults.

UPPER CRETACEOUS ROCKS

Overlying the lower part of the Great Valley sequence above 

a sheared contact, which is interpreted as another low-angle 

normal fault, is a unit of sandstone, shale, and conglomerate, 

as much as 1,615 m thick in the Sierra Azul block and as much 

as 335 m thick in the Santa Teresa block. From bottom to top 

(see pl. 1), the sedimentary section consists of conglomerate, 

lithic arkosic wacke, thinly interbedded arkosic wacke and 

shale turbidites, and massive shale and fine-grained sandstone 

containing local carbonate concretions.

Locally, the concretionary massive shale and sandstone 

near the top of the unit are overlain by a few meters of thinly 

interbedded, graded feldspathic sandstone and interbedded 

shale. Strata assigned to this unit in the Santa Teresa block 

are thin-bedded concretionary shale and sandstone that are 

undated but lithologically resemble the upper part of Upper 

Cretaceous rocks of the Sierra Azul block (see pl. 1). In the 

Sierra Azul block, megafossils occur in conglomerate of the 

lower part of the unit and in massive concretionary shale and 

sandstone of the upper part. These megafossils indicate a Late 

Cretaceous (Campanian) age, and the sedimentology of the 

enclosing strata suggests their deposition in a submarine-fan 

or slope setting (Elder, 1990). The conglomerate and thickbedded arkosic sandstone suggest a proximal or channeled, 

upper-midfan or slope setting, whereas overlying thin-bedded 

turbidites may have been deposited in deeper water in an 

interchannel, midslope region. Megafossils in the conglomerate were apparently transported downslope from a near-shore 

high-energy habitat; fossils in overlying concretionary shale 

and sandstone also were moved downslope, but from a more 

offshore shelf setting (Elder, 1990). Limited paleocurrent 

data suggest transport to the west-southwest. Tangentially 

crossbedded sandstone, locally interbedded with argillite in 

the uppermost beds of the succession, may record marine 

regression after deposition of the underlying transgressive 

succession (see pl. 1).

The conglomerate clast suite of the Sierra Azul block is 

generally well rounded and consists predominantly of cobbles 

to boulders of mafic and intermediate-composition porphyritic volcanic rocks, porphyritic silicic volcanic rocks, mafic 

to granitic plutonic rocks, and various minor sedimentary 

and metasedimentary rocks (McLaughlin and others, 1988b, 

1991a, b). The rocks of this unit are extensively deformed by 

fracturing, shearing, and folding; and in the Sierra Azul block, 

the unit is extensively laumontitized.

EOCENE SEQUENCE

In the Sierra Azul and Santa Teresa blocks, a sequence 

of shale and sandstone, mudstone, and bioclastic limestone 

overlies the Jurassic and Cretaceous Great Valley sequence 

and the Franciscan Complex. This younger sequence, possibly 

as much as 1,110 m thick, may be structurally repeated in the 

Sierra Azul block. In the Santa Teresa block, the sequence is 

as much as 670 m thick (see pl. 1). The base of the sequence 

is a sheared low-angle contact interpreted to be an attenuation 

fault, which places Eocene rocks over Great Valley sequence, 

Coast Range ophiolite, or Franciscan rocks.

Locally at the faulted base of the Eocene sequence is a 

unit that varies in composition from bioclastic limestone in 

the Santa Teresa block to glauconitic calcareous lithic grit or 

conglomerate in the Sierra Azul block. The unit commonly 

contains clastic material derived from the underlying Coast 

Range ophiolite and Franciscan Complex, indicating significant unroofing of this area of the Coast Ranges before the 

early Eocene. In addition, the unit locally contains shallowmarine mollusks, algal debris, and discocyclinid foraminifers. 

In the Santa Teresa and Sierra Azul blocks, shallow-water 

foraminifers are mixed with deeper-water fauna, indicating 

downslope displacement from a near-shore to bathyal setting. 

These foraminifers indicate that the basal unit is early Eocene 

(planktic zone upper P6b or P7; McDougall, 1989, 1991).

These basal beds are overlain by as much as 200 m of distinctive olive-green and red-and-green-mottled foraminiferal 

mudstone containing an early Eocene (planktic zones P8, P9) 

foraminiferal fauna indicative of a lower-bathyal to abyssal 

depositional setting (McDougall, 1989, 1991). The mottled 

mudstone and sandstone of Mount Chual (see pl. 1) is recogSTRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E20 EARTHQUAKE OCCURRENCE

nized over much of the Eocene California margin northeast 

and southwest of the San Andreas fault (McLaughlin and 

others, 1990, 1991b).

Overlying the Mount Chual unit is as much as 915 m of 

interbedded quartzofeldspathic sandstone and micaceous 

brown carbonaceous mudstone (see pl. 1), here called the 

sandstone and shale of Loma Chiquita Ridge, which typically 

are deeply weathered and locally are hydrothermally altered. 

Sparse foraminiferal faunas and rare, poorly preserved megafossils from these strata suggest an early Eocene (planktic zone 

P6b) to late Eocene age for the rocks, which overlaps the age of 

underlying Eocene strata. Paleobathymetry of the foraminifers 

indicates deposition in a middle- to lower-bathyal-slope setting. In most places, these rocks are in contact with underlying 

formations along attenuation faults, which locally cut out the 

underlying rocks at a low angle to bedding.

LOWER EOCENE TO LOWER MIOCENE 

SEQUENCE

The lower Eocene to lower Miocene sequence is recognized 

only along the southwest side of the Sierra Azul block adjacent 

to the San Andreas fault. This sequence is bounded to the northeast by a poorly defined fault that places the rocks over partly 

coeval Eocene strata of the Sierra Azul block. This northeast 

contact is interpreted to be a reverse or thrust fault (fig. 3) 

because of its southwestward dip and the apparent repetition 

of the lower Eocene part of the sequence across the boundary 

(cross secs. S–6, S–7, pl. 1; see pl. 2).

The lower Eocene to lower Miocene sequence consists of as 

much as 1,100 m of hard carbonaceous to cherty shale with 

minor interbeds of locally glauconitic quartzofeldspathic sandstone, a unit that is here referred to as the shale and sandstone 

of Highland Way. Microfossils from carbonaceous shale in the 

lower part of the sequence indicate an early Eocene (planktic 

zones P8 and P9 and nannofossil zone CP 11) age (McDougall, 

1989), indicating that the shale and sandstone of Highland Way 

is coeval with, but lithologically dissimilar from, the two structurally underlying Eocene marine units. Foraminifers from the 

upper beds of this sequence indicate a late Oligocene to early 

Miocene (Zemorrian to Saucesian) age, whereas the intervening 

strata are poorly dated. The benthic foraminifers suggest deposition at bathyal or greater depths, consistent with a midslope to 

lower-slope setting. These strata also are partly coeval with the 

Temblor Sandstone north of the Sierra Azul block (see pl. 1) 

but were deposited in a much deeper marine setting.

PRE-PLIOCENE SEQUENCE NORTHEAST OF THE 

EL SOMBROSO AND SIERRA AZUL BLOCKS

Two formations ranging in age from Oligocene to Miocene 

are exposed only in the area northeast of or along the Berrocal 

and Sargent faults.

TEMBLOR SANDSTONE

(OLIGOCENE TO MIDDLE MIOCENE)

The Temblor Sandstone, about 200 m thick, consists of 

thick-bedded to massive subfeldspathic to arkosic sandstone, 

minor mudstone, unsorted conglomerate, and dacitic tuff. 

The conglomerate occurs locally at or near the base of the 

formation, where it nonconformably overlies the Franciscan 

Complex and Coast Range ophiolite and contains locally 

derived angular detritus from the uplifted Franciscan Complex and Coast Range ophiolite and from the Eocene sequence 

of the Sierra Azul block (McLaughlin and others, 1991a, b). 

Reeflike accumulations of fragmented oysters, barnacles, 

bryozoans, and pectens occur locally within the basal conglomerate, suggesting a littoral to inner-shelf depositional 

setting and an Oligocene or early Miocene age (see pl. 1). 

In the New Almaden area (see pl. 1; fig. 2A), the upper beds 

of the Temblor are tuffaceous and grade upward into dacite 

tuff and tuff breccia. Map relations depicted by Bailey and 

Everhart (1964) imply that some of the dacite could be intrusive and postdate the volcaniclastic rocks. The tuff has been 

dated at about 15.6 Ma (D.H. Sorg and P.C. Russell, written 

commun., 1990). Locally, this tuff is hydrothermally altered 

and contains mercury mineralization (Bailey and Everhart, 

1964).

MONTEREY SHALE (MIDDLE MIOCENE)

The Monterey Shale consists of yellowish-white-weathering, 

hard, brown diatomaceous mudstone, porcelanite, and cherty 

porcelanite, as much as 424 m thick. The rocks conformably 

overlie the Temblor Sandstone and contain nannofossils and 

benthic foraminifers diagnostic of middle Miocene (nannofossil zone CN3 or CN4, Luisian) age (David Bukry and 

Kristin McDougall, written commun., 1989) and indicative of 

deposition at 500- to 1,500-m depth. The Monterey Shale is 

lithologically distinct from, and represents a shallower marine 

setting than, the lower Eocene to lower Miocene sequence in 

the Sierra Azul block (see pl. 2; fig. 3).

QUATERNARY AND TERTIARY NONMARINE 

DEPOSITS NORTHEAST OF THE 

SAN ANDREAS FAULT

SANTA CLARA FORMATION

(PLIOCENE AND PLEISTOCENE)

The Santa Clara Formation consists predominantly of 

fluvial coarse-grained gravel and fanglomerate; silty sandstone; silty, locally lithic carbonaceous mudstone; and 

laminated lacustrine mudstone. Cummings (1968) recognized 

eight lithofacies in the Santa Clara Formation, three of which 

crop out in the Loma Prieta region (see pl. 1). The Los Gatos 
E21

lithofacies (and its right-laterally offset equivalent, the Corte 

Madera lithofacies near Portola Valley), which is exposed in 

the Los Gatos area (figs. 1B, 2A), consists of fluvial gravel 

and fanglomerate containing clasts reworked from Upper 

Cretaceous conglomerate of the Sierra Azul block. The Stevens Creek lithofacies consists of gravel and fanglomerate, 

with a clast suite of graywacke and minor basalt, diabase, red 

and green metachert, and vein quartz, derived from adjacent 

uplands of the Franciscan Complex south of the Santa Clara 

Valley (see pl. 1). The Arastradero lithofacies consists of 

moderately well sorted, sandy, crossbedded pebble gravel 

and minor claystone and mudstone, deposited in braidedstream and flood-plain/overbank settings (Vanderhurst and 

others, 1982).

The Arastradero lithofacies, which is the stratigraphically 

lowest part of the Santa Clara Formation, includes laminated 

lacustrine mudstone adjacent to the San Andreas fault west of 

the Lexington Reservoir (see pl. 1). In the Saratoga area (fig. 

2A), the Arastradero lithofacies locally contains fossil-plant 

debris, freshwater mollusks, and vertebrate-fossil remains, 

including Felis (Cheetah) studeri of Blancan (1.9–4.8 Ma) 

age (Adam and others, 1982; Repenning and others, 1990). 

Beds overlying the Blancan part of the Santa Clara Formation are poorly dated because of their coarse, unfossiliferous, 

nonmarine characteristics, but the overlying section is dated 

at Pleistocene, partly on the basis of a 250-ka age assigned 

to the older fluvial-terrace and alluvial-fan deposits of Santa 

Clara Valley that unconformably overlie the Santa Clara Formation (McLaughlin and others, 1991a), and the presence 

of a 500–600-ka ash horizon (Rockland ash bed of SarnaWojcicki and others, 1991; Sarna-Wojcicki, 2000) in marine 

intertongues in the Santa Clara Formation near Woodside, 

Calif. (fig. 1B). If additional time is considered for deposition 

of the clastic section overlying the Rockland ash bed, then the 

age of Santa Clara Formation is approximately constrained to 

the range 0.3–4.8 Ma. Santa Clara strata adjacent to the San 

Andreas fault and along the frontal thrust-fault system bounding the east side of the Santa Cruz Mountains are steeply tilted 

and deformed into open, high-amplitude, northwest-trending 

folds. In the lowland areas and subsurface northeast of the 

mountain front, gravel mapped as the Santa Clara Formation 

shows little, if any, discordance to overlying, undeformed to 

mildly tilted fluvial deposits.

ALLUVIAL AND LANDSLIDE DEPOSITS

(PLEISTOCENE AND HOLOCENE)

Extensive alluvial-fan and stream deposits, younger than 

about 250 ka (McLaughlin and others, 1992), overlie the 

Santa Clara Formation with marked discordance. These 

deposits generally consist of poorly to moderately well sorted 

pebble to boulder gravel, sand, silt, and mudstone, deposited 

along ancient to modern stream channels and flood plains. 

The maximum age assigned to these deposits is based partly 

on the estimated 0.3-Ma age for the deformed upper part of 

the unconformably underlying Santa Clara Formation, and 

partly on the degree of weathering of the terrace materials 

(McLaughlin and others, 1992). In addition, oxygen-isotopic 

ratios plotted against time and correlated with eustatic sea-level 

fluctuations indicate global warming and eustatic highstands 

at 250 and 120 ka (Sarna-Wojcicki and others, 1991). This 

relation further constrains the age of the post-Santa Clara 

alluvial deposits. These highstands are the most likely times 

for cutting of the two highest and oldest post-Santa Clara 

fluvial terraces. Thus, the high terraces cut into the Santa 

Clara Formation along Los Gatos Creek between Los Gatos 

and the Lexington Reservoir are dated at about 250 ka. The 

upper part of the Pleistocene alluvial-fan- and stream-deposit 

section around the present San Francisco Bay margin locally 

contains vertebrate (Rancholabrean) fossils and radiocarbondated wood and plant materials dated at about 22 ka. Overlying Pleistocene to Holocene alluvial deposits have yielded 

radiocarbon ages of about 14 ka and younger (Wright, 1971; 

Lajoie and Helley, 1975).

Numerous large composite landslide deposits, ranging from 

rotational blocks of coherent rocks with intact internal stratigraphy to incoherent, unsorted, rubbly debris flows, occur 

in the Loma Prieta region (fig. 1), particularly along the San 

Andreas fault between the Lexington Reservoir and the divide 

between Soquel and Corralitos Creeks. Surface deformation 

in this area during the 1989 Loma Prieta earthquake probably 

resulted partly from reactivation of some of these landslides 

(Cotton and others, 1990; Spittler and Harp, 1990; Ponti and 

Wells, 1991).

STRUCTURE SOUTHWEST OF

THE SAN ANDREAS FAULT

The exposed structure southwest of the San Andreas fault 

zone consists of a series of nearly parallel faults and folds that 

trend more westerly than the San Andreas fault zone itself 

(see pl. 1; fig. 2). The more important folds are the Glenwood 

syncline and the Laurel anticline, between the Zayante and 

Butano faults, and the Summit syncline to the north (see pl. 

1; fig. 2). These folds apparently developed concurrently with 

displacements on the major faults. Major faults displacing the 

Tertiary section are the Zayante-Vergeles fault, which also 

forms the boundary between two major structural blocks, 

and the Butano fault. Hundreds of meters to more than 1 km 

south of the bedrock San Andreas fault in the Los Gatos and 

Laurel quadrangles are the faults in the Summit Road-Skyland 

Ridge area, which were reactivated during the 1989 Loma 

Prieta earthquake. The fault zone there consists of a series 

of geomorphically youthful echelon faults that trend more 

northerly or subparallel to the throughgoing San Andreas fault 

zone and locally appear to offset the Butano fault. The faults 

of the Summit Road-Skyland Ridge area are discussed below 

in the section entitled “San Andreas Fault Zone.”

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E22 EARTHQUAKE OCCURRENCE

ZAYANTE-VERGELES AND

RELATED BURIED FAULTS

The Zayante fault zone is the most important structure 

present between the San Andreas and San Gregorio faults. 

In the Santa Cruz Mountains (see pl. 1; fig. 1A), it is mapped 

for 55 km from west of Ben Lomond Mountain, where it 

is covered by upper Miocene strata of the Santa Margarita 

Sandstone and Santa Cruz Mudstone, southeastward to 

the vicinity of Watsonville, where locally it is covered by 

Quaternary deposits (Brabb, 1989). A gravity investigation 

(Clark and Rietman, 1973) demonstrated the continuity of 

this fault southeastward beneath Quaternary cover to the 

Vergeles fault in the northern Gabilan Range. The Vergeles 

fault, in turn, continues southeastward to the San Andreas 

fault south of San Juan Bautista. The combined ZayanteVergeles fault is 82 km long (figs. 1, 2).

Although the throughgoing strand of the Zayante fault 

is confined to a narrow zone, branching lineaments locally 

occupy a zone as much as 1 to 2 km wide (Clark, 1981, pl. 

2; McLaughlin and others, 1988b, sheet 2). The relatively 

straight course of this fault across canyons and ridges in 

the study area (see pl. 1) suggests that the fault plane dips 

steeply. Shear zones along the Vergeles fault south of San 

Juan Bautista dip 65°–85° S.(Ross, 1984, fig. 10). Gravity 

and aeromagnetic data (see Jachens and Griscom, this chapter) in the Santa Cruz Mountains suggest a southwestward 

dip to a depth of a few kilometers that steepens to near vertical at greater depth (cross sec. S–3, pl. 1; fig. 4).

South of the Zayante-Vergeles fault, the basement of the 

Salinia terrane is exposed or relatively shallow. Between 

the Zayante and San Andreas faults, the basement is not 

exposed in this part of the Santa Cruz Mountains, but 

gravity and magnetic data (see Jachens and Griscom, this 

chapter) suggest that the buried basement is most probably 

the northwestward continuation of the hornblende-quartz 

gabbro that crops out at Logan (figs. 1, 4). Thus, the ZayanteVergeles fault appears to be a major structure in the Loma 

Prieta region juxtaposing continental and oceanic basements, 

truncated by the San Andreas fault at its southeast end (see 

pl. 1; figs. 1, 4).

Ross (1970, 1984) postulated that the gabbroic rocks at 

Logan may be right-laterally offset some 300 to 350 km by 

the San Andreas fault from petrographically and chemically 

similar gabbroic rocks at Eagle Rest Peak (fig. 1A) in the 

San Emigdio Mountains at the southern tail of the Sierra 

Nevada. A south-dipping fault of the Pleito thrust zone (fig. 

1A) bounds the north side of the gabbroic rocks of Eagle 

Rest Peak; this fault could be equivalent to, or may at depth 

root into, a fault equivalent to the buried fault at the base of 

the gabbro of Logan (fig. 4). Although the gabbroic rocks of 

Eagle Rest Peak are covered to the south by Eocene marine 

strata, a southern segment of the Pleito thrust zone that 

cuts Miocene strata is aligned with the south contact of the 

gabbro. Faults of the Pastoria-Rand thrust system also are 

aligned with the south contact of the gabbroic rocks. Thus, 

the Pleito or Pastoria-Rand fault zones may correspond to 

the offset continuation of an ancestral Zayante-Vergeles fault 

(Ross, 1970, 1984). To juxtapose the Zayante-Vergeles fault 

with the Pleito thrust requires the restoration of about 350 

km of right slip on the San Andreas fault.

The juxtaposition of ≥91-Ma continental basement of 

the Ben Lomond block with 165–161-Ma oceanic gabbroic 

basement of the La Honda block along the Zayante-Vergeles 

fault (see pl. 1; fig. 1) seems to require major up-to-the-south 

dip-slip fault displacement. Geobarometric and fission-track 

data suggest that the basement of the Salinia terrane in the 

Ben Lomond block, southwest of the fault, was uplifted 

more than 13 km to within 5 km of the surface between 

90 and 61 Ma (Hansen and others, 1991; Arrowsmith and 

others, 1992; Bürgmann and others, 1994). South of the 

Zayante-Vergeles fault and outside the study area (fig. 1A), 

depositional remnants of the shallow- to deep-marine upper 

Paleocene Locatelli Formation overlie the basement of the 

Salinia terrane in the Ben Lomond block (Clark, 1981). The 

Locatelli Formation may also be present in the subsurface 

north of the Zayante-Vergeles fault. The basements of the 

La Honda and Ben Lomond blocks are also overlain north 

and south of the Zayante-Vergeles fault by about the same 

thicknesses of bathyal-marine strata of the lower to upper 

Eocene Butano Sandstone (see pl. 2; Clark, 1981). These data 

all suggest that initial juxtaposition of the La Honda and Ben 

Lomond blocks occurred between about 90 and 61 Ma.

Juxtaposition of the gabbro of Logan with Ben Lomond 

basement across the Zayante-Vergeles fault corresponds to 

a period of general uplift and unroofing in the Coast Ranges 

that seems to have begun southwest of the San Andreas 

fault (and thus farther south) as early as 91–60 Ma. From 

the latest Cretaceous (60 Ma) to the early Tertiary (52 Ma), 

oblique dextral convergence is believed to have occurred 

in northern California along the interface between the continental margin and rocks composing the Central belt of 

the Franciscan Complex (McLaughlin and others, 1988a). 

Some of this oblique strike-slip movement may have been 

accommodated within the Salinia terrane along the ZayanteVergeles-Pastoria fault.

As much as 3 to 4 km of vertical displacement along the 

Zayante fault during Oligocene time (see pl. 2; fig. 3) resulted 

in emergence of the Salinia terrane to the south, in deposition 

of the terrestrial Zayante Sandstone along and north of the 

fault, and in restriction of marine conditions to a northeastsloping La Honda Basin (Clark, 1981). Mafic volcanism 

was localized along this fault during late Oligocene to early 

Miocene time. Allen (1946) reported a remarkably similar 

history along the Vergeles segment of the fault in the San 

Juan Bautista area to the southeast.

In the Laurel quadrangle, the Purisima Formation and the 

lower part of the Butano Sandstone are juxtaposed along the 

Zayante fault, indicating at least 3.7 km of dip-slip separation, with the north side relatively downthrown. Model stud-
E23

ies of the gravity field (Clark and Rietman, 1973) suggest 

that the basement is displaced as much as 2.7 km by this 

fault in the southern Santa Cruz Mountains near Corralitos, 

but in their modeling these workers assumed a density of 

2.67 g/cm3 for the Salinian basement between the Zayante 

and San Andreas faults. More recent modeling (see Jachens 

and Griscom, this chapter), which takes into account the 

greater density and magnetic contrasts across the fault and 

assumes these contrasts to define the Zayante-Vergeles fault 

at depth, suggests as much as 4 to 6 km of vertical dip-slip 

displacement (up on the southwest), with at least 6 km of 

horizontal shortening.

Most displacement along the Zayante fault at the surface is 

in rocks older than middle Miocene. The fine clastic deposition 

recorded by the Lambert Shale north of the fault suggests that 

by early Miocene (Saucesian) time, the uplifted Ben Lomond 

block was no longer supplying coarse detritus and thus had 

been reduced to a lowland. Subsequent transgression by middle 

Miocene seas led to deposition of the Lompico SandstoneMonterey Formation sequence. Although most displacement 

appears to have occurred in Oligocene and earlier time, late 

Pliocene regional transpression reactivated the Zayante fault, 

because the Purisima Formation is offset along the Zayante 

fault in the southern part of the study area (see pl. 1). Late 

Pleistocene and possible Holocene displacement has been 

suggested for strands of the Zayante fault near Watsonville 

(Hall and others, 1974; Coppersmith, 1990), but in the Santa 

Cruz Mountains to the north, the Zayante fault lacks youthful 

geomorphic expression.

Aftershocks associated with the southwest-dipping Zayante 

fault at 7- to 10-km depth may indicate reactivation of the deep 

part of this fault during the earthquake. This deep slip did not 

propagate to the surface in the Santa Cruz Mountains (fig. 4).

Widespread basaltic volcanic rocks of Oligocene to early 

Miocene age in the Coast Ranges (including the Mindego Basalt 

and associated volcanic rocks of the La Honda Basin) have been 

cited as evidence for regional crustal extension resulting from 

a component of divergent motion (transtension) between the 

Pacific and North American plates during early development of 

the San Andreas transform fault (Clark and Rietman, 1973, pl. 

1; Stanley, 1985, 1990; Atwater, 1989; Tennyson, 1989). The 

stratigraphic record across the Zayante-Vergeles fault indeed 

indicates major up-on-the-southwest dip-slip movement during 

the Oligocene (see pl. 1; fig. 3). The volcanism in the La Honda 

Basin also seems to be best explained within the context of 

horizontal extension above thin crust. The present reverse-slip 

geometry of the Zayante-Vergeles fault both at the surface 

(see pl. 1; fig. 4) and at depth (see Jachens and Griscom, this 

chapter) is, however, difficult to relate to an extensional setting. 

Interpretation of the Zayante-Vergeles fault as a transtensional 

structure seemingly requires either that (1) the fault dips northeast at some depth or that it dipped northeast before about 

18 Ma and has since rotated to its present orientation, (2) an 

earlier normal fault has been overprinted by the present vertical 

to southwest-dipping Zayante-Vergeles thrust fault, or (3) the 

transtensional displacement was accommodated by faults not 

situated along the Zayante-Vergeles fault zone. This problem 

is presently unresolved and needs further study.

Compressional-deformation events were imposed on La 

Honda Basin fill deposits along the Zayante-Vergeles fault 

zone from 17 to 16 Ma (early Relizian), from 14 to 10 Ma (late 

Luisian to early Mohnian; Clark, 1981), and during the late 

Oligocene (fig. 3). Pliocene compression across the ZayanteVergeles fault may have resulted from a change in plate-motion 

vectors between the Pacific and North American plates that 

caused conspicuous transpression across the entire transform 

margin (Atwater, 1989; Harbert and Cox, 1989).

BUTANO FAULT

The Butano fault is poorly exposed in the southwestern part 

of the Los Gatos quadrangle, where it trends southeast into 

the San Andreas fault (see pl. 1; figs. 1B, 2A). In this area, the 

Butano fault lacks geomorphic expression and is mapped with 

difficulty, on the basis of the juxtaposition of different parts of 

the lower Eocene to lower Miocene sequence through which 

it passes. Where the Butano Sandstone is faulted against the 

Vaqueros Sandstone, dip separation is as much as 1.3 km, with 

the north side relatively downthrown. The Butano fault has 

been mapped northwestward from the study area as a series 

of bifurcating faults for a distance of as much as 37 km. In the 

northern Santa Cruz Mountains, these faults are reported to 

dip steeply southwest and to have a combined dip separation 

of as much as 1.8 km (Cummings and others, 1962).

Although the pre-Eocene slip history of the Butano fault 

is uncertain, sedimentary units of the lower Eocene to lower 

Miocene sequence that are almost identical across this 

fault have been offset by post-early Miocene reverse slip. 

Displacement may have occurred during the two Miocene 

compressional deformations that may have activated the Zayante-Vergeles fault, and during late Pliocene transpression. A 

lower member of the Purisima Formation in the northern Santa 

Cruz Mountains is the youngest stratigraphic unit cut by the 

Butano fault; in that area, stream terraces of Pleistocene age 

appear to extend across the fault zone without interruption 

(Brabb and Olson, 1986). The Butano fault does not appear to 

have moved during the earthquake, nor is there evidence for 

Holocene activity on the Butano fault in the study area (fig. 

1). Its eastern trace appears to be offset by the geomorphically 

youthful echelon faults in the Summit Road-Skyland Ridge 

area (see next section).

SAN ANDREAS FAULT ZONE

The San Andreas fault is the principal member of a transform-fault system that has accommodated as much as 460 

km of right-lateral motion between the Pacific and North 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
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Figure 5.—Schematic map of northwestern California and adjacent offshore area, showing partitioning of Miocene and younger dextral slip on faults o
E25

American plates since the early Miocene as it extended 

northward along the California margin (Huffman, 1972; 

Ross and others, 1973; Mathews, 1976; Fox and others, 1985; 

Stanley, 1987; Irwin, 1990). In northern and central California, the San Andreas fault system consists of numerous fault 

segments (fig. 5; table 2) that have accommodated different 

components of the total displacement at different times. Estimates of the amount and timing of slip that is partitioned to 

various segments of the fault system vary, partly owing to a 

poor understanding of the slip histories of the San Gregorio 

and Pilarcitos faults and faults in the eastern San Francisco 

Bay region.

Estimates of slip on the San Gregorio fault range from 115 

to 170 km (Graham and Dickinson, 1978; Clark and others, 

1984; Ross, 1984; Jachens and others, 1998). This slip is added 

to that in the San Andreas-Pilarcitos fault zone in the offshore 

area north of San Francisco (fig. 5; table 2).

At the intersection of the San Andreas and Calaveras faults 

near Hollister (fig. 1A), part of the remaining 310 km of slip 

on the San Andreas fault system is partitioned to the Calaveras 

fault (McLaughlin and others, 1990; Jones and Curtis, 1991). 

About 160 to 170 km of slip is partitioned to faults in the 

eastern San Francisco Bay region from the Calaveras fault, 

on the basis of offsets of the trend in ages of volcanic rocks 

and hydrothermal mineralization and displacement of the 

distinctive Permanente terrane of the Franciscan Complex 

across the Calaveras fault (McLaughlin and others, 1996). 

Additional evidence from gravity and aeromagnetic data 

refines the estimated offset on the southern section of the 

Calaveras fault (R.C. Jachens, oral commun., 2000) to about 

174 km. With this refinement, the remaining 136 km of right 

slip on the San Andreas fault system is apparently taken up by 

the San Andreas and Pilarcitos faults and (or) other unidentified faults on the San Francisco peninsula. Dextral offset of 

the Corte Madera lithofacies of the Santa Clara Formation 

across the Peninsular segment of the San Andreas fault from 

its source area in the Sierra Azul block (Cummings, 1968), 

together with displacement of the axis of a linear magnetic low 

associated with the Permanente terrane (Robert C. Jachens, 

oral commun., 2000), indicates that only 22 to 30 km of the 

remaining slip is attributable to the Peninsular segment of 

the San Andreas fault system. Thus, about 106 to 114 km of 

Miocene to Pliocene right slip remains to be accommodated 

by the Pilarcitos fault.

Most right-lateral slip across the Peninsular segment of 

the San Andreas fault appears to have occurred since 3 Ma. 

Modern right-lateral-strain rates on this segment based on 

geodetic measurements beginning before the earthquake 

(Thatcher and Lisowski, 1987; Thatcher, 1990) range from 5 

to 12 mm/yr. The slip rate across the entire transform margin 

at this latitude is about 30 mm/yr, apparently within the range 

indicated by geologic offsets (Brown, 1990).

Total Duration Slip rate Fault member displacement of faulting (mm/yr) (km) (Ma)

San Andreas fault south of intersection

with the east-bay fault system.

310 0–18 17

East-bay fault system-------------------------- 174 0–12 14

Peninsular segment of the San Andreas

fault.

22–30 0–3 10

Northern section of the San Andreas

fault system (Pilarcitos fault+San

Gregorio fault+Peninsular segment of

the San Andreas fault).

251–306 ~13

(0–3 and 8–18)

~19–22

(but 14–17 if slip

occurred between

18 Ma and present

Total, San Andreas fault system

(includes the San Gregorio fault, the

Pilarcitos fault, the Peninsular segment

of the San Andreas fault, and the east

bay fault system)

425–480 0–18 24–27

Total, Pilarcitos fault -------------------------- 106–114 ~13

(18–5)

~8–9

(but 5.8–6.3 if slip

occurred between

18 Ma and present

San Gregorio fault ----------------------------- 115–170 0–12 10–13

Table 2.—Miocene and younger offsets and slip rates for faults of the San Andreas fault system

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E26 EARTHQUAKE OCCURRENCE

SAN ANDREAS FAULT IN

THE LOMA PRIETA REGION

The San Andreas fault zone trends northwesterly through 

the study area (fig. 1) as a zone of deformation, ranging in 

width from about 610 m in the southeast to as much as 1,220 

m to the northwest (see pl. 1). The surface geology, as well 

as gravity and magnetic data (see Jachens and Griscom, this 

chapter), suggests that the fault zone is vertical or dips steeply 

southwest in the northwestern part of the study area. In the 

Loma Prieta quadrangle to the southeast, the upper part of the 

fault zone locally is considered to have a northeastward dip, 

on the basis of gravity and magnetic modeling (see Jachens 

and Griscom, this chapter), but the mapped fault zone follows 

a straight, apparently steep course at the surface (see pl. 1).

The San Andreas fault zone in the Loma Prieta region consists of two components. First, we recognize a “principal” or 

“main bedrock” component of the fault zone, on the basis of 

major differences in Pliocene and older stratigraphy across 

the fault. Thus, the principal trace of the fault zone corresponds to the long-term fault across which as much as 136 

km of right-lateral displacement is accommodated. Second, 

we recognize faults of the San Andreas fault zone that exhibit 

little evidence of large strike-slip offset but do exhibit strong 

evidence of repeated Quaternary surface displacements, on 

the basis of youthful geomorphic features (for example, 

aligned closed depressions and sag ponds, aligned narrow 

linear valleys, scarps, captured drainages).

Northwestward along the San Andreas fault zone from the 

southeast border of the study area (see pl. 1; fig. 2B) to the 

drainage divide between Corralitos and Soquel Creeks, youthful geomorphic features correspond closely to the principal 

trace of the San Andreas fault.

From the divide between Corralitos and Soquel Creeks 

northwestward to the boundary of the Los Gatos quadrangle, 

however, the San Andreas fault zone is obscured by extensive 

landslide deposits. In this area, the principal trace of the San 

Andreas fault forms the northeast boundary of the fault zone 

(see pl. 1; fig. 2B), and geomorphically youthful fault traces 

are mapped noticeably southwest of the principal fault trace 

(Sarna-Wojcicki and others, 1975; U.S. Geological Survey 

staff, 1989). The active, geomorphically youthful fault traces 

appear as photolineaments that cross the landslides and commonly align with structural and lithologic discordances in the 

Tertiary sedimentary section of the La Honda block exposed 

beneath the landslides.

From the divide between Corralitos and Soquel Creeks 

to the west boundary of the Los Gatos quadrangle, the 

principal trace of the San Andreas fault is exposed at only a 

few localities. Along Spanish Ranch Road in the northeast 

corner of the Laurel quadrangle, the fault is located to within 

15 m, where it juxtaposes Pliocene marine sandstone of 

the Purisima Formation on the southwest against lower to 

upper Eocene sandstone and shale sequences of the Mount 

Chual and Loma Chiquita Ridge areas to the northeast, and 

is draped by perched Quaternary alluvium (see pl. 1). The 

principal trace of the San Andreas fault is also exposed at 

several other isolated localities in tributary canyons south 

of Los Gatos Creek (see pl. 1; fig. 2B). At the southeast end 

of these localities (see pl. 1; fig. 2B), an 8-m-wide zone of 

gouge, including sheared serpentinite, separates the Rices(?) 

Mudstone Member of the San Lorenzo Formation on the 

southeast from Upper Cretaceous sandstone and shale of the 

Sierra Azul block to the northeast. In the two tributaries to 

the northwest (one adjacent to the community of Holy City 

and the other in Moody Gulch), the principal trace of the San 

Andreas fault separates Oligocene marine strata of the Vaqueros Sandstone on the southwest from lower Eocene sandstone 

and shale of the Sierra Azul block to the northeast.

At Holy City, Los Nietos (Union Oil) Holy City No. 1 well, 

located about 91 m southwest of the mapped principal surface 

trace of the San Andreas fault (see pl. 1; fig. 2B), apparently 

penetrated the fault at 1,108-m depth, indicating a dip of 84° 

SW. (cross sec. S–2, pl. 1).

No convincing evidence exists for primary coseismic movement along the mapped principal surface trace of the San 

Andreas fault from the drainage divide between Corralitos 

and Soquel Creeks to California Highway 17 during the 1989 

Loma Prieta earthquake, nor clear evidence that the main 

bedrock trace of the fault ruptured the surface in this area 

during the 1906 San Francisco earthquake. Although 1.4 m 

of right-lateral offset was recorded in the Wrights Railroad 

Tunnel (Lawson, 1908; Sarna-Wojcicki and others, 1975; 

Prentice and Ponti, 1994), this displacement was along a fault 

strand 150 to 300 m southwest and parallel to the principal 

trace of the San Andreas fault (see pl. 1; fig. 2B). Most of 

the surface deformation in 1989 and, possibly, 1906 was in 

the La Honda block, along secondary faults in the Summit 

Road-Skyland Ridge area, 1 to 2 km southwest of the principal trace of the San Andreas fault.

FAULTS IN THE

SUMMIT ROAD-SKYLAND RIDGE AREA

Geomorphically youthful faults in the Summit Road-Skyland Ridge area form a 6-km-wide zone trending N. 18°–65° 

W. Most of these faults trend N. 30°–40° W., subparallel to the 

principal trace of the San Andreas fault, and are more northerly 

than the Butano fault (trend, N. 65°–70° W.) and the axes of 

the Laurel anticline and Summit syncline. Most of the faults 

in this zone range from 300 to 800 m in length. One fault that 

cuts Morrill (Morrell) Road extends 3 km southeastward of 

Summit Road; another fault to the west of the Morrill Road 

fault appears to be the northwestward extension of a more 

nearly continuous fault that trends southeast into the Loma 

Prieta quadrangle (see pl. 1; fig. 2B).

These faults are mapped on the basis of structural discordances in the Butano Sandstone, San Lorenzo Formation, 

and Vaqueros Sandstone and, locally, of the juxtaposition of 
E27

different rock units (see pl. 1). Several of the faults, which 

were trenched in the vicinity of C.T. English School and Loma 

Prieta School along Summit Road, were found to juxtapose 

colluvium and Paleogene bedrock units of the La Honda 

block (Sarna-Wojcicki and others, 1975; Johnson and Associates, 1989). Most of the faults display youthful geomorphic 

features, such as closed depressions, swales, scarps, and 

vegetation lineaments along their extent (Sarna-Wojcicki and 

others, 1975), that have resulted from repeated displacements. 

Radiocarbon dating of offset soil horizons along one of these 

faults that trends beneath the Loma Prieta School grounds 

shows that some of the displacements are Holocene (Johnson 

and Associates, 1989).

Several faults along the west side of the Summit Road-Skyland Ridge area trend more westerly and parallel the strike 

of the Vaqueros Sandstone on the south limb of the Summit 

syncline (see pl. 1; fig. 2B). One of these faults is well exposed 

on the east side of California Highway 17, about 120 m south 

of Summit Road (see pl. 1; fig. 2B), and another was trenched 

south of Summit Road (670 m south of the NE. cor. sec. 21, T. 

9 S., R. 1 W.; see pl. 1). Both are bedding-plane normal faults, 

with the north sides relatively downthrown, and both faults 

slipped during the 1989 Loma Prieta earthquake. Multiple 

episodes of normal slip have occurred on these faults during 

the late Quaternary; the California Highway 17 fault has 

undergone about 8 m of cumulative displacement during the 

past 20 k.y. (Cotton, 1990; Cotton and others, 1990).

EARTHQUAKE DEFORMATION IN

THE SUMMIT ROAD-SKYLAND RIDGE FAULT ZONE

During the earthquake, numerous linear coseismic cracks 

and fissures, some as much as 600 m long, formed along the 

Summit Road-Skyland Ridge fault zone (Sarna-Wojcicki and 

others, 1975; Clark and others, 1989; Johnson and Associates, 

1989; McLaughlin and others, 1991a). Many of the coseismic cracks underwent sinistral extension during the 1989 

Loma Prieta earthquake, and some also were shown to have 

undergone sinistral extension during the 1906 San Francisco 

earthquake (Prentice and Schwartz, 1991). The overall surface 

orientation and sense of displacement of these coseismic 

cracks and fissures in the Summit Road-Skyland Ridge fault 

zone is inconsistent with the expected deformation field for 

the 1989 Loma Prieta earthquake inferred from first-motion 

studies of the main shock (fig. 4). The net updip-slip vector 

of the hanging-wall fault block, as determined from the rake 

of the first motion at the focal depth of 17 km, was toward 

N. 10° W. Net maximum horizontal extension at this depth 

was, therefore, at about N. 80° E.–S. 80° W. The direction of 

maximum horizontal extension derived from the observed 

surface-deformation field along Summit Road and Skyland 

Ridge was about N. 10°–30° E. to S. 10°–30° W. (Ponti and 

Wells, 1991)—a significant rotation from the expected orientation (fig. 4A).

Ponti and Wells (1991) analyzed and modeled the 1989 

deformation field of the Summit Road-Skyland Ridge fault 

zone. They found that the total extension across this zone after 

the earthquake was at least 35 times that predicted from the 

geodetically measured uplift and that the azimuths of extension statistically point downslope. They concluded that the 

coseismic cracking and fissuring was predominantly gravity 

driven and that recurrent ridgetop spreading and downslope 

movement of near-surface materials was triggered by seismic 

shaking. Other workers (for example, Cotton, 1990; Cotton 

and others, 1990) interpreted the cracks and fissures as 

predominantly bedding-parallel, bending-moment tension 

faults associated with preexisting folds in the flexed hanging-wall block of the Loma Prieta rupture, and discounted 

slope processes as a major cause of the deformation.

The large-scale collapse of the northeast-facing slopes 

between the Lexington Reservoir and Santa Rosalia Mountain 

(fig. 2B) that is indicated in our geologic mapping is consistent 

with long-term uplift and recurrent focused seismic shaking 

of the relatively soft, incompetent Tertiary rocks in the steeply 

dipping hanging-wall block southwest of the San Andreas 

fault. Numerous coseismic cracks and fissures in this area 

crossed large landslides whose headwalls are along the crests 

of Summit and Skyland Ridges.

Other coseismic cracks, however, are not so obviously relatable to landslide activation. Some of these coseismic cracks 

have straight traces that cross ridgecrests, are aligned for 

distances of 1 km or more with other subparallel features (see 

pl. 1; fig. 2B), and appear to cut across, and root into bedrock 

outside of and below, the basal-slip surfaces of landslides (see 

pl. 1). Other cracks with straight traces that are confined to 

landslides may have originated as faults before their incorporation into landslides and so are probably displaced downslope 

from inplace faults beneath the landslide basal-slip planes.

Prentice and Ponti (1994) suggested that previous interpretations of data from a post-1906 survey (Lawson, 1908) of the 

now-inaccessible Wrights Railroad Tunnel located beneath 

Summit Ridge erroneously concluded that the tunnel was 

deformed over a broad zone. They showed that the data can 

be reinterpreted to indicate that slip occurred along a single 

fault trace near the northeast end of the tunnel. If correct, this 

reinterpretation makes it unlikely that the Summit Road-Skyland Ridge fault zone ruptured to more than 200-m depth (the 

depth of the tunnel below Summit Ridge) in 1906 or 1989, 

although deeper seated earlier faulting is possible.

A prominent bulge, or convex-northeastward bend, in the 

principal trace of the San Andreas fault northeast of Summit 

Ridge is apparent from the fault configuration revealed in 

exposures beneath large landslides dissected by tributaries to 

Los Gatos Creek (see pl. 1; fig. 2B). If the mapped exposures 

of the San Andreas fault are in place beneath these landslides, 

as implied by our mapping, then this bulge may be local flexure or warping of the fault plane due to lateral spreading and 

extension within the adjacent Summit Road-Skyland Ridge 

fault zone. Alternatively, if the exposures of the San Andreas 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E28 EARTHQUAKE OCCURRENCE

fault mapped as in place are, in fact, underlain by deeper seated 

landslide slip surfaces, then the mapped fault is displaced 

downslope, and the principal trace is buried upslope beneath 

the overriding landslides.

Features mapped as faults on Summit Road and Skyland 

Ridge, which are not within landslides (see pl. 1; fig. 2B), 

imply some unknown component of sinistral extension unrelated directly to slope failure. A more thorough study of the 

composition, material properties, and deep structure beneath 

the Loma Prieta region may provide some further basis for 

interpreting how deformation was propagated from depth 

to the surface, as well as for separating tectonic from slopedriven components of deformation. Pertinent topics include 

the following data.

1. Vertical and lateral inhomogeneities in the crust.—The 

main shock and most aftershocks of the 1989 Loma Prieta 

earthquake occurred in crystalline basement rocks beneath 

the gabbro of Logan, and primary slip did not propagate 

above about 8-km depth but, instead, stopped below the 

gabbro (see Jachens and Griscom, this chapter). The 

compositions, seismic velocities, and inferred material 

properties of rocks above the earthquake focus also vary 

nonuniformly with depth, as well as laterally across the 

Zayante fault and the principal trace of the San Andreas 

fault (Eberhart-Phillips and others, 1990; T.M. Brocher 

and R.D. Catchings, oral commun., 1993; see Brocher and 

others, this chapter; Jachens and Griscom, this chapter). 

Furthermore, consideration of vertical changes in crustal 

composition can be important in reconciling disparities 

between seismically and geodetically derived fault-planerupture models (Lisowski and others, 1990; Marshall and 

others, 1991; Eberhart-Phillips and Stuart, 1992). Thus, 

these data raise the issue of how the deep focus of the 

main shock (17 km) and the complex vertical and lateral 

variations in crustal composition between the focus and 

surface might have affected the orientation of the surfacedeformation field.

2. Heterogeneous slip.—Another factor that could have significantly affected the location, orientation, amount, and 

sense of surface deformation is heterogeneous slip within 

the Loma Prieta rupture zone. Beroza (1991) showed that 

the earthquake ruptured primarily in a reverse sense to the 

northwest, beneath the Summit Road-Skyland Ridge area, 

and primarily right laterally to the southeast. One possible 

interpretation of these data would be that the hangingwall block underwent significant clockwise rotation in the 

plane of rupture. If such rotation were substantial, sinistral 

extension conceivably could occur in the hanging-wall 

block at the surface—an interpretation consistent with 

the observed deformation on Summit Road and Skyland 

Ridges (fig. 4B).

3. Rupture-end effect.—The hanging-wall rocks above an 

oblique dextral-reverse rupture that terminates before 

reaching the surface must accommodate the vertical slip 

that occurred along the underlying rupture plane. The horizontal component of motion must also be accommodated 

at the northwest and southeast ends of the rupture, so that 

horizontal stress is focused at the rupture ends. Some component of the sinistral extension exhibited in the Summit 

Road-Skyland Ridge surface-deformation field, therefore, 

might be viewed as an accommodation effect of horizontal 

and vertical stresses focused at the northwest termination 

of the Loma Prieta rupture.

Although these data do not establish a tectonic origin for the 

Summit Road-Skyland Ridge surface-deformation field, they 

do suggest a reasonable basis for suspecting the existence of 

a tectonic component to this deformation that could involve 

the interplay of one or more of the above factors.

STRUCTURE NORTHEAST OF

THE SAN ANDREAS FAULT

Northeast of the San Andreas fault, the San Francisco 

Bay region is also dominated by large fault-bounded blocks. 

The faults bounding these blocks trend west-northwest and 

have poorly constrained and differing displacement histories 

(figs. 1B, 2A). The exposed basements of these fault blocks 

are composite units, consisting of Franciscan rocks, Coast 

Range ophiolite, and Great Valley sequence. Most of the 

faults bounding the major structural blocks are delineated by 

discontinuous lenticular bodies of mafic to ultramafic igneous rocks derived from the Coast Range ophiolite that have 

conspicuous associated aeromagnetic anomalies. Although 

the tectonostratigraphy may vary in detail between the fault 

blocks, some elements of the pre-Quaternary stratigraphy 

are similar in most of the blocks between the San Andreas 

fault and the Silver Creek fault zone to the east. Northeast of 

the Silver Creek fault, however, many of the Cretaceous and 

younger rocks are dissimilar in age and tectonostratigraphic 

history to age-equivalent rocks southwest of the fault. Rocks 

northeast of the Silver Creek fault are, therefore, considered 

here to have undergone major displacement (probably dextral) 

relative to rocks to the southwest. The rocks present between 

the Silver Creek and San Andreas faults are herein referred 

to as the “San Francisco Bay block,” which is considered to 

include all of the subsidiary fault blocks between the San 

Andreas and Silver Creek faults. We discuss here only those 

faults between the San Andreas fault zone and the southwest 

side of the Santa Clara Valley.

Folds are poorly expressed in pre-Miocene rocks of the San 

Francisco Bay block and mainly occur within the thin Miocene 

to Pleistocene cover of marine and nonmarine deposits. These 

units locally display high-amplitude, open to tightly appressed, 

upright to northeastwardly overturned folds.

We have divided the San Francisco Bay block into several 

subsidiary fault blocks (figs. 1, 2), modified after the original 

subdivisions introduced by Bailey and Everhart (1964). In 

the following subsections, we describe the faults that bound 

major subsidiary fault blocks of the San Francisco Bay block, 
E29

and discuss the constraints on their geometries, the style and 

timing of their displacements, and, where evident, the geometry of the folding and its relation to the faulting.

SARGENT FAULT

The Sargent fault was first described by Allen (1946)

 in the area northwest of San Juan Bautista (fig. 1A), where 

it dips nearly vertically. In the Loma Prieta region, the Sargent fault dips 55°–85° SW. and generally strikes N. 60°–74° 

W. throughout most of the Los Gatos, Laurel, and Loma 

Prieta 7A-minute quadrangles. Beneath Loma Prieta peak, a 

segment of the Sargent fault zone dips 25°–35° SW. (Loma 

Prieta thrust fault) and places Jurassic rocks of the Coast 

Range ophiolite and a thin sliver of (Jurassic?) slate and phyllite of Loma Prieta peak over Eocene rocks. This low-angle 

thrust segment is truncated against a trace of the younger, 

geomorphically youthful, more steeply dipping Sargent fault 

that crosses the divide between the Los Gatos Creek and 

Uvas Creek drainages on the southwest side of Loma Prieta 

peak (fig. 2B). Moderately shallow to subhorizontal rakes of 

slickensides on some shear surfaces along the steeply dipping 

segment of the fault zone indicate a conspicuous component 

of dextral strike slip; other shear surfaces indicate dip slip 

(Sebrier and others, 1992). Thus, the Sargent fault is predominantly a reverse fault, up on the southwest side, overprinted by 

later dextral slip. Locally, faults that are subsidiary, parallel, 

and adjacent to the Sargent fault, particularly in the vicinity 

of Loma Prieta peak, display conspicuous silicified, slickensided, quartz-lined surfaces with multilayered fault mullion 

of different orientations, suggesting episodic changes in the 

relative amounts of dip slip and strike slip. Also, the presence 

of a geomorphically youthful, steeper, south-dipping segment 

of the Sargent fault that cuts the thin thrust sheet of ophiolitic 

and metamorphic basement suggests that fault movement has 

been distinctly episodic and has shifted to more steeply dipping fault segments during late Cenozoic time.

The Sargent fault displaces early Miocene rocks along 

its course. Linear geomorphology, entrainment of young 

alluvium, sag ponds, and trenchlike topography indicate that 

the fault has been active during late Quaternary time. The 

youngest topographic expression of the fault occurs along 

the most steeply dipping segments of the fault zone, notably 

southeast of California Highway 152, and to the northwest 

along the 15-km-long segment between Sveadal and Lake 

Elsman (see pl. 1; fig. 2B).

At the northwest end of this segment of the Sargent fault, for 

a distance of approximately 2.4 km, surface rupturing occurred 

during the 1989 Loma Prieta earthquake. East-west-oriented 

tension cracks in this area had components of dextral extension of as much as about 10 cm, with approximately equal 

components of vertical displacement (figs. 6A, 6B). Other 

east-west-oriented tension cracks with components of vertical 

displacement of as much as 76 cm had smaller components 

of dextral extension of about 2 to 3 cm. Many of the extensional cracks with large components of vertical displacement 

were relatable to slumps and slope failures. Less abundant, 

southwest-side-up, N. 55° W.-oriented compressional cracks, 

including “popups” and one mole track, also occurred. The 

reverse-sense compressional cracks typically indicated components of vertical displacement of 5 to 6 cm. On the basis of 

more detailed mapping of this deformation zone, Aydin and 

others (1990, 1992) concluded that the overall deformation 

pattern indicates southwest-side-up reverse-dextral slip.

Along most of its length, the Sargent fault entrains a thin 

zone, ranging from less than 30 to about 300 m in thickness, 

of mafic to ultramafic rocks assigned to the Coast Range 

ophiolite. Aeromagnetic data (see Jachens and Griscom, 

this chapter) indicate that a thicker magnetic body than that 

exposed at the surface underlies the Sierra Azul block at 

depth. We therefore infer that the surface exposures of ophiolitic rocks have been sliced away and displaced upward from 

the thicker slab of buried magnetic rocks which underlies the 

Mesozoic to Tertiary sedimentary section of the Sierra Azul 

block north of the Sargent fault. On the basis of modeling of 

the buried magnetic slab from northeast-southwest-oriented 

gravity and magnetic profiles (see Jachens and Griscom, this 

chapter) and our geologic mapping of the ophiolitic section 

(cross secs. S–3, S–5, pl. 1), the top of the ophiolite slab in 

the footwall of the Sargent fault is estimated at 1.4-km depth. 

The magnetic models alone suggest a somewhat greater depth 

to the top of the magnetic source, closer to 2.0 km, possibly 

owing to the presence of less magnetic rocks in the upper 

part of the ophiolite. The magnetic models also suggest that 

the ophiolite slab is generally thicker (1.5–3.0 km) than the 

0.8- to 1.7-km-thick slab depicted in our structure sections 

based on surface geology (see pl. 1). The downdip separation measured on the Sargent fault (cross secs. S–5, S–6, pl. 

1) is at least 1.5 to 1.8 km. On the basis of a dip range of 

50°–70°, the minimum horizontal shortening across the fault 

is estimated at 0.5 to 1.1 km.

The timing of most displacement across the Sargent fault 

is constrained by the youngest offset rocks, by K-Ar dating 

of hydrothermal alteration in the fault zone, and by fissiontrack data. Youthful geomorphology and faulted Miocene to 

Pliocene rocks along the fault between Gilroy and Mount 

Madonna (fig. 1B) indicate that the Sargent fault has remained 

active into the Quaternary (McLaughlin, 1974).

Potassic hydrothermal alteration that locally occurs along 

the Sargent fault zone is a result of fluids channeled along 

the fault, and so the age of the hydrothermal mineralization 

is a minimum for the age of faulting. Alteration in the fault 

zone northeast of Mount Madonna (fig. 1B) is K-Ar dated at 

about 17 Ma (McLaughlin and others, 1990, 1991b), indicating that slip was initiated by at least 17 Ma. To the northwest 

at Loma Prieta peak, where a thrust segment of the Sargent 

fault underlies the peak, potassic hydrothermal alteration in 

the hanging wall and footwall of the thrust is K-Ar dated at 

about 10 Ma (McLaughlin and others, 1990, 1991b).

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E30 EARTHQUAKE OCCURRENCE

Fission-track data (Arrowsmith and others, 1992; Bürgmann and others, 1994) in the vicinity of Loma Prieta peak 

indicate that the rocks northeast and southwest of the Sargent 

fault have been uplifted about 3.5 to 4.0 km since 4.0 Ma. Any 

major vertical displacement accommodated by the Sargent 

fault, therefore, must have predated 4.0 Ma (fig. 3).

The occurrence of potassic hydrothermal alteration in both 

the hanging wall and footwall of the Sargent fault zone at 

Loma Prieta peak indicates that vertical displacement across 

the fault since 10 Ma has been less than the vertical extent of 

the hydrothermal system. Fission-track data and radiometric 

ages indicate that the presently exposed quartz-adularia veins 

composing the hydrothermal zone at Loma Prieta peak were 

at a depth of at least 3 km at 4 Ma. Well-studied quartzadularia vein systems elsewhere in the northern California 

Coast Ranges and in New Zealand have generally formed in 

upper-mesothermal and epithermal depth ranges of 0 to 2 

km (Browne and Ellis, 1970; McLaughlin and others, 1985; 

Peabody and Einaudi, 1992; Rytuba, 1993). At 3- to 4-km 

depth, veins in the hanging wall and footwall of the thrust 

fault beneath Loma Prieta peak thus are probably in the lowermost part of the hydrothermal system. This hydrothermal 

system probably could not have been appreciably deeper 

when it formed at 10 Ma, and so most vertical displacement 

may have been accommodated by the Sargent fault between 

17 and 10 Ma.

In summary, the Sargent fault has accommodated at least 

1.4 km of uplift, 1.5 to 1.8 km of reverse slip, and 0.5 to 1.1 

km of shortening. This deformation began as early as 17 Ma 

and may have largely ceased by 10 Ma (fig. 3).

In the sections above on stratigraphy, we noted that the 

slate and phyllite of Loma Prieta peak occurs in a thin fault 

sliver along the sole of the Loma Prieta thrust and that these 

metamorphic rocks have no known Coast Range source. 

The composition, metamorphic texture, and low greenschist 

metamorphic grade of these rocks are nearly identical to 

those of rocks of the Jurassic Mariposa Slate in the western 

Sierra Nevada foothills. It therefore was proposed (Blake and 

others, 1991; McLaughlin and others, 1991b) that the slate 

and phyllite of Loma Prieta peak is part of the western Sierra 

Nevada metamorphic belt of crystalline rocks, which continues from the Sierra Nevada foothills southwestward beneath 

the Franciscan Complex of the Coast Ranges (Wentworth and 

others, 1987; Fuis and Mooney, 1990). If this model is valid, 

then the slate and phyllite of Loma Prieta peak has undergone 

considerable vertical displacement because basement rocks 

with a velocity structure (Fuis and Mooney, 1990; Blake and 

others, 1991) similar to that of crystalline basement in the 

Sierra Nevada foothills lie at 12- to 15-km depth just northeast 

of the San Andreas fault. The Sargent fault, however, appears 

to have accommodated no more than about 1.4 km of offset 

of the ophiolitic basement of the Sierra Azul block. Thus, 

any uplift of the slate and phyllite of Loma Prieta peak from 

more than about 1.4-km depth and accompanying shortening must largely predate 17 Ma and probably occurred on a 

fault(s) other than the Sargent. The results from fission-track 

studies (Arrowsmith and others, 1992; Bürgmann and others, 

1994) suggest that the fault block between the Sargent and 

San Andreas faults has been uplifted about 3 to 4 km since 

5–3 Ma, yielding uplift rates of 0.6 to 1.3 mm/yr—almost 3 

times that determined across the Sargent fault and, therefore, 

possibly related to deeper seated structures.

Kinematic data suggest that subequal or larger amounts 

of strike slip have occurred along the Sargent fault during 

the latest stages of its offset history, although, as determined 

from the surface geology, the amount of dextral strike slip 

on the fault is relatively unconstrained and, therefore, unresolvable at this time.

HOOKER GULCH-SIERRA AZUL FAULT ZONE

The Hooker Gulch fault is a reverse fault, about 4.8 km 

long, that is truncated by the San Andreas fault on the west. 

The Hooker Gulch fault dips vertically to about 60° SW. and 

is offset by a younger, north-south-trending normal fault south 

of Mount Umunhum (fig. 1A). To the east of the offsetting 

normal fault, the continuation of the Hooker Gulch fault is 

known as the Sierra Azul fault. Reverse separation across the 

Hooker Gulch-Sierra Azul fault, measured from structure 

sections (cross sec. S–5, pl. 1), may be about 1.8 km, on the 

basis of separation of the top of the Coast Range ophiolite 

across the fault. Shortening of the Sierra Azul section across 

the fault is about 1 km, and the associated uplift or vertical 

displacement of the section is about 1.5 km. Thus, the amounts 

of reverse slip, uplift, and shortening across the Hooker 

Gulch and Sargent faults are similar. The Hooker Gulch fault 

appears to be a part of the same reverse-slip-fault system as 

the Sargent fault, as indicated by its subparallel orientation 

and nearly identical sense and amount of displacement. The 

relative component of dextral strike slip, if any, is unknown 

because no kinematic indicators were measured.

The dextral normal fault that offsets the Hooker GulchSierra Azul fault zone appears to have a dip-slip component 

of about 2 km. The fault extends into the basal part of the 

Coast Range ophiolite and locally, at its north end, extends 

into, and dies out in, underlying metaclastic rocks of the 

Franciscan Complex, as indicated in exposures on Mount 

Umunhum (cross sec. S–5, pl. 1). Here, the crossfault dips 

about 60° SW., with slickensides that rake 20° NW. and 

suggest that, at least locally, about 80 percent of the slip is 

dextral. Elsewhere, along the trend of the fault, slip vectors indicate components of strike slip in the range 50–90 

percent.

The timing of displacement on the Hooker Gulch-Sierra Azul 

fault is post-early Eocene, on the basis of offset of the Eocene 

sequence. The similarity of the fault zone in its orientation and 

displacement to the Sargent fault suggests that it may be of the 

same Miocene or younger generation. The timing of movement 

along the oblique-normal crossfault south of Mount Umunhum 
E31

also is Miocene or younger, possibly considerably later than 

that along the Hooker Gulch-Sierra Azul fault zone.

ALDERCROFT FAULT

The Aldercroft fault is the principal local boundary between 

rocks of the Sierra Azul block on the south and the Franciscan Complex to the north, and as such, it occupies the same 

structural position as the Coast Range fault of Jayko and 

others (1987), which extends throughout the Coast Ranges. 

As pointed out by Jayko and others, the Coast Range fault is 

a highly modified subduction boundary with an early record 

of Jurassic and Cretaceous compressional tectonism associated with oblique subduction of the Franciscan Complex. 

During the latest Cretaceous to early Eocene, low-angle 

normal faulting associated with uplift and attenuation of the 

Coast Range ophiolite and with unroofing of the Franciscan 

Complex was focused along this boundary. In the study 

area (fig. 1), the Aldercroft fault exhibits compressional and 

extensional structures that correspond to those described 

along the Coast Range fault (Jayko and others, 1987). In the 

Sierra Azul block, the extensional structures may be as young 

as Oligocene because subparallel extensional faults at a low 

angle to bedding cut Eocene, but not Miocene, rocks higher 

in the section. Thus, the Aldercroft fault is here presumed to 

have been coextensive with the Coast Range fault until the 

time of extensional deformation from the Late Cretaceous 

(possibly Campanian) to the Oligocene.

On the geologic map and structure sections (cross secs. 

S–3, S–5, pl. 1), the lowest fault in the Aldercroft fault zone is 

depicted as a southwest-dipping thrust fault, because it places 

older (Jurassic) ophiolitic rocks over younger (Cretaceous) 

Franciscan rocks along its length. However, structurally 

higher faults in the zone, which locally form the upper contact 

of the ultramafic section of the Coast Range ophiolite, either 

consistently place younger rocks down on the ophiolite or cut 

out intervening ophiolite section. These faults are extensional 

rather than compressional structures and were superposed on 

and, in places, may root into or below the basal thrust of the 

Aldercroft fault zone (cross secs. S–3, S–5, pl. 1).

Although the Aldercroft fault is similar in orientation to 

the Hooker Gulch-Sierra Azul fault zone and merges with 

these faults along strike, no convincing field evidence exists 

for early Miocene or younger reactivation of the Aldercroft 

fault as a thrust. The latest observed displacement along the 

segment of the Aldercroft fault between its intersection with 

the San Andreas fault and Mount Umunhum is normal, up on 

the north side. Locally, the fault is overturned and dips 45° 

NE. Near Mount Umunhum (fig. 1A), where the fault dips 

75° SW., some of this normal displacement may be accommodated by the Soda Springs fault, although the Soda Springs 

fault appears to have only minor displacement.

The amount of normal slip on the Aldercroft fault is 

undetermined in the study area (fig. 1), partly because slip 

mostly predated displacements along other faults (including the Sierra Azul, Soda Springs, and Berrocal faults) that 

cut or merge with the Aldercroft fault and break the Sierra 

Azul block into several subsidiary blocks. Late (Miocene 

and younger) offsetting compressional faults that followed 

earlier extensional deformation along the Aldercroft fault 

(cross sec. S–3, pl. 1) suggest that the Aldercroft fault is 

presently inactive.

Structure sections drawn across the Aldercroft fault and 

the Coast Range ophiolite northeast of Mount Thayer (fig. 2), 

based on an interpretation of the surface geology (cross sec. 

S–3, pl. 1), depict the Aldercroft fault as dipping about 55° 

SW., intersecting the San Andreas fault at about 6-km depth. 

This configuration generally agrees with the magnetic models 

of Jachens and Griscom (this chapter), although interpretation 

of aeromagnetic data suggests a more complex “wedging” 

relation between the Coast Range ophiolite and underlying 

Franciscan Complex at depth.

Gravity and magnetic models (see Jachens and Griscom, this 

chapter) suggest that the Coast Range ophiolite is segmented 

at depth by wedges of nonmagnetic rocks interpreted to be part 

of the Franciscan Complex. Jachens and Griscom’s interpretation of the aeromagnetic data implies that the surface trace of 

the Aldercroft fault may be the steeply tilted roof thrust to one 

such wedge of Franciscan rocks. Their model further suggests 

that the wedge composed of Franciscan Complex originally 

was emplaced eastward into a thick subhorizontal slab of 

ophiolite. This wedge then peeled up and incorporated thinner 

slabs of the ophiolite into the roof of the wedge complex. The 

present southwestward dip of the Aldercroft fault zone would 

have resulted from later southwestward tilting superimposed 

on the wedge complex.

This structural style is consistent with the interpretation 

(McLaughlin and others, 1988a) that the Coast Range ophiolite and Great Valley sequence are interleaved with rocks of the 

Central belt of the Franciscan Complex along folded imbricate 

low-angle faults. A model for eastward “wedge” emplacement 

(see Jachens and Griscom, this chapter) is also consistent with 

the field observation that the present Aldercroft fault indicates 

down-to-the-southwest displacement. The gravity and magnetic models do not agree, however, with our interpretation 

that this down-to-the-southwest displacement is due to extension, as suggested by structures (boudinage and attenuation) 

in the ophiolite. If the displacement resulted from thrusting 

associated with emplacement of a wedge, we would expect to 

see evidence of compression, such as eastward-vergent folds 

or duplexing, which are unrecognized.

FAULTS OF THE SOUTHWESTERN

SANTA CLARA VALLEY THRUST BELT

The thrust belt along the southwest side of the Santa Clara 

Valley includes several anastomosing and branching fault 

zones, including the Lime Kiln, Berrocal, Shannon, and Monte 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E32 EARTHQUAKE OCCURRENCE

Vista faults. Focused compressional deformation occurred 

across the Monte Vista and Berrocal-Shannon fault zones 

during the earthquake (Haugerud and Ellen, 1990; D.H. Sorg 

and F.J. Groffie, unpub. data, 1990), strongly suggesting that 

these thrust zones root into the San Andreas fault and (or) that 

reverse slip on these thrusts is driven by horizontal shortening 

across the San Andreas fault zone. The faults of this thrust belt 

dip predominantly southwest at moderately low to steep angles 

(35°–60°), as judged from the curvilinear upcanyon V-patterns 

of their surface traces, the measured attitudes of shear surfaces 

in the fault zones, and the subsurface exposures of faults in 

mines of the New Almaden mercury district (Bailey and Everhart, 1964). These faults vary in dip along strike, however, and 

locally are overturned to the northeast. Faults of the Berrocal 

and Shannon fault zones include those dipping 40°–55° NE. 

that can mostly be interpreted as subsidiary normal faults in 

the hanging-wall blocks of the major southwest-dipping reverse 

faults (cross secs. S–1, S–2, S–5, pl. 1). Map relations along 

the Berrocal, Shannon, and Lime Kiln fault zones suggest that 

these faults include segments of older, folded, sheetlike fault 

slivers, whose boundaries are commonly delineated by thin 

serpentinite bodies (cross sec. S–2, pl. 1). Some of the thin, 

deformed serpentinite bodies in these fault zones are derived 

from the lower part of the Coast Range ophiolite.

The various faults of the southwestern Santa Clara Valley 

thrust belt traverse some of the more densely populated areas 

of the San Francisco peninsula and locally display relatively 

clear evidence of Quaternary displacement. The relation of the 

thrust belt to the San Andreas fault, however, has remained 

unclear. In particular, crustal shortening and related deformation that occurred across the Berrocal, Shannon, and Monte 

Vista fault zones during the 1989 Loma Prieta earthquake have 

raised the questions of (1) earthquake potential, (2) the extent 

to which faulting along the thrust belt coincides with large 

earthquakes on the San Andreas fault (Haugerud and Ellen, 

1990; D.H. Sorg and F.J. Groffie, unpub. data, 1990), and (3) 

whether faulting and seismicity may occur independently of 

events on the San Andreas fault.

The Berrocal, Shannon, and Monte Vista faults have offset 

Pliocene and Pleistocene strata of the Santa Clara Formation 

at numerous localities, including exposures observed in several 

trenches (fig. 6). The distribution, lithofacies variations, and 

deformation of the Santa Clara Formation and some younger 

nonmarine Quaternary deposits indicate that the deposition 

of these largely fluvial units was penecontemporaneous with 

uplift, shortening, and faulting across the thrust belt. The 

youngest deposits observed to be cut by faults of the thrust 

system are present along a northeast-dipping normal-fault 

segment of the Shannon fault zone, exposed in a trench at the 

Senator Mine, west of New Almaden, in 1989 (figs. 2A, 6). 

At this locality, a 15- to 18-ka paleosol (Emcon Associates, 

1990) that formed on the eroded Temblor Sandstone is cut by a 

north-dipping shear whose sense and amount of displacement 

are uncertain (fig. 6). This fault is subsidiary to a conspicuous 

northeast-dipping normal fault within the hanging-wall block 

of the main southwest-dipping segment of the Shannon fault. 

Radiocarbon ages from an overlying lower Holocene colluvial 

unit that is uncut by shearing suggest that the faulting probably 

was pre-Holocene.

At another locality, along the Berrocal fault zone, near the 

intersection of Wood Road and Santa Cruz Avenue in the town 

of Los Gatos (figs. 2A, 6), undated Pleistocene terrace sand 

and gravel, estimated at 120–250 ka (see sections above on 

stratigraphy), are cut by a 58°– 60° SW.-dipping segment of 

the Berrocal fault, which thrusts volcanic rocks of the Franciscan Complex upward and northeastward over the terrace 

gravel. The hanging-wall Franciscan rocks have a postterrace 

reverse-slip component of ≥0.7 m. The fault scarp is draped 

with Holocene colluvium, which is unfaulted.

The Lime Kiln fault in this thrust belt is not as well studied as the Berrocal and Shannon faults. However, the Lime 

Kiln fault locally exhibits moderately strong physiographic 

linearity in the vicinity of Lime Kiln Canyon southeast of the 

Lexington Reservoir (fig. 2A). In this area, poorly exposed 

elevated terrace gravel may be offset along the fault (see pl. 

1), but more detailed study of these exposures is needed to 

determine whether offset has occurred.

Thus, the southwestern Santa Clara Valley thrust belt 

locally exhibits evidence for surface faulting during the 

latest Pleistocene (20–12 ka). The fault that exhibits this 

evidence is a secondary north-dipping normal fault. A widespread pulse of faulting appears to have occurred since the 

late Pleistocene (250–120 ka) along the length of the thrust 

belt. Several of the faults now exposed at the surface within 

the thrust belt appear not to have ruptured to the surface 

during all or part of the Holocene, as suggested by colluvial 

draping of Pleistocene fault scarps. We note, however, that 

many of the thrust faults of this northeastward-vergent fault 

system deform Miocene to Pleistocene hanging-wall rocks 

into open to tightly appressed folds with axes that parallel 

the curvilinear trend of the thrust belt. Many of the faults 

now exposed at the surface may have propagated upward 

from depth as blind thrusts within the range-front thrust belt. 

In a fold-and-thrust belt, blind thrusts generally propagate 

upward, deforming the overlying sedimentary package into 

folds that take up vertical and compressional components of 

slip above the depth of rupture propagation. Typically, such 

faults steepen as they propagate to the surface and (or) rotate 

to steeper positions after reaching the surface, as newly active 

blind thrusts begin propagating upward and outboard from 

the root zone of the thrust system at depth. The open, highamplitude folds in the Santa Clara Formation and underlying 

Miocene strata in the low uplands between the Monte Vista 

and Berrocal fault zones (Cupertino-Saratoga area, fig. 2A) 

probably owe their origin to slip on blind thrusts of the Monte 

Vista and Berrocal fault zones during the late Pleistocene 

(Sorg and McLaughlin, 1975). The Monte Vista fault zone is 

probably a former blind thrust that has breached the axis of its 

hanging-wall anticline, broken the surface, and rotated into a 

steeper, southwest-dipping orientation. If this interpretation 
E33

is correct, then active blind thrusts that root into the San 

Andreas or Monte Vista fault at depth might exist beneath 

the Santa Clara Valley northeast of the main exposures of 

the Monte Vista fault.

Sebrier and others (1992) reported that the southwestern 

Santa Clara Valley thrust belt has undergone a multiphase 

kinematic history. The chronology of outcrop-scale fault 

features and fault-slip indicators in the Monte Vista, Berrocal, and Shannon fault zones indicates that the earliest slip 

along these faults was reverse and up to the southwest. This 

primary reverse slip is overprinted by a later generation of 

predominantly strike slip features, which include both dextral and sinistral components. Excluding the 1989 coseismic 

compressional deformation (Haugerud and Ellen, 1990), a late 

pre-Loma Prieta generation of features indicative of normal 

(transtensional) displacement is also recorded at localities in 

the Monte Vista and Shannon-Berrocal fault zones (Sebrier 

and others, 1992). The chronology of change in slip modes 

on these faults, from dominantly thrusting to strike slip, is 

interpreted to have occurred as faults within the thrust belt 

propagated toward the surface and rotated into more steeply 

dipping orientations. The late, extensional indicators are 

questionably related by Sebrier and others to secondary 

extension associated with hanging-wall folds or to transfer 

of slip from older, structurally high thrust faults to structurally lower, more recently activated blind thrusts beneath the 

Santa Clara Valley.

AGE OF THRUST-BELT INITIATION

AND AMOUNT OF DISPLACEMENT

Earliest slip on the Berrocal and Sargent faults is constrained 

by structural relations and by the ages of hydrothermal Kfeldspar veinlets along the Sargent fault zone, as previously 

discussed. The 17.8-Ma K-Ar age on adularia from the fault 

zone near Mount Madonna (fig. 1B) was determined on rocks 

from within a thin fault sliver bounded to the northeast by 

the southwest-dipping Berrocal fault and to the southwest by 

the structurally higher Sargent fault. Hydrothermal alteration occurs sporadically northward along both the Sargent 

and Berrocal faults, as indicated by local concentrations of 

silica-carbonate rock, an alteration product of serpentinite, 

and by hydrothermally bleached and veined clastic rocks. We 

interpret this alteration to indicate that the Sargent and Berrocal faults were both conduits for hydrothermal circulation 

at 18–17 Ma. Juxtaposition of rocks in the Sierra Azul block 

with the Temblor Sandstone and the Monterey Shale along the 

Berrocal fault also indicates that major reverse-slip displacement postdated the middle Miocene.

Local unroofing and uplift of the Coast Ranges before late 

early Miocene time is reflected in an unconformity at the 

base of the Temblor Sandstone in the New Almaden area (fig. 

2A). Low-angle attenuation faults that cut Eocene and older 

rocks in this area are unrecognized in the Miocene section, 

suggesting that a change from extensional faulting to northeastward-vergent thrusting occurred during the Oligocene to 

early Miocene. Although the data are equivocal, we find no 

convincing evidence in the depositional or structural record 

to indicate that displacements or hydrothermal circulation 

occurred along the Sargent or Berrocal faults any earlier than 

17.8 Ma. We therefore speculate that deep-seated thrusting 

was initiated on the Sargent and Berrocal faults at 18–17 

Ma. This early contractional deformation apparently was 

overprinted by younger thrusting by 3.0 Ma along the Monte 

Vista, Berrocal, and Shannon faults that cut strata of the Santa 

Clara Formation.

A major episode of Miocene to Pliocene shortening, accompanied or followed by strike-slip faulting along the thrust belt, 

is suggested by major structural and depositional discordance 

between the Santa Clara Formation and underlying Miocene 

rocks, as well as by the depositional-facies contrasts between 

coeval Miocene deposits northeast and southwest of the 

Berrocal-Sargent fault zone, as noted above in the sections 

on stratigraphy. The contrast in depositional settings across 

the thrust belt indicates that the Miocene deep-water margin 

was displaced relatively landward (northeastward) over the 

adjacent Miocene continental slope and shelf. By analogy 

with the average width of the modern Continental Shelf and 

upper slope of California and Oregon, as much as 50 km of 

the Miocene shelf and slope width appear to have been tectonically removed. Since the Berrocal and Sargent faults both 

lie within 6 to 10 km of the San Andreas fault, we conclude 

that a considerable amount of the missing Miocene shelf to 

slope must have been removed by strike-slip motion along 

and (or) within the thrust belt. If the missing margin were 

completely the result of horizontal shortening, considerably 

greater topographic relief should be present northeast of 

the San Andreas fault. Nevertheless, a major unconformity 

between the Miocene section and the overlying Santa Clara 

Formation suggests that major erosion, probably associated 

with uplift and horizontal shortening, occurred after deposition of the Monterey Shale before and concurrent with the 

onset of Santa Clara deposition (5–2 Ma).

The aggregate components of reverse and strike slip across 

the southwestern Santa Clara Valley thrust belt are relatively 

large, although the amount of strike slip remains obscure. A 

gravity study across the Monte Vista segment of the thrust 

belt in the Cupertino and Los Altos Hills areas suggests 

that Franciscan rocks are thrust eastward at least 2.3 km 

over low-density Miocene and Pliocene to Pleistocene strata 

along a 9.7-km-long segment of the fault zone (Fleck, 1967; 

McLaughlin, 1974), indicating at least 2 to 3 km of horizontal 

shortening in those areas since the Miocene (10 Ma). Assuming this amount of horizontal shortening, as well as an average 

dip of 45° across the Monte Vista segment of the thrust belt 

(Sorg and McLaughlin, 1975), about 3 km of uplift and 4 km 

of reverse slip have been accommodated along the Monte 

Vista fault zone. Uplift and horizontal shortening rates have 

thus been about 0.3 mm/yr, if displacement is distributed over 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E34 EARTHQUAKE OCCURRENCE

Figure 6.—Pleistocene and younger faulting along reverse faults northeast of the San Andreas fault in the Loma Prieta region, Calif. (fig. 1). 

A, Coseismic surface rupture during the 1989 Loma Prieta earthquake along the Sargent fault on north side of the Elsman Reservoir. Rightstepping cracks within fault zone indicate dextral extension with a 10-cm right-slip component of up-to-the-north slip. View westward. B, 

Closeup of coseismic surface rupture in figure 6A, showing right-stepping dextral cracking along the Sargent fault at the Elsman Reservoir. 

Extension direction is N. 75° E.–S. 75° W. C, Thrust faulting within the Santa Clara Formation at construction site along Permanente Road 

in Cupertino (fig. 2A). Main trace of the Monte Vista fault is to southwest (right). Photograph taken in 1973. D, Closeup of thrust fault in 

figure 6C. View southeastward. E, Faulted Pleistocene fluvial gravel and sand along the Berrocal fault at construction site on Wood Road 

near downtown Los Gatos (fig. 2A). Fault is up on southwest (left) side. Photograph taken in October 1989. F, Closeup of faulted fluvial 

deposits in figure 6E. Colluvium overlies faulted contact between Franciscan volcanic rocks and Pleistocene deposits. G, South strand of the 

Shannon fault trenched by Emcon Associates at the Guadalupe Landfill on the site of the Senator Mine, 1.3 km west of Guadalupe Creek. 

Hammer straddles north (left)-dipping fault in the Temblor Sandstone, which extends upward into dark paleosol, estimated at 15–18 ka. 

View southeastward.

A

B

C

D

E
E35

the past 10 m.y. (that is, if uplift began soon after deposition 

of the Miocene strata). If most of the uplift and horizontal 

shortening have occurred since the onset of deposition of the 

Santa Clara Formation, about 5–2 Ma, then late Quaternary 

uplift and horizontal shortening rates could have approached 

or exceeded 1.0 mm/yr, an interpretation supported by youthful topographic relief, by tight high-amplitude folds in the 

Santa Clara Formation, and by locally conspicuous angular 

depositional discordance of the Santa Clara Formation with 

Miocene and older rocks.

Studies along the Monte Vista fault zone in Cupertino (fig. 

2A; Sorg and McLaughlin, 1975; E.J. Helley, oral commun., 

1991) show that a 120-ka fluvial terrace is vertically displaced 

about 43 m across the Monte Vista fault at the mouth of Stevens 

Creek Canyon, yielding a post-120-ka uplift rate of about 0.4 

mm/yr. This uplift rate is close to that indicated by gravityderived horizontal shortening in the past 10 m.y. (Fleck, 1967) 

and supports an interpretation that the uplift rate has been 

relatively constant since the late middle Miocene. Alternatively, 

the data could be interpreted as indicating a high uplift and 

horizontal-shortening rate of about 1.0 mm/yr between 3 and 

0.12 Ma, which subsequently decreased to 0.3 mm/yr.

In summary, structural relations and gravity data suggest 

that as much as 3 km of uplift and horizontal shortening and 

4 km of reverse slip have been accommodated across the 

southwestern Santa Clara Valley thrust belt since 10 Ma. 

Uplift and horizontal-shortening rates have been about 0.4 

mm/yr since 120 ka, about the same as those calculated by 

assuming a constant slip rate since 10 Ma (0.3 mm/yr). The 

stratigraphic and structural discordance between the Santa 

Clara Formation and Miocene and older formations, however, 

strongly suggests that uplift and horizontal-shortening rates 

increased markedly with the initiation of deposition of the 

Santa Clara Formation between 4.8 and 1.9 Ma. If most of the 

observed uplift and horizontal shortening were accommodated 

during this later episode of deformation (4.8–0.12 Ma), then 

F G

Figure 6.—Continued

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E36 EARTHQUAKE OCCURRENCE

rates were about 0.6 to 1.7 mm/yr and decreased to about 0.3 

mm/yr after 120 ka.

The amounts of horizontal shortening, uplift, and reverse 

slip across the major thrust zones between the southwestern 

margin of the Santa Clara Valley and the San Andreas fault 

are listed in table 3 in an attempt to estimate the total Miocene 

and later contractional deformation. These data suggest that 

approximately 5 km of horizontal shortening, 6 km of uplift, 

and 8 km of reverse slip have been accommodated between the 

southwestern Santa Clara Valley thrust belt and the San Andreas 

fault (fig. 7). The observed amounts of horizontal shortening 

and uplift clearly do not approach what would be necessary to 

cause the observed loss of the Miocene marine margin, and so 

we conclude that a 45-km-wide segment of this margin has been 

eliminated through a large, but as yet ill-defined, component 

of strike slip. The total uplift and horizontal-shortening rates 

of 0.4 to 0.5 mm/yr since 18 Ma are consistent with regional 

uplift rates of less than 1 mm/yr for Quaternary marine terraces 

southwest of the San Andreas fault (Lajoie and others, 1979), 

with rates calculated for the southwestern Santa Clara Valley 

thrust belt since 10 Ma, and with the rate calculated from uplift 

of the 120-ka fluvial terrace across the Monte Vista fault. These 

rates are low, however, in comparison with uplift rates near the 

San Andreas and Sargent faults based on fission-track studies, 

which are about 0.8 to 1.0 mm/yr since 5–3 Ma (Arrowsmith 

and others, 1992; Bürgmann and others, 1994), and in comparison with uplift and horizontal-shortening rates of 0.6 to 1.7 

mm/yr for the thrust belt based on the assumption that most 

deformation occurred between 4.8 and 0.12 Ma. Structural 

relations and K-Ar dating of hydrothermal alteration suggest 

that most vertical and contractional displacement across the 

Sargent fault may have occurred before 10 Ma. Similar fissiontrack-derived uplift rates northeast and southwest of the Sargent 

fault must, therefore, reflect displacement above deeper faults 

beneath the Sargent fault. Correspondence of the 4.8–0.12-Ma 

uplift rate for the southwestern Santa Clara Valley thrust belt 

with the fission-track rate for approximately the same time 

frame suggests that rapid uplift adjacent to the San Andreas 

fault could have been accommodated by the southwestern Santa 

Clara Valley thrust belt. In figure 7, we suggest a model for the 

partitioning of contractional deformation normal to the San 

Andreas fault among faults of the Sierra Azul block and those 

of the southwestern Santa Clara Valley thrust belt.

LEXINGTON FAULT ZONE

The Lexington fault zone (see pl. 1; figs. 1, 2) strikes northsouth and is approximately 4 km long by 0.4 km wide. This 

fault zone is of particular interest because it appears to link 

the dextral-slip San Andreas fault zone with the southwestern 

Santa Clara Valley thrust belt (McLaughlin and others, 1991a, 

1992). Faults of the Lexington fault zone dip both east and 

west, but the principal slip has been along a reverse fault at 

the east side of the zone dipping about 70° E. Thus, the fault 

zone is regarded as predominantly contractional (McLaughlin 

and others, 1991a, 1992; Sebrier and others, 1992). Los Gatos 

Creek is incised along the Lexington fault zone, and the main 

fault is mappable as a conspicuous north-south lineament on 

aerial photographs.

The Lexington fault zone cuts the Santa Clara Formation. 

Near the Lexington Reservoir (fig. 2), the Santa Clara Formation 

consists of a basal lacustrine lithofacies, which dips southwest 

toward the San Andreas fault. The lacustrine lithofacies interfingers eastward with coarse-grained fanglomerates that are 

tilted steeply, truncated, and juxtaposed with Franciscan rocks 

Fault zone Shortening

(km)

Uplift

(km)

Reverse slip

(km)

Duration

of faulting

(Ma)

Sargent fault--------------- 1.1 1.4 1.8 18–10

Hooker Gulch fault------- .9 1.6 1.8 18–10

Southwestern Santa

Clara Valley thrust

belt.

3.0 3.0 4.2 10–0 or 4.8–0

Total---------------------- 5.0 6.0 7.8 18–0

Lexington fault ----------- 1.1–1.4 2.9–3.8 3.1–4.0 4.8–0

Table 3.—Estimated parameters for faults located between the San Andreas fault and the 

southwestern Santa Clara Valley thrust belt

[Values for the Lexington fault were not added to total fault displacement. The Lexington fault is nearly orthogonal to the 

main thrust front, and so its displacement may merely be partitioned off from the other fault zones]
0 10 KILOMETERS

0 5 10 MILES

5

SCALE

UPLIFT TOTAL 6.0 KM 

SAN ANDREAS SOUTHWESTERN SANTA CLARA VALLEY THRUST BELT 

HOOKER GULCH SARGENT

3.0

1.10.9

1.6

1.4

3.0

SHANNON FAULT ZONE BERROCAL FAULT ZONE HOOKER GULCH FAULT SARGENT FAULT

SAN ANDREAS FAULT ZONE

ZAYANTE

Topography

Graniticrocks

La Honda

Block

Basin fill

Gabbroicrocks

Depth (km)

SEA LEVEL

5

10

15

1.8

1.8

4.2

?

?

?

SW Santa Clara Valley Thrust Belt

Uplift across the San Andreas fault (schematic)

FAULT ZONE

A B

MAINSHOCK

Figure 7.—San Andreas fault system in the Loma Prieta region, Calif. (fig. 1). Plus, movement into plane; minus, movement out of plane. A, HypotheticAndreas fault zone, showing deep southwest- and northeast-dipping components inferred from seismic evidence (Olson and Hill, 1993). B, Schematic cof shortening, reverse slip, and uplift to thrust faults between the San Andreas fault and the southwestern Santa Clara Valley thrust belt since 18 Ma.
E38 EARTHQUAKE OCCURRENCE

along the main trace of the Lexington fault on the east side of 

the reservoir. These relations suggest that the Lexington fault 

was active, and that vertical displacement occurred, during the 

deposition of fanglomerates of the Santa Clara Formation. The 

fanglomerates apparently were shed westward from the growing Lexington fault scarp into a large deep lake along the San 

Andreas fault. The depression in which the lacustrine sediment 

was deposited may have been formed by a landslide dam or, 

alternatively, a structural barrier that blocked northward drainage along the San Andreas fault.

Flat-lying fluvial-terrace gravel unconformably overlies 

folded and tilted strata of the Santa Clara Formation along the 

Lexington fault zone. These younger Pleistocene gravel deposits 

are incised by Los Gatos Creek and are now at an elevation 

of about 200 m. The high elevation of this terrace relative to 

other fluvial terraces in the Santa Clara Valley that postdate the 

0.3–4.8-Ma Santa Clara Formation, the degree of soil development, and the degree of weathering of terrace clasts suggest 

that the terrace was cut during the 250-ka eustatic highstand 

(see sections above on stratigraphy). This age may also mark 

the time of capture of northwestward drainage along the San 

Andreas fault by the headward incision of northeastwardly 

draining Los Gatos Creek. As along the southwestern Santa 

Clara Valley thrust belt, uplift and faulting of the Santa Clara 

Formation occurred between about 4.8 and 0.25 Ma (McLaughlin and others, 1992). Subsequently, the 250-ka terrace gravel 

was incised approximately 152 m along the fault zone by Los 

Gatos Creek, yielding an incision rate and an approximate uplift 

rate of about 0.6 mm/yr. If continuous for 4.8 m.y., this rate 

would result in about 2.9 km of uplift accommodated across 

the Lexington fault zone. This much uplift, however, is likely 

a minimum because deformation of the Santa Clara Formation 

appears to have been focused between 4.8 and 0.25 Ma. If the 

high uplift rates calculated from fission-track data near the San 

Andreas fault also apply here, then the amount of uplift could 

be as much as 3.8 km. Similarly, shortening accommodated 

across the Lexington fault zone since 4.8 Ma may be 1.1 to 

1.4 km, and reverse slip has been 3.1 to 4.0 km, on the basis of 

minimum and maximum uplift rates, respectively. Kinematic 

data from the Lexington fault zone (Sebrier and others, 1992) 

confirm that the principal early slip occurred no earlier than 

4.8 Ma within the fault zone and that it was reverse slip. On 

the basis of kinematic indicators on fault surfaces, a distinctly 

later episode of dextral slip has been superposed on the reverse 

slip (Sebrier and others, 1992).

MODELING 

THE STRUCTURAL EVOLUTION

OF THE LOMA PRIETA REGION

STRUCTURAL AND STRATIGRAPHIC 

CONSTRAINTS

The areas to the northeast and southwest of the principal 

trace of the San Andreas fault in the Loma Prieta region 

were as much as 150 km apart when movement on the San 

Andreas fault was initiated during the early Miocene (20 Ma). 

Stratigraphic units of the two fault blocks, especially those 

that were deposited before 20 Ma, therefore, have different 

stratigraphic and structural histories.

The structural and stratigraphic data suggest that pre-Tertiary plutonic and metamorphic basement rocks southwest of 

the San Andreas fault underwent 13 to 15 km of unroofing 

between 95 and 61 Ma (fig. 3). Additional uplift (≤4 km) and 

erosion southwest of the San Andreas fault occurred during 

the Oligocene to early Miocene in association with major 

dip slip on the Zayante fault and with volcanism. Transpressional deformation (see pl. 2; fig. 3) could have been active 

southwest of the San Andreas fault as early as the Oligocene, 

although widespread volcanic rocks of this age suggest that 

extensional tectonism was dominant during the Oligocene to 

early Miocene (Stanley, 1987).

Northeast of the San Andreas fault, Franciscan rocks 

underwent unroofing during the Paleocene (65–56 Ma). This 

deformation may have been accompanied by extension that 

continued into the Oligocene, as indicated by widespread 

low-angle attenuation faults in rocks of Eocene age. Transpressional tectonism has been active northeast of the San Andreas 

fault zone only since the early Miocene; most transpression 

occurred during the Pliocene to middle Pleistocene (4.8–0.3 

Ma). Since the Miocene, a cumulative uplift of about 6 km and 

horizontal shortening by about 5 km have occurred between 

the San Andreas fault and the southwestern Santa Clara 

Valley thrust belt. Oligocene and younger transpressional 

deformation across the San Andreas fault zone apparently 

occurred largely in conjunction with the 460 km of Miocene 

and younger aggregate right slip across faults of the modern 

San Andreas fault system (fig. 5; table 2).

Late Cretaceous and early Cenozoic unroofing and attenuation of rocks along the length of the Coast Ranges were taken 

up by low-angle detachment faults and accompanying folds 

that predated the San Andreas fault system. These attenuation faults probably were driven at greater depths by thrust 

faults along the Late Cretaceous to early Cenozoic convergent 

margin between the North American and Farallon plates. These 

low-dipping early Cenozoic structures strongly influenced 

the orientation of later compressional structures, and some 

of the early faults probably have been reactivated during the 

Neogene and Quaternary as compressional low-angle faults 

with accompanying folds.

DEEP-CRUSTAL CONSTRAINTS

Deep seismic reflection and refraction data from recent 

surveys in central and northern California (Wentworth and 

others, 1984, 1987), first-motion studies of earthquakes 

(Oppenheimer and others, 1988), and borehole-breakout 

data (Mount and Suppe, 1987) indicate that the northern 

and central Coast Ranges are presently being deformed 
PACIFIC (+FARALLON?) OCEANIC PLATE

(Moho - Page and Brocher, 1993)

LITHOSPHERIC MANTLE

(Fuis and Mooney, 1990)

ASTHENOSPHERE ASTHENOSPHERE

UNDERPLATED

CRUST IN OLDER

SLAB WINDOW

(Furlong and others, 1989)

TOP OF SLAB WINDOW

LITHOSPHERIC MANTLE

CRYSTALLINE (SIERRAN ?) BASEMENT

BRITTLE DEFORMATION

(DUCTILE

DEFORMATION)

(DUCTILE

DEFORMATION)

sea level

sea level

sea level PACIFIC PLATE SGF shoreline ZAY

SAFSAR

CAL

Diablo Range Great Valley TESLA-ORT

~120 km

Santa

Cruz

Ben

Lomond

Mountain

Loma

Prieta

Peak

Mount Hamilton

Los Gatos

Santa

ClaraValley

~90 km

San Andreas fault to tip ofCoast Range Wedge

SEA LEVEL

5

10

15

2030

35

40

4550

25

55

Contintental Slope to San Andreas fault

DEPTH (KM)

(Topography and elevations above sea level schematic)

GABBRO OF LOGAN

(Furlong and others, 1989)

?

?

?

? ?

?

??

?

?

?

?

?

?

?

?

?

?

? ?

?

?

? ? ? ? ? ?

?

?

?

?

? ? ?

? ? ?

?

?

?

?

S(F(O(S(C(PpPp(MUs1mMM

?

Figure 8.—Schematic cross section across the California margin at latitude of Loma Prieta (fig. 1), showing hypothetical deep structure of the San Anwedging, and plate boundary relations. Depth, thickness, and compositions of crust and mantle units and location of midcrustal decollement are preflection and refraction models of Fuis and Mooney (1990), Page and Brocher (1993), and Brocher and others (this chapter). Depth to present top and Snyder, 1979), configuration of lithified materials underplated in older, shallower roof area of window, and hypothetical boundary relation beAmerican plates are based on thermal and seismic models of Furlong and others (1989). CAL, Calaveras fault; SAF, San Andreas fault; SAR, Sargefault; TESLA–ORT, Tesla-Ortigalita fault; ZAY, Zayante fault.
E40 EARTHQUAKE OCCURRENCE

by compression normal to the San Andreas fault (Zoback 

and others, 1987). Interpretations of the seismic reflection 

and refraction data suggest that significant compression is 

accommodated by blind thrust faults within an eastwardly 

tapered, composite series of tectonic wedges which are herein 

referred to as the “wedge” or wedge complex. The “wedge” 

is thought to consist largely of the Franciscan Complex and 

to have been driven northeastward into the lower part of the 

Great Valley sequence and the Coast Range ophiolite during 

the Late Cretaceous to early Tertiary. The wedge complex is 

also considered to be presently active under the eastern Coast 

Ranges (Wentworth and others, 1984; Wentworth and Zoback, 

1990; Unruh and Moores, 1992). Roof thrusts of the “wedge” 

dip east near its tip and root into a master zone of west-dipping 

floor thrusts (fig. 8).

A major unresolved uncertainty is the question of how major 

strike-slip faults in the San Andreas fault system interact with 

the basal floor-thrust zone of the wedge complex. Details of 

the deformation history of the wedge complex also remain 

obscure. Nevertheless, such a model provides a reasonable 

mechanism to unify relations among strike-slip, uplift, and 

crustal shortening in the Coast Ranges.

WEDGE TECTONIC MODEL

A hypothetical wedge tectonic model for the Loma Prieta 

region and adjacent San Francisco Bay region is illustrated in 

figure 8. In this model, faults of the southwestern Santa Clara 

Valley thrust belt represent a part of the roof thrust zone of the 

wedge complex. The southwestward dip of the thrust belt is 

explained as an effect of southwestward tilting and northeastsouthwestward shortening in the roof-thrust zone of the wedge 

complex, resulting from eastward “wedge” propagation.

The geometry of the “wedge” tip, the thickness of the 

“wedge,” and the depth to its tip are based on reflection and 

refraction profiling (Fuis and Mooney, 1990; Wentworth and 

Zoback, 1990). These models suggest that the tip of the wedge 

complex is about 90 km east of the San Andreas fault, beneath 

the San Joaquin Valley. The 5- to 6-km/s-velocity material 

that is assumed to be Franciscan Complex, which composes 

the bulk of the “wedge,” is interpreted to be underlain below 

its floor-thrust zone by higher-velocity mafic oceanic rocks, 

including rocks correlative with the Coast Range ophiolite 

and (or) the gabbro of Logan. From the seismic refraction 

and reflection modeling (Fuis and Mooney, 1990), we also 

infer that rocks of the Jurassic western Sierra Nevada slate 

belt continue in the subsurface from their outcrops east of the 

Great Valley to the San Andreas fault, where they underlie 

the Sierra Azul block and rocks of the Coast Range ophiolite 

at 12- to 15-km depth. The slate and phyllite of Loma Prieta 

peak and overlying ophiolitic section in the hanging-wall block 

of the Sargent fault zone at Loma Prieta peak are interpreted 

in this model to be a part of this sub-Franciscan basement. 

Aeromagnetic data (see Jachens and Griscom, this chapter) 

and structure sections (see pl. 1), however, constrain the total 

vertical displacement of the Coast Range ophiolite by the 

Sargent fault to less than 2 km. Uplift of the slate and phyllite 

of Loma Prieta peak and overlying ophiolitic section from 

greater depths, therefore, must have occurred along another, 

possibly flatter fault beneath the Sierra Azul block.

Further support for this interpretation is indicated by the 

widespread occurrence of laumontite in the Cretaceous and 

older section of the Sierra Azul block but its absence in 

younger rocks, and by the occurrence in the Eocene sequence 

of the Sierra Azul area of detritus derived from high-pressuremetamorphic rocks of the Franciscan Complex. The distribution of laumontite and of high-pressure mineral detritus in the 

section requires burial of the Cretaceous and older rocks to 

3- to 11-km depth (Dickinson and others, 1969; Blake and 

others, 1988), followed at 70–56 Ma by 8 to 9 km of unroofing, 

to raise the eroded top of the Cretaceous and older section to a 

depth shallower than the stability field of laumontite. Attenuation faulting, which downfaulted the unlaumontitized Eocene 

sequence onto various pre-Eocene rock units, appears to have 

occurred as late as 30 Ma but also earlier, during unroofing of 

the pre-Eocene section (Jayko and others, 1987).

Interpreted within the context of a wedge tectonic model, 

the late Mesozoic and Cenozoic uplift pattern was related to 

oblique northeastward emplacement of a wedge complex, 

resulting in 13 to 15 km of uplift by 61 Ma and an additional 8 

to 9 km of uplift by 56 Ma. Speculatively, the slate and phyllite 

of Loma Prieta peak was uplifted along flat thrusts in the roof 

of the wedge complex from a sub-Franciscan basement position to within 4 km of the surface during this pre-early Eocene 

stage of “wedge” formation. Concurrent with northeastward 

“wedge” propagation, folding and tilting occurred above 

blind thrusts in the roof- and floor-thrust zones of the wedge 

complex, beneath the shallower attenuation structures.

Movement on the San Andreas transform fault was initiated 

about 18 Ma as the Mendocino Fracture Zone propagated 

to the north end of the Salinia terrane. Concurrently, hot 

asthenosphere moved upward into the lower crust and upper 

mantle (fig. 8) beneath the North American plate south of the 

Mendocino triple junction, to form a slab window in place 

of the subducted Farallon plate north of the triple junction 

(Dickinson and Snyder, 1979). At that time, the La Honda and 

Ben Lomond blocks were about 300 km to the south, and the 

San Francisco Bay block was about 170 km south of its present position. The northeastward-vergent Sargent and Berrocal 

faults, as well as faults of the southwestern Santa Clara Valley 

thrust belt, may also have formed at this time in response to 

transpression across the propagating north end of the triple 

junction. Surface fault offsets, Quaternary terrace-uplift 

rates, and fission-track data (Bürgmann and others, 1994) 

northeast of the San Andreas fault suggest that cumulative 

uplift of about 6 km and shortening of about 5 km have been 

accommodated by these thrusts in the roof-thrust zone of the 

wedge complex. In the Loma Prieta region (see pl. 1), these 

thrusts root in the main San Andreas transform boundary at 
E41

8- to 10-km depth and activate parts of the buried Oligocene 

and older roof-thrust zone of the wedge complex. Southwest 

of the San Andreas fault, fission-track data (Bürgmann and 

others, 1994), stratigraphy, and paleobathymetry suggest as 

much as 3 km of uplift since 30–24 Ma.

From at least 18 Ma to the present, dextral slip along the 

Pilarcitos fault, faults of the east-bay fault system, and the 

Peninsular segment of the San Andreas fault have broken 

the “wedge” into the Salinia terrane, the San Francisco Bay 

block, and the eastern San Francisco Bay block. Strike-slip 

motion between the San Andreas and east-bay fault systems 

may partly be accommodated by oblique slip along the floorthrust zone of the “wedge.” An undetermined, but probably 

small, component of the 5 km of horizontal shortening and 

8 km of uplift west of the east-bay fault system also may be 

accommodated by movement along reactivated parts of this 

floor-thrust zone.

The Pacific-North American plate boundary has migrated 

eastward to the east-bay fault system (that is, the Calaveras, 

Hayward, and related faults) since the late Miocene. This 

eastward step in the plate boundary was facilitated by underplating and cooling of the lithospheric mantle and lower crust 

above the slab window formed south of the Mendocino triple 

junction (fig. 8), as the transform and locus of asthenospheric 

upwelling propagated northward and the plate boundary 

became increasingly transpressional (Furlong and others, 

1989).

1989 LOMA PRIETA EARTHQUAKE

The main shock of the 1989 Loma Prieta earthquake was 

initiated southwest of the San Andreas fault, with rupture 

occurring from 17- to 8-km depth, along a 70° SW.-dipping 

blind dextral-reverse fault (U.S. Geological Survey staff, 

1990). The faulting was initiated near the base of the North 

American plate (figs. 4, 8), near the floor-thrust zone of the 

wedge complex. Seismicity above 10-km depth is considered 

to have resulted from accommodation of the deeper slip by 

upwarping and secondary slip on the thrust-fault system 

northeast of the San Andreas fault. Beneath the Sargent fault, 

secondary slip was taken up to within 4 km of the surface.

Within the context of a wedge tectonic model, the 1989 

Loma Prieta earthquake may be interpreted as resulting from 

activation of the hypothetical floor-thrust zone of the wedge 

complex, in response to transpression across the Pacific-North 

American plate boundary. The earthquake rupture propagated 

steeply upward and northeastward from the floor thrust to 

8–10-km depth. At shallower levels to the northeast, distributed secondary slip was accommodated across eastwardvergent oblique-dextral thrusts of the southwestern Santa 

Clara Valley thrust belt. Southwest of the principal bedrock 

trace of the San Andreas fault, secondary slip was accommodated by extensional faults in the Summit Road-Skyland 

Ridge area that appear to have ruptured from the surface 

downward to less than 200-m depth and that may not have 

roots in the deep crust.

CONCLUSIONS

1. The Loma Prieta earthquake of October 17, 1989 occurred 

on a steep, southwest-dipping, right-lateral reverse fault that 

did not rupture to the surface. The fault rupture was initiated 

southwest of the principal trace of the San Andreas fault at 

about 17-km depth, near or at the base of the continental 

crust and, possibly, along the floor-thrust zone of a tectonic 

wedge complex. The primary rupture propagated upward 

to 8- to 10-km depth beneath the principal surface trace 

of the San Andreas fault. Whether this primary rupture 

was on the principal trace of the San Andreas fault or on 

a subsidiary fault that merges upward with a vertical San 

Andreas zone is still undetermined.

2. Geophysical modeling (see Brocher and others, this chapter; Jachens and Griscom, this chapter) indicates that the 

primary rupture occurred entirely within buried crystalline 

basement rocks beneath a slab of magnetic rocks, as much 

as 5 km thick, believed to represent the subsurface continuation of the gabbro of Logan.

3. Above the 8- to 10-km depth of primary rupture, slip was 

accommodated southwest of the principal trace of the San 

Andreas fault (a) by upwarping in the hanging-wall block 

of the fault; (b) by secondary shallow extension, sinistral 

slip, and associated ridgetop spreading in the Summit 

Road-Skyland Ridge fault zone; and (c) by shaking-induced 

massive slope failure. Most of the secondary surface deformation during the earthquake was focused in the Summit 

Road-Skyland Ridge fault zone, off and southwest of the 

principal surface trace of the San Andreas fault. Geologic 

data indicate that similar deformation occurred in the 

Summit Road-Skyland Ridge area both in 1906 and earlier 

in the Quaternary, and so it should be expected to recur here 

during earthquakes of similar magnitude in the future.

4. The Sargent fault immediately northeast of the principal 

trace of the San Andreas fault shows evidence of major 

reverse slip that occurred from about 18–10 Ma. A late 

phase of right-lateral slip indicated from slip-surface data 

may be later than 10 Ma. During the 1989 Loma Prieta 

earthquake, a small component of right-lateral extension 

and compression occurred near the northwest end of the 

Sargent fault near Lake Elsman.

5. Fission-track data (Bürgmann and others, 1994) suggest 

that about 3 to 4 km of uplift have occurred beneath the 

Sierra Azul block northeast of the San Andreas fault since 

3 Ma, indicating a post-3-Ma uplift rate of about 1 mm/yr. 

Constraints on the timing and duration of movement on the 

Sargent fault and on the southwestern Santa Clara Valley 

thrust belt suggest that uplift since 3 Ma has been largely 

accommodated by the southwestern Santa Clara Valley 

thrust belt and that most uplift and shortening occurred 

STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E42 EARTHQUAKE OCCURRENCE

before about 250 ka. Speculatively, younger blind thrusts 

could be present beneath the Santa Clara Valley that root 

southwestward into the exposed southwestern Santa Clara 

Valley thrust belt or the San Andreas fault. Other southwestward-vergent blind thrusts might also exist beneath the 

Santa Clara Valley that merge at depth with the Calaveras 

fault or a more eastward segment of the floor- or roofthrust zone of the active tectonic wedge complex believed 

to underlie this region.

ACKNOWLEDGMENTS

We are particularly grateful to Earl Brabb, who participated 

in parts of the mapping and provided much unpublished 

biostratigraphic information. We thank Robert Jachens and 

Andrew Griscom for extensive discussions and insights into 

the implications of their geophysical studies of crustal structure, and Robert Jachens and Carl Wentworth for their reviews 

of the manuscript. Zenon Valin prepared digital files for many 

of the figures and for the geologic map and structure sections 

(see pl. 1). Some of the illustrations were originally prepared 

by Kathryn Nimz. Finally, we are particularly indebted to Carl 

Wentworth for the many hours he spent helping the first author 

edit and simplify the digital map (see pl. 1).

REFERENCES CITED

Adam, D.P., McLaughlin, R.J., Sorg, D.H., Adams, D.B., Forester, 

R.M., and Repenning, C.A., 1982, An animal and plant-fossil 

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STRATIGRAPHY AND STRUCTURE ACROSS THE SAN ANDREAS FAULT ZONE
E48 EARTHQUAKE OCCURRENCE
E49

CONTENTS

————— Page

Abstract------------------------------------------------------------------------- E49

Introduction-------------------------------------------------------------------- 50 

Data sources------------------------------------------------------------------- 50

 Geology------------------------------------------------------------------ 50

 Magnetic field----------------------------------------------------------- 50 

 Gravity field------------------------------------------------------------- 51

 Physical properties----------------------------------------------------- 51

Basement blocks of the Loma Prieta region------------------------------- 51

 Crust southwest of the San Andreas fault--------------------------- 51

 Block boundaries southwest of the San Andreas fault------------ 56

 Crust northeast of the San Andreas fault---------------------------- 59

 Block boundaries northeast of the San Andreas fault------------- 60

 Three-dimensional structure notheast of the San Andreas fault-- 61

Implications for origin and tectonic evolution---------------------------- 65

 Crust southwest of the San Andreas fault--------------------------- 65

 Crust northeast of the San Andreas fault---------------------------- 67

Crustal structure and seismicity--------------------------------------------- 70

Conclusions-------------------------------------------------------------------- 74

Acknowledgments------------------------------------------------------------ 77

References cited--------------------------------------------------------------- 77

 —————

ABSTRACT

To characterize the geophysical and geologic setting of the 

1989 Loma Prieta earthquake, we have constructed a model 

for the subsurface structure and composition of the midcrust 

and upper crust beneath the Santa Cruz Mountains and 

adjacent southern Santa Clara Valley. This study was based 

mainly on an interpretation of local anomalies in the gravity 

and magnetic fields, relying heavily on geologic information to constrain the gravity and magnetic modeling, and on 

velocity and structural information from seismic tomography. 

We subdivide the study area into seven distinctive basement 

blocks on the basis of geophysical and geologic expression, 

provide a three-dimensional description of each basement 

block and its bounding faults by using calculated models, 

discuss implications about the tectonic development of the 

Loma Prieta region, and analyze the relation between the 

aftershock distribution and our crustal model.

Southwest of the San Andreas fault, from southwest to 

northeast, are three basement blocks: the nonmagnetic Ben 

Lomond block, a terrane of Mesozoic granitic and metamorphic rocks; the Logan block, with a Mesozoic basement of 

magnetic, high-density gabbro that is separated from the Ben 

Lomond block by the Zayante fault, which dips southwest 

at an average angle, in places, of less than 45°; and, to the 

north, another nonmagnetic block, presumably composed 

of granitic rocks, interpreted to be in fault contact with the 

Logan block. The Zayante fault may be a former Mesozoic 

thrust fault reactivated by Cenozoic tectonism. The western parts of these three basement blocks are terminated by 

the San Gregorio fault, which locally dips 50°–70° E. The 

eastern parts of all three blocks are terminated by the San 

Andreas fault. Magnetic Pliocene sedimentary rocks in the 

shallow subsurface are thrust as much as 1 km northeastward 

beneath the rocks on the opposite side of the San Andreas 

fault. To the south, near exposures of the gabbro of Logan, 

the fault is nearly vertical, but a few kilometers farther south 

it dips 70° SW. to at least 6-km depth, where low-density 

Pliocene sedimentary rocks are thrust beneath Mesozoic 

granitic rocks.

Northeast of the San Andreas fault, the area extending to the 

Silver Creek fault is subdivided into four blocks of Franciscan 

Complex that, in turn, are overlain near the two bordering 

faults by the Mesozoic Great Valley sequence (including the 

Coast Range ophiolite) and Tertiary sedimentary rocks. These 

faulted blocks are primarily delineated by curvilinear magnetic boundaries. The two southwestern blocks dip southwest 

and contain thick (2–4 km) slabs of magnetic serpentinite 

(Coast Range ophiolite) separated by northeastward-directed 

tectonic wedges of Franciscan rocks. The two northeastern 

blocks, though deformed in detail, are interpreted from their 

geophysical expression to be regionally horizontal, each block 

underlain by a subhorizontal floor thrust. These subhorizontal 

floor thrusts are interpreted to be wedge faults on the basis of 

their associations with the Coast Range ophiolite.

The geologic units northeast of the San Andreas fault, their 

geophysical expression, and their history and style of tectonic 

wedging are similar to those along the 600-km-long Franciscan-Great Valley contact 40 km east of the study area. We 

suggest that the blocks of Franciscan rocks in the Loma Prieta 

region were originally situated along the Franciscan-Great 

Valley contact near its present truncation by the San Andreas 

fault (at lat approx 36° N.) and were offset northwestward 

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE,

INFERRED FROM GRAVITY AND MAGNETIC ANOMALIES

By Robert C. Jachens and Andrew Griscom, U.S. Geological Survey

E49
E50 EARTHQUAKE OCCURRENCE

during post-middle Miocene time along the San Andreas fault 

system to their present positions.

Comparison of our crustal model with the aftershock distribution indicates that nearly all hypocenters can be assigned to 

only three major faults: the San Andreas, Zayante, and Sargent 

faults. Many aftershocks whose epicenters plot away from the 

traces of these faults are associated with dipping fault segments. In particular, the San Andreas fault surface is locally 

curved, dipping gently northeast in the uppermost part of the 

crust and then bending back steeply southwest in the deeper 

parts. Epicenters of aftershocks on the curved segments of the 

San Andreas fault plot on both sides of its trace. Hypocenters 

appear to avoid our modeled configuration of the gabbro of 

Logan, but the mechanism by which this basement block 

influences the regional seismicity is unknown.

INTRODUCTION1

The 1989 Loma Prieta earthquake and its aftershocks 

occurred in a region of marked crustal complexity and diversity. The San Andreas fault itself marks a first-order break 

between a moderately deformed Mesozoic plutonic terrane 

(Salinia terrane) to the southwest and a highly deformed complex of accretion-related terranes (Franciscan Complex) and 

the Great Valley sequence to the northeast. Past movements 

on other major faults on both sides of the San Andreas fault 

have juxtaposed numerous crustal blocks with contrasting 

histories and properties. The crust that has resulted from 

assembly of these diverse blocks is characterized by profound 

variations in composition and physical properties, both laterally and vertically.

In a broad sense, the 1989 Loma Prieta earthquake and its 

aftershocks are a manifestation of continuing movement on 

the San Andreas fault system in the Santa Cruz Mountains. In 

detail, however, the specific relation of these events to the San 

Andreas fault proper is not so clear. No throughgoing pattern 

of surface rupture on the San Andreas fault was observed after 

the earthquake, and the aftershock distribution shows strong 

planar alignments that in some places lie well off the presumed 

subsurface fault position (Dietz and Ellsworth, 1990).

Because of the discrepancies noted above, questions have 

been raised concerning which fault actually ruptured during 

the earthquake and what role crustal structure may have played 

in localizing or modifying the rupture pattern. To address these 

questions and to provide a general picture of the geophysical 

and geologic setting of the earthquake, we have constructed 

a model for the subsurface structure and composition of the 

midcrust and upper crust beneath the Santa Cruz Mountains 

and adjacent southern Santa Clara Valley. This study was 

based mainly on an interpretation of local anomalies in the 

gravity and magnetic fields, relying heavily on the geologic 

information of McLaughlin and Clark (this chapter) and R.G. 

Coleman (written commun., 1991) to constrain our modeling. 

We also used some velocity and structural information from 

seismic tomography (Foxall and others, 1993). Our model by 

no means completely characterizes the midcrust and upper 

crust in the Loma Prieta region because it focuses mostly on 

those features that are reflected in the gravity and magnetic 

fields. Everywhere, however, we have attempted to link our 

crustal model directly to the surface geology and, in so doing, 

have projected the geology into the subsurface rather than 

simply inferring physical properties. In addition to providing 

structural information about the midcrust and upper crust, the 

resulting image of the subsurface carries important implications previously unrecognized about the tectonic development 

of the Loma Prieta region, as discussed below.

DATA SOURCES

GEOLOGY

The geologic maps (pls. 3A, 3B) were compiled principally 

from the data of Wentworth and Blake (1991; C.M. Wentworth 

and M.C. Blake, Jr., written commun., 1991) and from 7 1/2-

minute quadrangle maps (McLaughlin and others, 1988b, 

1991; Clark and others, 1989). Additional information was 

compiled from the geologic map of California (Jennings and 

Strand, 1958; Jennings and Burnett, 1961; Rogers, 1966) and 

from the reports by Clark and Rietman (1973), McCulloch and 

Greene (1990), and Wagner and others (1991). Information 

on the terranes consisting of Franciscan Complex was provided by R.J. McLaughlin (written commun., 1992). We have 

considerably simplified these geologic maps by generalizing 

the stratigraphy into only a few units that are relevant to our 

geophysical interpretation, by eliminating many faults deemed 

to be relatively minor, and by enlarging certain serpentinite 

masses to sizes that can be depicted at these map scales.

MAGNETIC FIELD

The primary source of magnetic-field data for this study is a 

detailed aeromagnetic survey of a 25-km-wide transect across 

the Santa Cruz Mountains, the southern Santa Clara Valley, 

and the southern Diablo Range flown by the U.S. Geological 

Survey (USGS) (Abrams and others, 1991) during late 1989 

and early 1990 (see pl. 3A). Total-magnetic-field data were 

collected with a fixed-wing aircraft along flightlines (N. 75° 

W.) spaced 400 m apart at a nominal height of 150 m above 

the ground. Actual flight heights varied, depending on the 

ruggedness of the topography and the local population density 

(nominal flight height, 300 m over populated areas). This aeromagnetic survey covers the epicentral area of the 1989 Loma 

————— 1This study and the composition of the manuscript were completed in 

November 1993, and so this paper uses only the data and cited references 

that were available at that time.
E51

Prieta earthquake and its aftershocks. Detailed profiles were 

extracted from the aeromagnetic survey to serve as the basis 

for most of the magnetic modeling discussed below.

Additional magnetic-field data of generally lower quality 

and less detail were assembled to provide a regional geophysical context for the detailed aeromagnetic survey described 

above. These data, which were from aeromagnetic surveys 

with a wide variety of flight heights and flightline spacings, 

were analytically merged into a consistent single data set that 

represents the magnetic field at 300 m above the ground (see 

pl. 3B). Jachens and Roberts (1993) provided details of the 

aeromagnetic surveys and merging procedures used to produce 

the regional magnetic map (pl. 3B).

GRAVITY FIELD

The regional gravity map of the southern San Francisco Bay 

region (pl. 3C), including the Loma Prieta region, was constructed mainly from data compiled for the isostatic-residualgravity map of California and offshore southern California 

(Roberts and others, 1990), supplemented by more than 

1,100 new observations on the San Francisco peninsula and 

in the mountains surrounding the southern San Francisco Bay 

region. More complete descriptions of the original data and 

their sources were given by Clark and Rietman (1973) or can 

be traced through the references listed on the following sheets 

of the California Division of Mines and Geology’s “Bouguer 

Gravity Map of California”: San Francisco sheet (Chapman 

and Bishop, 1968), San Jose sheet (Robbins and others, 1976), 

and Santa Cruz sheet (Bishop and Chapman, 1967).

The data are plotted on plate 3C as isostatic-residual-gravity values to emphasize those anomalies caused by density 

distributions in the midcrust and upper crust (uppermost 

approx 10–15 km) at the expense of anomalies from deeper 

density distributions that buoyantly support the topography 

in a manner consistent with the concept of isostasy (Jachens 

and Griscom, 1985). Standard procedures and formulas (for 

example, Telford and others, 1976) were used to reduce the 

observed gravity data to complete Bouguer gravity values, 

and an Airy-Heiskanen model of isostatic compensation 

(Heiskanen and Moritz, 1967) was used to further reduce the 

data to isostatic-residual-gravity values. Details of the gravity 

reduction procedures applied to the data were given by Roberts 

and Jachens (1993).

PHYSICAL PROPERTIES

Information on the densities and magnetic susceptibilities 

of the major lithologic units within the study area (see pls. 

3A, 3B) was obtained mainly from measurements on hand 

specimens (table 1), although in a few places magnetic-susceptibility measurements were made directly on outcrops 

with portable susceptibility meters. The data of Ross (1972) 

on Mesozoic granitic rocks, of Bailey and others (1964) on 

sandstone and metasandstone of the Great Valley sequence 

and the Franciscan Complex, and of Brabb (1960) on Tertiary 

sedimentary rocks of the Santa Cruz Mountains, together with 

about 200 measurements by V.E. Langenheim (unpub. data, 

1992) and additional measurements on about 500 samples 

from the collections of various U.S. Geological Survey geologists, constitute the 1,269-element physical-property data base 

available to this study. Summaries of density data relevant to 

this study were published by Chapman and Bishop (1968), 

Clark and Rietman (1973), Robbins and others (1976), and 

Carle and Langenheim (1990).

BASEMENT BLOCKS OF

THE LOMA PRIETA REGION

A convenient way of characterizing the crustal structure 

of the Loma Prieta region is in terms of distinct basement 

blocks, defined on the basis of their geophysical and geologic 

expressions. At least seven such basement blocks occur in 

the study area (see pl. 3A). These basement blocks are not 

necessarily tectonostratigraphic terranes according to the 

standard definition (Howell and others, 1985), although some 

of them satisfy that definition and encompass such terranes. 

Many of the basement blocks are bounded by faults or probable faults, but others have bounding structures whose nature 

remains unclear.

CRUST SOUTHWEST OF

THE SAN ANDREAS FAULT

Three major basement blocks make up the crust southwest 

of the San Andreas fault: (1) the Ben Lomond block (see 

McLaughlin and Clark, this chapter), bounded on the west by 

the San Gregorio fault and on the northeast by the Zayante 

fault and its southwestward continuation, the Vergeles fault 

(hereinafter referred to simply as the Zayante fault) and floored 

by Late Cretaceous granitic rocks that are exposed in the Ben 

Lomond pluton and the northern Gabilan Range (see pl. 3A; 

Clark and Rietman, 1973; Ross, 1984); (2) another block in the 

northern part of the study area, floored by unknown basement 

that might be a southward continuation of Late Cretaceous 

granitic rocks of the Montara pluton exposed on the San 

Francisco peninsula; and (3) the Logan block, an intervening 

block whose basement is almost totally concealed beneath 

Cenozoic sedimentary rocks and deposits but whose geophysical signature implies a primarily mafic igneous composition 

(figs. 1–3). The second and third blocks compose the La Honda 

block of McLaughlin and Clark (this chapter).

The basement of the Ben Lomond block, where exposed in 

the Ben Lomond pluton and the northern Gabilan Range (see 

pl. 3B), is composed of minor amounts of metasedimentary 

rocks intruded by Late Cretaceous plutons varying widely in 

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E52 EARTHQUAKE OCCURRENCE

Figure 1.—Isostatic-residual-gravity profile (A) (circles, measured data; curve, calculated values) and density 

model (B) along line R–R' (pl. 3) across the Logan block between the Zayante and San Andreas faults. Most 

reliable part of density model is between km 0 and 25. Pluses, locations of gravity observations. Inverted 

triangles, gravity stations on granitic basement. Numbers on units are densities (in grams per cubic centimeter). 

Vertical density structure for Tertiary sedimentary rocks here and in figures 3, 5, 9, and 16 was determined by 

using variation in seismic velocity with depth (Foxall and others, 1993), converting to a variation of density 

with depth (Gardner and others, 1974) after adjusting resulting curve on the basis of densities of samples 

collected from surface outcrops and from cores from base of a 3.4-km-deep drill hole. Most important features 

of density model include the contorted Zayante fault, the vertical San Andreas fault, and a wedge-shaped mass 

of high-density (2.91 g/cm3) gabbro (gb) below basin of low-density Tertiary sedimentary rocks between these 

two faults. High-density slab at base of model is included to account for high background gravity values over 

the Salinia terrane; actual density distribution responsible for this high background level may be a thin gabbroic 

body or may include entire lower crust. No vertical exaggeration. gr, granitic and older metamorphic rocks; 

Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range ophiolite; KJf, Cretaceous and Jurassic 

Franciscan complex; KJs, Cretaceous and Jurassic sedimentary rocks; Qal, Quaternary alluvium; Ts, Tertiary 

sedimentary rocks. -10 0 10 20 GRAVITY ANOMALY, IN MILLIGALS

0 5 10 15 20 25 30

15 10 5 0

DEPTH, 

IN KILOMETERS

DISTANCE, IN KILOMETERS

gb 2.91

Ts 2.67

gr 2.67

Ts 2.59

Ts 2.56

Ts 2.52

Ts 2.48

Ts 2.37

Qal 2.17

Ts+KJs+KJf

2.60

Jou

2.77

KJf 2.67

2.87

gr? 2.67

Jou 2.72

Zayante fault

San Andreas fault

Sargent fault

R R'

A

B

gb?
E53

composition, mainly quartz diorite in the Ben Lomond pluton 

and the northernmost Gabilan Range but more felsic farther 

south in the Gabilan Range (Ross, 1972, 1984; Clark and 

Rietman, 1973). Similar crystalline basement rocks penetrated 

by drill holes in the onshore area between the Ben Lomond 

pluton and the northern Gabilan Range (Clark and Rietman, 

1973) are evidence for the continuity of this basement block in 

the subsurface between the two outcrop areas. The crystalline 

rocks of the Ben Lomond block are effectively nonmagnetic 

and produce no magnetic anomalies (see pl. 3B). The basement is overlain by a thin section of Cenozoic strata (typically 

<1 km thick), mainly of middle Miocene age and younger 

except along the north edge of the block, where the Butano 

Sandstone of Eocene age is exposed (Clark and Rietman, 

1973; Wagner and others, 1991).

The Logan block is bounded at the surface on the northeast 

by the San Andreas fault, on the west by the San Gregorio fault, 

and on the southwest by the Zayante fault. Its north boundary 

is covered by Tertiary sedimentary rocks and defined by a 

magnetic gradient (see pl. 3B; fig. 2) but is difficult to locate 

precisely. The Logan block is not clearly associated with any 

major exposed structures. Its basement is exposed only as a 

thin sliver along the San Andreas fault near the southeast end 

of the block (see pls. 3B, 3C) where it is composed primarily of 

-200 0 200

MAGNETIC ANOMALY,

IN NANOTESLAS

0 5 10 15 20 25 30 35 40

DISTANCE, IN KILOMETERS

15 10 5 0

DEPTH,

IN KILOMETERS

gr

gr? gr?

KJf

gb 0.0024

Ts

?

Zayante fault

San Andreas fault

Q Q'

A

B

Figure 2.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) along line Q–Q' (pl. 3) across 

the Logan block between the Zayante and San Andreas faults. Pluses, locations of magnetic observations. In figure 2B, rectangular body of 

magnetic gabbro (gb) underlies sedimentary basin bounded by the Zayante and San Andreas faults; number is magnetization (in electromagnetic 

units per cubic centimeter). Anomaly between km 24 and 29 is caused by an ophiolitic body bounded by magnetic boundary 1A and is 

located northeast of the San Andreas fault. No vertical exaggeration. gr, granitic and older metamorphic rocks; KJf, Cretaceous and Jurassic 

Franciscan complex; Ts, Tertiary sedimentary rocks.

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E54 EARTHQUAKE OCCURRENCE

two unusual plutonic rock types: a hornblende-quartz gabbro 

and a less common, foliated anorthositic gabbro (Ross, 1970, 

1972). According to Ross (1970, 1972), these rocks are unlike 

any other plutonic rocks in the Coast Ranges southwest of the 

San Andreas fault (Salinia terrane). A U-Pb zircon age of 161 

Ma for these gabbros (James and others, 1986), significantly 

older than the Late Cretaceous ages typical of granitoids of 

the Salinia terrane, is additional evidence for the unusual composition of these rocks. Both types of gabbro are extremely 

dense relative to other plutonic rocks of the Salinia terrane 

(table 1). Although the hornblende-quartz gabbro also is highly 

magnetic (as implied by the strong magnetic anomaly over its 

outcrop area), the anorthositic gabbro appears to be, at most, 

only weakly magnetic, on the basis of the absence of a strong 

magnetic anomaly over the southern part of its outcrop area. 

The magnetic properties inferred from the magnetic map (pls. 

3A, 3B) are supported by magnetic-susceptibility measurements on a few samples of both types of gabbro (table 1).

The highly magnetic properties of the hornblende-quartz 

gabbro relative to other plutonic and sedimentary rocks in 

the study area (see pls. 3A, 3B) permit this body to be traced 

by its magnetic anomaly even where concealed beneath many 

-40 -20 0 20 40

GRAVITY ANOMALY,

IN MILLIGALS

0 5 10 15 20 25 30 35 40

15 10 5 0

DEPTH,

IN KILOMETERS

Ts 2.37

Ts 2.48

 Ts 2.55

Ts 2.61

gr 2.67 Ts 2.64

gb 2.91

KJf 2.68

Qal+Ts 2.17

(schematic only)

Qal 2.04

2.87

gr? 2.67

San Andreas fault

DISTANCE, IN KILOMETERS gr? 2.67

gb?

?Zayante fault

Q Q'

A

B

Figure 3.—Isostatic-residual-gravity profile (A) (circles, measured data sampled from gridded gravity map; curve, calculated values) and 

density model (B) along line Q–Q' (pl. 3) across the Logan block between the Zayante and San Andreas faults. Pluses, locations where 

gravity was calculated. Numbers on units are densities (in grams per cubic centimeter). Density model is consistent with magnetic model 

(fig. 2B) and with sedimentary-rock densities modeled in figure 1. See figure 1 for explanation of high-density slab at base of density model. 

The Zayante fault does not cause a gravity gradient, because of young thin sedimentary basin shown on left side of fault. In figure 3A, thin 

curve is gravity calculated from density model with sedimentary section stripped away—that is, with all density contrasts in sedimentary 

section set to zero. No vertical exaggeration. gb, gabbro; gr, granitic and older metamorphic rocks; KJf, Cretaceous and Jurassic Franciscan 

complex; Qal, Quaternary alluvium; Ts, Tertiary sedimentary rocks.
E55

kilometers of nonmagnetic material. Modeling of the magnetic data along line S–S' across the gabbro outcrop (fig. 4) 

indicates that, as Ross (1972) speculated, the outcrop is the 

top of a thin (1–2 km wide) sliver within the San Andreas 

fault zone. Our model also shows, however, that this thin 

sliver connects at depth with a more extensive, flat-lying 

gabbro body (in places nearly 5 km thick) that makes up, 

in part, the basement of the Logan block. Similar modeling 

and inferences based directly on the magnetic map (pl. 3B) 

indicate that the magnetic gabbro extends in the subsurface 

westward from the San Andreas fault to approximately the 

location of the major magnetic boundary (dashed line, pl. 

3B). Note that this magnetic boundary represents the inferred 

minimum westward extent of the hornblende-quartz gabbro 

and, as such, indicates the minimum westward extent of the 

Logan block in the subsurface. Nonmagnetic basement rocks 

or deeper magnetic basement rocks of the Logan block could 

extend beyond this limit.

All the magnetic models of the basement of the Logan block 

indicate that the magnetic gabbro of Logan is a slablike body 

which does not extend below about 10-km depth. The crust 

immediately below the gabbro is significantly less magnetic, 

has a density that is poorly constrained by gravity data (most 

likely avg >2.67 g/cm3), and has a P-wave velocity of more 

Table 1.—Densities and magnetic susceptibilities of rock samples collected in the San Francisco Bay area, Calif.

[Do., ditto]

Magnetic susceptibility Density (g/cm3

) Number (10−3 cgs units)

Rock type of

samples

Mean±1σ Median Mean±1σ Median

Franciscan Complex

Chert ---------------------------------------- 13 2.62±0.06 2.63 0.0 0.0

Greenstone--------------------------------- 29 2.88±0.09 2.90 .1±0.2 . 1

Limestone---------------------------------- 7 2.66±0.05 2.67 . 0 . 0

Sandstone and metasandstone---------- 448 2.66±0.11 2.66 – – – – – –

Do-------------------------------------- 75 – – – – – – .0±0.0 . 0

Salinia terrane

Ben Lomond Mountain pluton --------- 80 2.68±0.06 2.66 – – – – – – Do-------------------------------------- 47 – – – – – – .0±0.1 . 0

Gabilan Range plutons ------------------ 74 2.67±0.04 2.66 – – – – – –

Montara Mountain pluton--------------- 48 2.68±0.06 2.69 – – – – – –

Do-------------------------------------- 25 – – – – – – .1±0.1 . 1

Gabbro of Logan-------------------------- 11 2.91±0.08 2.92 – – – – – –

Do-------------------------------------- 3 – – – – – – 1.4±0.4 1.3

Coast Range ophiolite

Mafic rocks -------------------------------- 8 2.68±0.10 2.85 .3±0.5 . 1

Serpentinite-------------------------------- 56 2.42±0.27 2.41 1.3±0.9 1.2

Cenozoic and Mesozoic sedimentary rocks

Sandstone (Miocene and Oligocene)-- 40 2.30±0.22 2.29 . 0 . 0

Sandstone (Eocene)----------------------- 61 2.49±0.12 2.53 . 0 . 0

Sandstone (Cretaceous)------------------ 14 2.59±0.07 2.59 – – – – – –

Do-------------------------------------- 4 – – – – – – . 0 . 0

Sandstone of the Purisima Formation- 15 – – – – – – .7±0.3* . 6

*Values of 0.4×10−3 to 0.6×10−3 cgs units were obtained from modeling the topography magnetically where sandstone of the Purisima Formation makes up the topography, 

and then comparing the results to the aeromagnetic map (pls. 3A, 3B). Similar values were obtained from more than 100 direct measurements on outcrops and in roadcuts.

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E56 EARTHQUAKE OCCURRENCE

than 6 km/s (Michael and Eberhart-Phillips, 1991; Foxall and 

others, 1993). Although the magnetic properties of the crust 

immediately beneath the gabbro of Logan serve to distinguish 

it from the gabbro, the seismic velocity and poorly defined 

density of the gabbro of Logan are compatible with those of 

several different rock types (Christensen, 1982).

Most of the basement of the Logan block is overlain by a 

thick Tertiary section of predominantly marine clastic rocks, 

ranging in age from Paleocene or Eocene to Pliocene, that, 

in turn, is covered in places by a thin mantle of Quaternary 

deposits (Clark and Rietman, 1973). The Tertiary section alone 

is more than 6 km thick and has been deformed into both broad 

and tight folds with axes subparallel to the trend of the San 

Andreas fault and the long axis of the Logan block. The Tertiary sedimentary rocks, which are significantly less dense than 

either granitic rocks of the Ben Lomond block or basement 

rocks of the Logan block (table 1), are the source of the linear 

gravity low over the Logan block that extends northwestward 

from the north edge of the Gabilan Range, a distance of more 

than 45 km. As discussed below, the absence of a pronounced 

gravity low over the north and, to some extent, south ends of 

the Logan block probably reflects a thickening of the slab of 

high-density gabbro basement rather than a thinning of the 

overlying sedimentary sequence. The sedimentary rocks are 

mostly nonmagnetic except for the Pliocene Purisima Formation, which, on the basis of field observations, contains high 

concentrations of detrital magnetite. Sedimentary rocks of the 

Purisima Formation are the source of both the linear magnetic 

anomaly that lies just northeast of and parallel to the Zayante 

fault (see pl. 3A), and the linear magnetic anomaly 8 km to the 

east, immediately southwest of the San Andreas fault, as well 

as of most of the low-amplitude, short-wavelength magnetic 

anomalies on the Ben Lomond block (see pl. 3A), because 

here Purisima rocks make up the topography and the magnetic 

anomalies reflect this topographic surface.

The northernmost block within the study area (see pl. 3A), 

southwest of the San Andreas fault, abuts the Logan block 

to the south and is bounded by the San Gregorio fault on the 

southwest and by the San Andreas and Pilarcitos faults on the 

northeast. The basement boundary between these two blocks 

is completely concealed beneath Tertiary sedimentary strata. 

The gravity and magnetic maps (pl. 3), however, indicate by 

their gradients that the basement rocks of the northernmost 

block are significantly less magnetic and probably less dense 

than those of the Logan block, properties similar to those of 

Late Cretaceous granitic rocks of the Montara pluton (see pl. 

3B), the nearest basement outcrop north of the Logan block. 

The contact, which trends northwest, is interpreted to be a 

fault because of its linearity and the greater age of the gabbro 

of the Logan block relative to other nearby plutonic rocks of 

the Salinia terrane (James, 1992). The associated magnetic 

anomaly indicates that this contact is subparallel to the Butano 

fault but offset from it by about 4 to 5 km to the north-northeast. Although the parallelism between the Butano fault trace 

and the buried north-northeast edge of gabbroic basement of 

the Logan block suggests that the two features are related, 

their coincidence at depth is difficult to support because the 

relative positions of the Butano fault and the buried basement 

contact would require a moderate northeastward dip, whereas 

geologic evidence indicates that the Butano fault dips steeply 

southwest (Cummings and others, 1962).

BLOCK BOUNDARIES SOUTHWEST OF

THE SAN ANDREAS FAULT

The faults or other structures that separate the crustal 

blocks southwest of the San Andreas fault typically are sites 

of abrupt changes in density and (or) magnetic susceptibility 

and so are generally characterized by gravity and (or) magnetic 

anomalies. Where such anomalies occur, the potential-field 

data provide the means for defining the subsurface position 

and configuration of the block boundaries through modeling. Although all such model interpretations are inherently 

ambiguous, within the study area (see pl. 3C) independent 

constraints on the modeling imposed by measured lithologic 

properties, geologic mapping, and other geophysical interpretations effectively reduce the ambiguity, such that meaningful 

geologic conclusions can be reached. The following geophysical discussion of the subsurface configuration of block 

boundaries progresses from southwest to northeast.

The San Gregorio fault at the mouth of Monterey Bay (see 

pl. 3A) is marked over a distance of more than 60 km by a 

magnetic gradient (high to the southwest), indicating that 

nonmagnetic basement rocks of the Ben Lomond block are in 

contact with a magnetic crustal block to the southwest (see pl. 

3B). The magnetic gradient extends northwestward and continues even to where the Logan block abuts the San Gregorio 

fault, probably because the top of magnetic crust southwest 

of the fault is much shallower than the magnetic basement of 

the Logan block. Models along the lines crossing this anomaly 

(for example, fig. 6) indicate that within the study area the 

magnetic boundary, presumably the San Gregorio fault, dips 

50°–70° NE. Seismic-reflection surveys of the offshore area 

immediately southwest of the San Gregorio fault reveal a 15-

km-wide zone of young folds and similar northeast-dipping 

thrust faults that approximately parallels the San Gregorio 

fault (Lewis, 1996).

Both gravity and magnetic modeling require that the Logan 

block extends southwestward beneath the Ben Lomond block, 

implying that the Zayante fault dips southwest. The inferred 

minimum southwest limit of magnetic basement rocks of 

the Logan block (see pl. 3B) in places is more than 10 km 

southwest of the trace of the Zayante fault. Because the top 

of crystalline basement of the Logan block lies at about 5- to 

7-km depth, on the basis of seismic tomography (Foxall and 

others, 1993; see Eberhart-Phillips and Michael, this chapter) 

and magnetic modeling (fig. 2), we conclude that the Zayante 

fault dips southwest along much of its length, in places at an 

average angle of less than 45°. This conclusion is supported 
E57

by the gravity profile along line R–R' (fig. 1), which shows 

that early Tertiary sedimentary rocks of the Logan block here 

continue laterally at least 6 km beneath a thin lip of granitic 

basement of the Ben Lomond block (see pl. 3C). Also, a threedimensional velocity model of this area (Foxall and others, 

1993; Eberhart-Phillips and Michael, this chapter) shows lowvelocity (<5 km/s) material continuing southwestward of the 

Zayante fault trace to 3- to 5-km depth, where drill-hole logs 

indicate crystalline basement to be generally less than 1 km 

deep (Clark and Rietman, 1973). The Zayante fault, however, 

lies near the southwest limit of the crustal volume that is well 

resolved in the velocity models.

Although both gravity and magnetic data indicate that the 

Logan block continues in the subsurface southwestward of 

the Zayante fault, the two data sets reflect different elements 

of this block. Specifically, the magnetic anomalies indicate 

the shape and position of the gabbroic basement, whereas the 

gravity anomalies primarily reflect the distribution of Tertiary 

sedimentary rocks.

Although the San Andreas fault does not produce conspicuous, continuous geophysical anomalies in the study area 

(see pl. 3), isolated gravity and magnetic anomalies provide 

important information on the subsurface configuration of the 

fault surface. The gravity profile along the line R–R' (fig. 1) 

-200

0 200

MAGNETIC ANOMALY,

IN NANOTESLAS

 0 5 10 15 20 25 30 35

10

5

0

DEPTH,

IN KILOMETERS

DISTANCE, IN KILOMETERS

gb 0.0011

gb

0.0008

0.0005

0.0007

Jou

(schematic only) San Andreas fault

1

� ��

�

�

Figure 4.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) along line S–S' (pl. 3) across the 

Logan block between the Zayante and San Andreas faults. Pluses, locations of magnetic observations. Numbers on units are magnetizations 

(in electromagnetic units per cubic centimeter). Most important features of magnetic model include calculated dip of the San Andreas fault 

and shapes of sliver of gabbro (gb) and wedge-shaped mass of gabbro beneath basin of Tertiary sedimentary rocks. Note magnetic boundary 

1, magnetic slab southwest (left) of it, and absence of magnetic slab extending to the San Andreas fault—that is, “hole” in slab near San 

Andreas fault. No vertical exaggeration. Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range ophiolite.

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E58 EARTHQUAKE OCCURRENCE

indicates that the San Andreas fault is vertical in the upper 3 

to 4 km of the crust and probably remains vertical to at least 

6- to 7-km depth. Starting about 4 km southeast of the point 

where line R–R' crosses the San Andreas fault, and extending 

southeastward along the fault a distance of about 10 km, a 

linear magnetic anomaly (see pl. 3A) indicates that magnetic 

sedimentary rocks of the Purisima Formation, which here are 

confined to the Logan and Ben Lomond blocks southwest of 

the fault, extend in the shallow subsurface as much as 1 km 

northeastward of the fault trace (see pl. 3A). The magnetic 

sedimentary rocks here are only 1 to 1.5 km thick, and so 

the deeper parts of the fault are undefined. Nevertheless, the 

fact that Purisima rocks are thrust beneath older sedimentary 

rocks northeast of the fault suggests that, in places, the nearsurface position of the San Andreas fault could be well to the 

northeast of its mapped position, although the distribution 

of aftershocks (Dietz and Ellsworth, 1990) suggests that, at 

depth, the fault dips steeply southwest. Farther southeast, the 

San Andreas fault forms the northeast face of the sliver of 

magnetic basement rocks of the Logan block exposed along 

the fault zone. The magnetic model along line S–S' (fig. 4B), 

which crosses the northwest end of this sliver, indicates that 

in the well-constrained upper 3 km of the model, the fault dips 

steeply southwest. Near the north end of the Gabilan Range, 

Figure 5.—Isostatic-residual-gravity profile (A) (circles, measured data sampled from gridded gravity map; curve, calculated values) and 

density model (B) along line S–S' (pl. 3) across the Logan block between the Zayante and San Andreas faults. Pluses, locations where gravity 

was calculated. Numbers on units are densities (in grams per cubic centimeter). Model is consistent with magnetic model (fig. 4B) and with 

sedimentary-rock densities modeled in figure 1. Most important features of density model include dip of the Zayante fault and configuration 

of sedimentary basin of the Logan block. Block of low-density material to right (northeast) of the San Andreas fault is schematic. See figure 

1 for explanation of high-density slab at base of model. No vertical exaggeration. gb, gabbro; gr, granitic and older metamorphic rocks; KJf, 

Cretaceous and Jurassic Franciscan complex; Qal, Quaternary alluvium; Ts, Tertiary sedimentary rocks. -40 -20 0 2

0

GRAVITY ANOMALY,

 IN MILLIGALS

 0 5 10 15 20 25 30 35

10

5

0

DEPTH,

IN KILOMETERS

gb

gb 2.91

Qal+Ts 2.17 2.87

2.91

KJf San Andreas fault

(schematic only)

DISTANCE, IN KILOMETERS

Qal+Ts 2.17

gr 2.67

Ts 2.37

Ts 2.48

Ts 2.55

Ts 2.61

Ts 2.64

Ts 2.67

2.87

gr? 2.67

Zayante fault

gb?

S S'

A

B
E59

about 15 to 20 km southeast of line S–S', the San Andreas fault 

separates Cenozoic sedimentary rocks on the northeast from 

plutonic rocks of the Ben Lomond block (see pl. 3C). The 

fault trace lies near the bottom of the gravity low caused by 

Cenozoic sedimentary rocks (see pl. 3C), indicating that these 

sedimentary rocks extend in the subsurface southeastward 

beneath the plutonic rocks. Modeling of this gravity anomaly 

(Pavoni, 1973) suggests that the fault dips about 70° SW. at 

least to the depth resolved by this model, about 6 km.

CRUST NORTHEAST OF

THE SAN ANDREAS FAULT

The part of the study area (see pl. 3B) northeast of the San 

Andreas fault is difficult to subdivide clearly into crustal 

blocks because it is structurally much more complex than that 

to the southwest. Using a combination of gravity, magnetic, 

and geologic data, however, we can subdivide the exposed 

crust northeast of the San Andreas fault and southwest of the 

Silver Creek and Calaveras faults into four major basement 

blocks bounded by faults. Although these basement blocks, 

to some extent, appear to be represented by tectonostratigraphic terranes as mapped by Blake and others (1984) and 

McLaughlin and others (1991), detailed geologic mapping 

of terranes in this area is by no means without uncertainties, 

and the final movements along some or, possibly, all of our 

block-bounding faults appear to postdate faulting that took 

place when the terranes were originally amalgamated.

We summarize the geology of this part of the study area (see 

pl. 3B) principally from the work of McLaughlin and others 

(1991) (see McLaughlin and Clark, this chapter). Adjacent to 

the San Andreas fault and bounded on the northwest side by 

segments of the Aldercroft, Sierra Azul, and Berrocal faults 

and the southern section of the Sargent fault (see pl. 3A) is 

an isolated layered sequence of Mesozoic and Cenozoic rocks 

(Sierra Azul block of McLaughlin and Clark, this chapter; 

see McLaughlin and others, 1991) consisting of (1) the basal 

Coast Range ophiolite (unit Jou), with its lower part of serpentinized ultramafic rocks and its upper mafic part overlain 

Figure 6.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) along line D–D'–D" (pl. 3) across 

the San Gregorio fault. Pluses, locations of magnetic observations. Numbers on units are magnetizations (in electromagnetic units per cubic 

centimeter). Magnetic-anomaly sources west (to left) of the San Andreas fault appear to have considerable depth extent. Most important 

feature is unexpectedly shallow northeastward dip of the San Gregorio fault. See figure 7 for details in northeast half of model. No vertical exaggeration. gb, gabbro; gr, granitic and older metamorphic rocks; Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range 

ophiolite; KJf, Cretaceous and Jurassic Franciscan complex; KJs, Cretaceous and Jurassic sedimentary rocks; Qal, Quaternary alluvium; 

Ts, Tertiary sedimentary rocks. -200 0 200 MAGNETIC ANOMALY, IN NANOTESLAS

-40 -30 -20 -10 0 10 20 30

DISTANCE, IN KILOMETERS

10 0

DEPTH,

IN KILOMETERS

gr

Ts+KJs

KJf

Jou

gb 0.002

0.0005

 0.0004 0.0006 0.00025

San Andreas fault

gr?

0.00022 San Gregorio fault

?

? Ts

?

D D' D''

A

B

Zayante fault

Tp

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E60 EARTHQUAKE OCCURRENCE

by (2) sequences of Jurassic and Cretaceous sedimentary 

rocks (unit KJs) that include as much as about 200 m of the 

lower part of the Great Valley sequence, overlain by as much 

as 1,600 m of Late Cretaceous sedimentary rocks, which, in 

turn, are overlain by (3) a sequence of Tertiary sedimentary 

rocks (unit Ts) of Eocene and Oligocene or early Miocene age. 

Northeast of this isolated block is a series of terranes that here 

collectively form the Franciscan Complex, a Late Jurassic to 

Cretaceous subduction complex consisting predominantly of 

metamorphosed turbiditic graywacke, siltstone, shale, and lesser 

amounts of chert and volcanic rocks. Additional rocks that characteristically occur in melange of the complex are serpentinite, 

limestone, and blocks of high-pressure-metamorphic rocks. All, 

or nearly all, of the serpentinite is considered to be derived from 

the Coast Range ophiolite (R.G. Coleman, written commun., 

1991; McLaughlin and others, 1991). For structural reasons, 

following McLaughlin and others (1991), we have subdivided 

the Franciscan Complex on our geologic maps (pls. 3A, 3B) 

into three northwest-trending belts of rocks, described as follows, from southwest to northeast: (1) the Bald Mountain-El 

Sombroso terrane (unit KJf1), predominantly metasandstone 

and argillite, with associated metabasalt flows and radiolarian 

chert, especially along the eastern part of the belt; (2) the Permanente terrane (unit KJf2), predominantly metabasalt flows, 

with lesser amounts of limestone, chert, argillite, and graywacke 

(Sliter and others, 1991); and (3) unspecified terranes (unit 

KJf3), predominantly argillite and graywacke, metabasalt, and 

serpentinite.

In the northeast corner of this part of the study area are two 

smaller patches of Mesozoic sedimentary rocks (unit KJs, pl. 

3A) correlative with the Great Valley sequence and in fault 

contact with the Franciscan Complex. Overlying the larger 

of these two patches (in the Santa Teresa Hills) are Tertiary 

(Eocene) sedimentary rocks, similar to those in the isolated 

block (previously described) adjacent to the San Andreas fault. 

In this northeast corner near the Silver Creek fault are also several small patches of mafic rocks associated with serpentinite 

(unit Jou); collectively, these rocks are considered to be Coast 

Range ophiolite (R.G. Coleman, oral commun., 1991) Small 

isolated patches of younger Tertiary (early to middle Miocene) 

sedimentary rocks (unit Ts) that are distributed irregularly 

throughout this part of the study area lie unconformably and 

(or) are faulted over Franciscan rocks. A series of Pliocene mafic 

volcanic rocks (unit Tv) are locally associated with the Silver 

Creek and Calaveras faults. Pliocene and Quaternary alluvial 

deposits (unit Qal) cover some older rocks.

Magnetic properties of the lithologic units in this part of the 

study area (see pls. 3A, 3B) are reflected both in magnetic-susceptibility measurements on rock samples and in the distribution 

of magnetic anomalies relative to the underlying rocks. Magnetic-susceptibility measurements on rock samples (table 1) 

indicate that the serpentinite and serpentinized ultramafic rocks 

are highly magnetic, whereas the mafic rocks of the Coast Range 

ophiolite are only weakly magnetic. The metavolcanic rocks 

(greenstone) of the Franciscan Complex mostly range from 

nonmagnetic to weakly magnetic because the low-grade metamorphism has destroyed the bulk of their original magnetite. 

The sedimentary rocks and deposits of all ages are essentially 

nonmagnetic, with the local exception of the Pliocene Purisima 

Formation, as described above in the subsection entitled “Crust 

Southwest of the San Andreas Fault.” These results from rock 

samples are fully confirmed by examining the correlation of 

the magnetic map with the geologic map (pl. 3A). Nearly all 

outcrops of serpentinized ultramafic rocks produce magnetic 

anomalies, whereas the large areas of Franciscan metavolcanic 

rocks (not shown on pl. 3) within the Permanente terrane do not 

produce significant magnetic anomalies. The mafic components 

of the Coast Range ophiolite near the San Andreas fault do not 

display conspicuous magnetic anomalies, even where the rocks 

form topographic highs and so were close to the magnetometer. 

The much younger volcanic rocks in restricted amounts near 

the Silver Creek and Calaveras faults produce high-amplitude, 

narrow magnetic highs and lows, attesting to their intense 

remanent magnetizations.

The densities of the rock types that occur northeast of the 

San Andreas fault are listed in table 1. Metavolcanic rocks 

(greenstone) of the Franciscan Complex and mafic rocks of 

the Coast Range ophiolite are significantly denser than most 

other rocks in the study area (see pl. 3). The densities of serpentinite samples (serpentinized ultramafic rocks) vary widely, 

as indicated by the large standard deviation associated with 

the measurements on 56 samples, but on average are lower 

than those of most pre-Tertiary rocks. The sedimentary rocks 

generally increase in density with increasing age; the average 

densities of Pliocene and Quaternary deposits are lower than 

those shown for Oligocene and Miocene sandstones.

BLOCK BOUNDARIES NORTHEAST OF

THE SAN ANDREAS FAULT

Certain major curvilinear magnetic anomalies (see pls. 3A, 

3B) extend across the part of the study area northeast of the 

San Andreas fault. To locate the boundaries of the anomaly 

sources, we used an automatic technique (Cordell and Grauch, 

1985; Blakely and Simpson, 1986). After integrating the magnetic map to produce the pseudogravity map, this technique 

calculates the locations of the maximum horizontal gradients 

on the anomaly margins and at these locations plots a sinuous 

series of dots that represents the pseudodensity boundaries—in 

other words, the boundaries of the magnetic rocks. This method 

depends on the theoretical observation that gravity-anomaly 

(or pseudogravity) gradients tend to be steepest over the edges 

of gravity sources. These boundaries were then extended and 

connected visually on an overlay of the magnetic map (pl. 3A) 

to form the magnetic boundaries that subdivide the northeastern 

part of the study area into four major blocks; these magnetic 

boundaries are numbered from 1 to 3 and display two lesser 

branches numbered 1A and 2A. We consider these three magnetic boundaries to represent major faults, and the subsequent 

calculated geophysical models display our interpretation of 

their geologic significance.
E61

Comparisons of these magnetic boundaries and blocks with 

the geologic map (pl. 3A) yield the following associations 

and relations.

1. The block between the San Andreas fault and magnetic 

boundary 1 is characterized by a major magnetic high and 

is composed of two different structural elements: (1) the 

northwestern part of the block, which consists of an isolated 

geologic terrane of Mesozoic ophiolite and sedimentary 

rocks overlain by Tertiary sedimentary rocks collectively 

referred to as the Sierra Azul structural block by McLaughlin and others (this chapter); and (2) the southeastern part 

of the geophysically defined block, which is composed of 

the Franciscan Complex. Magnetic boundary 1 is defined 

as the northeast contact of a large mass of magnetic rocks 

and closely follows parts of the Aldercroft, Sierra Azul, and 

Berrocal faults where they bound the Sierra Azul structural 

block. Moderately continuous masses of serpentinite occur 

along these parts of magnetic boundary 1, apparently in 

association with the lower part of the Coast Range ophiolite. 

Nevertheless, along magnetic boundary 1 to the southeast, 

its southeast end diverges eastward from the Berrocal fault 

and extends across the Franciscan Complex, approximately 

(but not exactly) following the southeast side of the Permanente terrane. The magnetic high associated with this block 

(see pl. 3B) is terminated in the southeast by the Calaveras 

fault, which also terminates magnetic boundary 1. Magnetic 

boundary 1 is probably the strongest and most consistent 

magnetic feature within the study area northeast of the San 

Andreas fault, and we regard its substantial southeastward 

divergence from the Berrocal fault as an important factor in 

our structural interpretations of this part of the study area. 

This divergence indicates that it is incorrect to interpret the 

source of the magnetic high simply to be serpentinite in 

the layer of Coast Range ophiolite at the base of the lower 

part of the Great Valley sequence (combined in unit KJs, 

pl. 3A) because to the southeast the magnetic rocks mostly 

underlie the Franciscan Complex. A separate northwesterly 

branch off magnetic boundary 1 (magnetic boundary 1A, 

pl. 3A) extends into the next block to the northeast but, as 

for magnetic boundary 1, forms the northeast boundary of 

a mass of magnetic rocks. This magnetic boundary follows 

the Berrocal fault and then swings westward along the Lime 

Kiln fault in a bend similar to that in magnetic boundary 1 

at the Aldercroft fault; the similarity in shape thus suggests 

a relation between these two boundaries. Farther northwest 

(see pl. 3B), magnetic boundary 1A converges with and is 

ultimately truncated by the San Andreas fault. Along much 

of its northern part, magnetic boundary 1A appears to follow 

the southwest side of the Permanente terrane and is intermittently associated with small serpentinite masses.

2. The block between magnetic boundaries 1 and 2 is composed 

of the Franciscan Complex, specifically the Permanente 

terrane and the Bald Mountain-El Sombroso terrane. 

Magnetic boundary 2 approximately follows the northeast 

geologic contact of the Permanente terrane and correlates 

with the northeast side of a zone of sheared melange originally mapped by Bailey and Everhard (1964) that contains 

numerous blocks of blueschist. Magnetic boundary 2 is 

defined geophysically as the southwest border of a series 

of nearly continuous magnetic anomalies that are caused by 

nearly continuous curvilinear masses of serpentinite associated with the third block and lying northeast of magnetic 

boundary 2. Northwest of line D–D' (pl. 3A), the magnetic 

boundary is located within the Permanente terrane because 

here the magnetic serpentinite masses lie within the terrane. 

The block bounded by magnetic boundaries 1 and 2 contains 

few magnetic anomalies other than those along the southwest side of magnetic boundary 1A, which are caused by 

a belt of serpentinite.

3. The block between magnetic boundaries 2 and 3 consists of an assemblage of Franciscan metavolcanic and 

metasedimentary rocks and abundant small masses of 

serpentinite, which cause various small, irregular magnetic 

highs. Magnetic boundary 2A, which is located within this 

block, is similar to magnetic boundary 2 in that both are 

located as the south or southwest boundaries of magnetic 

serpentinite belts. The north boundary of the block (magnetic boundary 3) is a continuous linear magnetic boundary 

on the south side of a coherent magnetic feature. Where it 

coincides with the south margin of exposed serpentinite, 

this magnetic boundary may partly correspond to the Shannon fault.

4. The block between magnetic boundary 3 and the Silver 

Creek fault is composed predominantly of magnetic serpentinite and mafic components of the Coast Range ophiolite, 

together with metasedimentary rocks of the Franciscan 

Complex and, importantly, some sedimentary rocks of 

the Great Valley sequence (lower part of unit KJs, pl. 3A) 

that are not unexpected, considering the large amount of 

basal Coast Range ophiolite exposed in this block. The 

metasedimentary rocks were mapped as part of the Central 

belt of the Franciscan Complex by Blake and others (1984), 

which they defined as the melange matrix enclosing other 

Franciscan terranes.

In addition to the curvilinear magnetic features and the obvious correlations of magnetic highs with serpentinite, a smaller 

but significant number of other subcircular and irregular magnetic anomalies are associated with the localities of Franciscan 

rocks northeast of the San Andreas fault. We suggest that most 

of these other anomalies are caused by isolated masses of serpentinite that have not yet been mapped or are buried but that, 

on the basis of magnetic-gradient studies, are present at depths 

of less than a few hundred meters below the surface.

THREE-DIMENSIONAL STRUCTURE

NORTHEAST OF THE SAN ANDREAS FAULT

In the preceding section, we described the magnetic map 

(pls. 3A, 3B) and its relations to geology mainly in terms of 

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E62 EARTHQUAKE OCCURRENCE

plan view. Here, we show two-dimensional geophysical models 

of the data in order to describe the geologic and geophysical 

structure of the study area in three dimensions. We display 

about half of the models that we have actually calculated, using 

those that best illustrate the subsurface structure.

The magnetic anomaly bounded to the northeast by magnetic 

boundary 1 is modeled from the magnetic profiles along lines 

D–D' (fig. 7A) and E–E' (fig. 8A). The magnetic models show 

a magnetic slab, 2 to 3 km thick, that dips 45°–55° SW. until 

cut off by the San Andreas fault. The slab is too magnetic to 

contain significant amounts of weakly magnetic mafic material 

from the Coast Range ophiolite, and so is presumed to be composed substantially of serpentinite or partially serpentinized 

ultramafic rocks. This slab cannot contain significant amounts 

of unserpentinized ultramafic rocks, or it would produce a 

correlative gravity high. Examination of the gravity map (pl. 

3C) shows that at the northwest end of magnetic boundary 1 

(along the Aldercroft fault), the gravity high in this area (12- 

and 14-mGal contours) extends northward of both magnetic 

boundaries 1 and 1A, indicating that the source of the gravity 

high is not the magnetic slab but must, in fact, lie below both 

it and the dipping magnetic slab forming magnetic boundary 

1A (fig. 7). The southeastward extension of magnetic boundary 1 into the Franciscan Complex shows that here the dipping 

magnetic slab is both overlain and underlain by Franciscan 

rocks—that is, it is embedded within the Franciscan Complex. Additionally, in those areas where the slab underlies the 

Great Valley sequence (figs. 7, 8), this Cretaceous sequence 

and ophiolite are stratigraphically so thin (McLaughlin and 

others, 1991) that they presumably rest in fault contact on a 

wedge-shaped mass of the Franciscan Complex that, in turn, 

overlies the magnetic slab. We similarly interpret the magnetic 

slab southwest of magnetic boundary 1A (fig. 7), and magnetic 

models indicate that this second slab is thinner (1 km thick) 

and probably branches off from the main slab of magnetic 

boundary 1. In summary, the magnetic anomalies bounded on 

the northeast by magnetic boundaries 1 and 1A are interpreted 

to be caused by southwest-dipping serpentinite slabs that are 

generally embedded within the Franciscan Complex, but 

locally the northeast tip of the major slab contacts (or coincides 

with) mafic and ultramafic components of the Coast Range 

ophiolite where they are exposed near the southwest side of 

magnetic boundary 1.

In the southeastern part of the magnetic slab of magnetic 

boundary 1, adjacent to the San Andreas fault, is a broad 

magnetic low (−100-nT contour, pl. 3B), about 20 km long 

and 10 km wide. The magnetic model across this low (fig. 

4B) indicates that the magnetic slab is either missing here or 

of negligible thickness. The northeast boundary of this hole 

in the slab is shown on plate 3, and its southwest boundary is 

the San Andreas fault (fig. 9).

The magnetic profile along line M–M' (fig. 9) demonstrates 

some complexities in the geophysical expression of the rocks 

lying southeast of magnetic boundary 1. Magnetic boundary 1, 

which is determined from magnetic data, is located about 1 km 

northeast of the calculated (from gravity data) northeast contact 

of the concealed, but near-surface, Coast Range ophiolite. This 

discrepancy is the result of a local steep circular magnetic feature, caused by a small, near-circular mass of magnetic rocks, 

probably serpentinite, located at approximately km 19 on the 

profile (see pl. 3A; fig. 9). Another complicating factor, lying 

directly above the ophiolite slab, is a concealed mass of highdensity (2.87 g/cm3

) rocks (probably basaltic volcanic rocks) 

that also correlate with a moderate magnetic anomaly caused 

by a near-surface source and located between approximately 

km 15 and 17 on the profile (see pl. 3A; fig. 9). Although this 

presumed magnetic basalt is shown to be very near the surface 

by both gravity and magnetic data, the geologic maps (pls. 

3A, 3B) show only a broad area of Franciscan sedimentary 

rocks (Wentworth and Blake, 1991) in this area. This unusual 

rock type (high density, moderately magnetic) may be a mafic 

component of the Coast Range ophiolite or, less likely, a mafic 

volcanic rock type within the Franciscan Complex.

Magnetic boundary 2 is located near the northeastern margin 

of the Permanente terrane. This geologic contact corresponds 

well to the lower part of a locally steep gravity gradient (see 

pl. 3C) near the 2- to 4-mGal contours to the northwest and the 

6-mGal contour to the southeast. Because of this correspondence between the geologic contact and the gravity gradient, 

and because rocks of the Permanente terrane have a relatively 

high average density (estimated at 2.83 g/cm3), we consider 

that both the gravity high southwest of magnetic boundary 2 

and the gravity gradient are caused primarily by exposed rocks 

of the Permanente terrane and its inferred buried southwestward extension. This gravity high is a conspicuous feature 

that extends along the northeast side of the San Andreas fault 

(see pl. 3C) for about 100 km. In the density model along line 

M–M' (fig. 9B), we show that the gravity anomaly can be modeled by a slab of the Permanente terrane, about 2 km thick and 

dipping about 30° SW., approximately parallel to the magnetic 

slab (magnetic boundary 1) that on this cross section overlies 

the Permanente terrane.

The magnetic models crossing magnetic boundary 2, along 

line E–E' (fig. 8B), indicate that the magnetic rocks forming 

a nearly continuous belt on the northeast side have a moderate depth extent relative to their width. Similarly, modeling of 

magnetic anomalies in the area between magnetic boundaries 

2 and 3 indicates that they have shallow sources with a vertical 

extent of generally less than 1 km. Examination of the geologic 

map (pl. 3B; Wentworth and Blake, 1991) between magnetic 

boundaries 2 and 3 shows large, relatively equidimensional 

areas of metavolcanic rocks that are 1.5 by 2.0 km in plan 

view or larger; however, the gravity map (pl. 3C) displays no 

significant gravity highs here, even though a thickness of only 

1 km of basalt (density, 2.88 g/cm3) should produce an anomaly 

of nearly 10 mGal. Indeed, the gravity field here is relatively 

featureless and has a background level of about 0 mGal, not 

much different from levels over metasedimentary rocks of 

the Franciscan Complex in the Diablo Range 25 km farther 

east. We conclude that the exposed units of the block between 
E63

magnetic boundaries 2 and 3 must generally be relatively flat 

lying, that the serpentinite and metavolcanic rocks are less than 

1 km thick, and that the underlying material is nonmagnetic and 

composed of moderate-density, presumably metasedimentary 

rocks. Because gravity levels are slightly higher (6 mGal) in 

the eastern part of this block between magnetic boundaries 2 

and 3, we conclude that here the metavolcanic rocks may be 

as much as 1 to 2 km thick.

The major change in structural style across magnetic boundary 2, southwest-dipping units to the southwest and different, 

relatively flat lying rocks to the northeast, indicates a structural boundary as well as a fold in this area and confirms the 

importance of the shear zone of melange on the southwest 

side of boundary 2.

Magnetic models across magnetic boundary 3 and the 

block to the north are shown in figures 7B and 8B. Again, the 

structures are flat lying, except close to the Silver Creek fault. 

As shown in figure 8B, the serpentinite sheet is essentially 

continuous across the Santa Clara Valley beneath the alluvium, 

as previously inferred from geologic studies (R.G. Coleman, 

written communs., 1990, 1991). Magnetic boundary 3 is thus 

the south edge of a nearly continuous subhorizontal sheet of 

Figure 7.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) along line D'–D'' (pl. 3) across the 

Logan block and area northeast of the San Andreas fault. Pluses, locations of magnetic observations. Numbers on units are magnetizations 

(in electromagnetic units per cubic centimeter). Gabbroic body southwest of the San Andreas fault is consistent with other models (figs. 

1, 4–6) across the Logan block. Three areas of the Franciscan Complex surrounding thick, dipping slab (km 15–19) and thin, dipping slab 

(km 18–23) are considered to be wedges. Magnetic boundary 3 corresponds here to the Shannon fault; to the northeast, a horizontal slab of 

Coast Range ophiolite overlies a horizontal wedge contact with Franciscan rocks. No vertical exaggeration. gb, gabbro; gr, granitic and older 

metamorphic rocks; Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range ophiolite; KJf, Cretaceous and Jurassic Franciscan 

complex; KJs, Cretaceous and Jurassic sedimentary rocks; Tp, Pliocene Purisima Formation; Ts, Tertiary sedimentary rocks. -200

0 200

MAGNETIC ANOMALY,

IN NANOTESLAS

0 5 10 15 20 2 30 35

10

5

0

DEPTH,

IN KILOMETERS

DISTANCE, IN KILOMETERS

Jou

0.0004

0.0006 0.0005

KJf

KJf

KJf

gr Ts

gr?

gb 0.002

0.00025

Shannon fault

Lime Kiln fault

Berrocal fault

Aldercroft fault

San Andreas fault

Zayante fault

1 1A 2 3

Jou KJf

Ts+KJs

D' D''

A

B

Tp

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E64 EARTHQUAKE OCCURRENCE

serpentinite, the lower part of a mass of Coast Range ophiolite. 

This sheet may originally have extended southwestward to 

magnetic boundary 2 but has since been mostly eroded away 

to leave the thin small serpentinite masses presently observed 

here. This interpretation implies that magnetic boundary 3 is a 

fault with a throw of about 1 km down to the north to account 

for the missing 1 km of serpentinite on the south side. This 

interpreted fault at least partly correlates with the Shannon 

fault. The gravity field in the area northeast of magnetic 

boundary 3 is a gradient that is caused by a large mass of 

low-density rocks on the northeast side of the Silver Creek 

fault. These rocks, possibly Tertiary sedimentary rocks, extend 

in the subsurface as far as 6 km southwestward horizontally 

beneath the Franciscan terranes along the southwest-dipping 

Silver Creek fault (Jachens and others, 1990), but the details 

are beyond the scope of this paper.

Figure 8.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) along line E–E' (pl. 3) across the 

Logan block and area northeast of the San Andreas fault. Pluses, locations of magnetic observations. Numbers on units are magnetizations 

(in electromagnetic units per cubic centimeter). Gabbroic body southwest of the San Andreas fault is consistent with other models (figs. 1, 

4–7) across the Logan block. Thick, dipping slab bounded by magnetic boundary 1 between km 9 and 15 is considered to have wedge contacts 

on each side. Magnetic boundary 3 at km 23 may correspond to the Shannon fault. Horizontal slab of Coast Range ophiolite overlying a 

horizontal wedge contact is between magnetic boundary 3 and the Silver Creek fault. Note wedge of magnetic Purisima Formation (sandstone) 

projecting in subsurface northeastward beneath trace of the San Andreas fault. No vertical exaggeration. gb, gabbro; gr, granitic and older 

metamorphic rocks; Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range ophiolite; KJf, Cretaceous and Jurassic Franciscan 

complex; KJs, Cretaceous and Jurassic sedimentary rocks; Tp, Pliocene Purisima Formation; Ts, Tertiary sedimentary rocks. -300 -100 100 300 MAGNETIC ANOMALY, IN NANOTESLAS

0 5 10 15 20 25 30 35

10

5

0

DEPTH,

IN KILOMETERS

DISTANCE, IN KILOMETERS

Zayante fault

San Andreas fault

Shannon fault

Silver Creek fault

? ?

gr

Ts

Ts+KJs

KJf

KJf

Jou

0.0002 0.0003

0.0002

Tp

gb 0.002

0.0004

0.0004

0.0017

0.0015

Jou

gr?

KJf

E E'

A

B

?

1 2 3
E65

IMPLICATIONS FOR

ORIGIN AND TECTONIC EVOLUTION

CRUST SOUTHWEST OF

THE SAN ANDREAS FAULT

Ross (1970) pointed out the unusual petrology of the gabbroic rocks exposed at the southeast end of the Logan block 

in comparison with granitoids of the rest of the Salinia terrane 

and with gabbros within the Franciscan Complex. He also 

drew attention to the striking similarities between the gabbro 

of Logan and gabbroic rocks exposed on the opposite side 

of the San Andreas fault at Gold Hill and near Eagle Rest 

Peak, about 150 and 300 km, respectively, southeast along 

the fault. On the basis of these observations, he proposed that 

Figure 9.—Isostatic-residual-gravity profile (A) (circles, measured data sampled from gridded gravity map; curve, calculated 

values) and density model (B) along line M–M' (pl. 3) across the Logan block and area northeast of the San Andreas fault. 

Pluses, locations where gravity was calculated. Numbers on units are densities (in grams per cubic centimeter). Most 

important features are two southwest-dipping bodies with densities of 2.83 and 2.87 g/cm3. 2.83-g/cm3 body, which is 

the Permanente terrane of the Franciscan Complex, is here shown to be main source of gravity high (pl. 3) extending 100 

km along northeast side of the San Andreas fault. 2.87-g/cm3, moderately magnetic body may be Permanente terrane or 

a mafic part of the Coast Range ophiolite. Intervening material is magnetic slab associated with magnetic boundary 1 at 

km 20.5 and terminated by hole in slab at km 14.5. Circled numbers, magnetic boundaries. Bodies southwest of the San 

Andreas fault (sedimentary rocks on gabbro) are somewhat schematic because the Zayente fault is concealed here. No 

vertical exaggeration. gb, gabbro; gr, granitic and older metamorphic rocks; Jou, Jurassic ultramafic rocks and serpentinite 

of the Coast Range ophiolite; KJf, Cretaceous and Jurassic Franciscan complex; KJs, Cretaceous and Jurassic sedimentary 

rocks; Tp, Pliocene Purisima Formation; Qal, Quaternary alluvium; Ts, Tertiary sedimentary rocks. -20

0 20

GRAVITY ANOMALY,

IN MILLIGALS

0 5 10 15 20 25 30

10

5

0

DEPTH,

KILOMETERS

DISTANCE, IN KILOMETERS

Qal 2.17 Qal 2.17

Ts 2.48

Ts 2.52 

Ts

Ts 2.59 

Ts

gb

gr 2.67

gr? 2.67

2.91

 2.67

 2.56 

2.87

Ts 2.48KJs+Ts 2.52

KJf 2.87

KJf 2.83

Jou 2.67

KJf

2.67

Jou 2.67

Permanente terran

KJf 2.67

e

gb?

San Andreas fault

Zayante fault

M M'

A

B

1 2

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E66 EARTHQUAKE OCCURRENCE

the gabbroic bodies had once been part of a continuous belt 

which was dismembered by movement along the San Andreas 

fault system. The fact that the gabbros of Logan and Gold 

Hill are thin slivers within the fault zone led him to mention 

an alternative interpretation, namely, that the entire Salinia 

terrane is underlain by similar gabbro and that these two 

outcrops represent thin slivers that leaked up along the fault 

zone. Identical U-Pb ages of 161 Ma on zircon from samples 

of the gabbros of Logan and Eagle Rest Peak (James and 

others, 1986) are older than those of Salinian granitoids and 

support the correlation proposed by Ross (1970). Stable- and 

radiogenic-isotopic and trace-element data from samples of 

the gabbros of Logan, Gold Hill, and Eagle Rest Peak (Johnson 

and O’Neil, 1988) also support this correlation, suggesting 

that these rocks represent oceanic crust subjected in the past 

to seawater hydrothermal alteration. Recent isotopic studies 

(James, 1992) show that the timing of plutonism and metamorphism, initial Pb- and Sr-isotopic ratios, and the style of 

zircon discordance of basement rocks of the Ben Lomond 

block are nearly identical to those of basement rocks in the 

southern Sierra Nevada, supporting the idea that these rocks 

also were once closer together.

The geophysical data and interpretations presented here provide further support for the idea that the gabbros of Logan and 

Eagle Rest Peak were once joined. First, the magnetic anomaly 

over Eagle Rest Peak (Youngs and Mattison, 1989) is nearly 

identical in shape and amplitude to the elliptical magnetic 

anomaly over the northern part of the Logan block (see pl. 

3B), especially when the depth of burial of the Logan source 

(relative to the exposed Eagle Rest Peak source) is taken into 

account. Second, gravity values over Eagle Rest Peak and the 

northern part of the Logan block are nearly identical, again 

when the effects of depth of burial are taken into account. In 

the vicinity of Eagle Rest Peak, the isostatic-residual-gravity 

values over granitic rocks of the southern Sierra Nevada average about +20 mGal but increase to more than +30 mGal near 

the outcrop of the gabbroic body (Roberts and others, 1990). 

Corresponding granitic rocks in the Ben Lomond pluton 

(+20 mGal) and the inferred gabbroic body (+30 mGal) show 

gravity values nearly identical to those in the vicinity of Eagle 

Rest Peak when the gravity effects of the overlying thickness 

of Cenozoic deposits are removed (fig. 3). Finally, though not 

directly supporting the proposed former connection between 

the gabbros of Logan and Eagle Rest Peak, the magnetic data 

do show that the thin sliver of the exposed gabbro of Logan is 

actually part of a basement block, 90 km long and as much as 

20 km wide, which is separated from the Ben Lomond block 

to the south by the Zayante fault. These magnetic data indicate 

that the gabbro of Logan is restricted to a single structural 

block within the Salinia terrane. The absence of deep-source 

magnetic anomalies associated with granitoids of the Salinia 

terrane argues against an interpretation that magnetic gabbro 

everywhere underlies the Salinia terrane.

An additional implication of the probable pre-San Andreas 

tie between the gabbros of Logan and Eagle Rest Peak stems 

from the interpretation, based on magnetic anomalies, that 

the basement of the gabbro of Logan abuts the San Andreas 

fault for a horizontal distance of about 65 km (see pl. 3C), 

whereas the gabbro of Eagle Rest Peak appears to abut the 

San Andreas fault for a maximum horizontal distance of only 

35 km (Youngs and Mattison, 1989). Immediately northwest 

of the gabbro of Eagle Rest Peak, an approximately 30-kmwide, northwest-trending magnetic anomaly terminates at the 

San Andreas fault (Youngs, 1989). This anomaly is the south 

end of the Great Valley magnetic anomaly (Griscom, 1966), 

a huge feature that extends northwestward along the entire 

Great Valley for a distance of 700 km and whose source is 

completely concealed beneath sedimentary fill. The combined 

San Andreas fault-bounded length of the gabbro of Eagle Rest 

Peak and the adjacent source of the Great Valley magnetic 

anomaly is about 60 to 70 km, similar to the San Andreas 

fault-bounded length of the gabbro of Logan. Therefore, we 

speculate that the gabbro of Eagle Rest Peak is correlated with 

the source of the Great Valley magnetic anomaly and that the 

gabbroic basement of the Logan block is also correlated with 

this source. The general form of the mass causing the Great 

Valley magnetic anomaly is that of a thick, west-dipping slab; 

interpreted to represent ophiolite obducted from the west onto 

the North American Continent during Jurassic time (Griscom, 

1982; Griscom and Jachens, 1990).

Geophysical modeling of the Zayante Fault (figs. 1, 4, 5) 

indicates substantial uplift and horizontal shortening across this 

fault. Crystalline basement is uplifted at least 4 to 6 km probably along most of the length of the fault, and the southward 

continuation of Tertiary sedimentary rocks beneath the granitic 

basement (fig. 1) suggests horizontal shortening of at least 6 

km in this area. Gravity and magnetic models along line S–S'

(figs. 4B, 5B) also show sedimentary rocks extending southward 

beneath granitic basement near the south end of the Logan 

block, but in this area the overlap is not well constrained by the 

gravity data. Magnetic basement of the Logan block extends in 

the subsurface more than 10 km southward of the surface trace 

of the Zayante fault (see pl. 3B), suggesting horizontal shortening of at least this amount. As discussed below, however, some 

of this horizontal shortening may reflect pre-Tertiary movement 

on a predecessor to the Zayante fault and thus predate thrusting 

of the sedimentary rocks beneath the block to the south. The 

timing of offset on the Zayante fault, as summarized by Clark 

and Rietman (1973), indicates that dip-slip movement affecting 

the Tertiary sedimentary section began in early Oligocene time 

and had mostly ended by early Miocene time, although some 

movement continued at least into the Pliocene. Furthermore, 

correlation of clasts in sedimentary rocks north of the Zayante 

fault with source rocks in the northern Gabilan Range suggests no major strike-slip movement along this fault since late 

Oligocene time (Clark and Rietman, 1973).

An unusual feature of the Zayante fault revealed by the 

geophysical data is that although this fault bounds the Logan 

block and its highly magnetic gabbroic basement for a distance of more than 90 km, no magnetic gabbro occurs in the 
E67

upthrown block—that is, no appropriate magnetic anomalies 

are present even in covered areas. This feature indicates that 

if the Zayante fault originated in Tertiary time after deposition of the Eocene Butano Sandstone, then either it broke 

outboard of all gabbro or subsequent movement, or erosion 

has effectively eliminated all gabbro locally from the Ben 

Lomond block, or both. Alternatively, Tertiary movement on 

the Zayante fault merely reflects reactivation of an older fault 

that, by Eocene time, already had juxtaposed the markedly 

different basement rocks of the Logan and Ben Lomond 

blocks. We suggest that the absence of magnetic gabbro in 

the Ben Lomond block argues against an Eocene or later 

origin for the Zayante fault.

That the Zayante fault may have fortuitously cut the basement outboard of all gabbro seems unlikely in light of the 

fact that the mapped or inferred trace of the fault is quite 

close to (never more than 10 km from and typically much 

closer to) the southwest edge of gabbroic basement along 

the entire 90 km reach of the fault. Eocene or later uplift 

and erosion sufficient to strip off all the gabbro also seems 

unlikely because, at least along the north third of the fault, 

Eocene sedimentary rocks of the Butano Sandstone are preserved on the upthrown block (Clark and Rietman, 1973). 

Furthermore, erosion of gabbroic basement that in places 

would have been at least 2 km thick should have supplied 

mafic sedimentary debris into the basin to the north, but the 

Tertiary sedimentary rocks do not contain such debris.

An explanation of the absence of gabbro south of the Zayante fault by strike-slip movement along this fault is possible 

but would constrain the allowable timing and magnitude 

of such movement. The proposed movement must have 

occurred earlier than the late Oligocene clast-source-rock 

crossfault tie identified by Clark and Rietman (1973), and 

probably earlier than the early Oligocene initiation of uplift 

south of the fault. Right-lateral offset of at least 100 km on 

the Zayante fault would have carried all gabbro (including 

counterparts to the gabbro of Eagle Rest Peak) westward of 

the current trace of the San Gregorio fault, and subsequent 

movement on the San Gregorio and San Andreas faults would 

have swept the outboard block northward to the vicinity of 

Point Reyes (Griscom and Jachens, 1989), where detailed 

aeromagnetic coverage west of the San Andreas fault is 

limited to a strip about 20 km wide paralleling the fault 

(Chase and Youngs, 1979). In the entire strip extending from 

about 30 km south to about 70 km north of Point Reyes, only 

one small magnetic anomaly is present (on the Point Reyes 

Peninsula), associated with a pendant of magnetic metamorphic rocks within the Cretaceous pluton, not with gabbro. 

Although the absence of gabbroic bodies in this strip does 

not rule out right-lateral strike-slip offset as an explanation 

for the absence of gabbro southwest of the Zayante fault, it 

does suggest that for this explanation to be viable, the offset 

would probably have to greatly exceed 100 km. Similarly, 

the absence of magnetic gabbro or any large unexplained 

magnetic anomalies in the southern Sierra Nevada batholith 

south of the gabbro of Eagle Rest Peak and its associated 

magnetic anomaly (Youngs, 1989; Youngs and Mattison, 

1989) indicates that hundreds of kilometers of left-lateral 

offset would be required to explain the absence of gabbro in 

the Ben Lomond block. Therefore, lacking any positive evidence for gabbro in places compatible with Eocene or later 

offset across the Zayante fault, we favor the interpretation 

that a proto-Zayante fault existed which was subsequently 

reactivated during Tertiary time.

The existence of a proto-Zayante fault is uncertain because, 

in the Santa Cruz Mountains, it would have been severely 

modified by later movement and its counterpart in the southern Sierra Nevada may be concealed. The contact between 

the gabbro of Eagle Rest Peak and Sierran granitoids to the 

south occupies a structural position comparable to that of the 

Zayante fault, is buried beneath Tertiary sedimentary rocks, 

and dips south, according to the aeromagnetic data (Youngs 

and Mattison, 1989). This contact could be a reverse or thrust 

fault that juxtaposed deep-seated batholithic rocks against 

the gabbro by large-scale vertical movement, possibly along 

an earlier suture between the gabbro and Sierran country 

rocks. According to McLaughlin and Clark (this chapter), 

evidence from fission-track data suggests that by about 70 

Ma, Ben Lomond granitic rocks were within 3 to 5 km of 

the surface, even though geobarometry of intrusive rocks 

and metamorphic roof pendants suggests earlier intrusion at 

20- to 30-km depth. Regardless of the possible existence of 

a proto-Zayante fault, reactivation of a south-dipping structure by north-southward compression in the middle Tertiary 

would account both for the northward-directed thrusting 

indicated by geophysical models and for the absence of 

gabbroic basement south of the Zayante fault in the Santa 

Cruz Mountains.

CRUST NORTHEAST OF

THE SAN ANDREAS FAULT

In previous sections, we have described the four geophysically 

defined crustal blocks and their three-dimensional structure 

northeast of the San Andreas fault. Here, we briefly describe 

the structure of the contact between the Franciscan Complex 

and the Great Valley sequence where it has been mapped along 

the east side of the California Coast Ranges, emphasizing 

the recent discovery of eastward-tapering tectonic wedges 

that underlie the western parts of the Great Valley sequence 

(Wentworth and others, 1984; Wentworth and Zoback, 1990). 

Then, we show that the structures deduced in the study area 

(see pl. 3A) are largely similar to those deduced to be caused 

by tectonic wedging, and we conclude that the part of the 

study area between the San Andreas and Silver Creek faults 

consists of tilted and broadly folded wedge structures presently 

exposed at the surface.

The contact between the Franciscan Complex (a subduction complex) and the Great Valley sequence (a forearc-basin 

complex) extends for 600 km along the east side of the Coast 

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E68 EARTHQUAKE OCCURRENCE

Ranges and passes about 40 km east of the study area (see 

pl. 3A). This contact is a fault, situated at the base of the 

fragmented Coast Range ophiolite, which forms the former 

oceanic basement of forearc-basin strata of the Great Valley 

sequence. This fault, according to plate-tectonic theory, was 

originally believed to represent an exposed east-dipping subduction zone (Hamilton, 1969; Ernst, 1970). More recently, 

seismic reflection and refraction data (Wentworth and others, 

1984; Walter, 1990; Wentworth and Zoback, 1990), together 

with gravity and magnetic data (Griscom and Jachens, 1990), 

have shown that this contact is rather more complex. In the 

subsurface below and east of the contact are a series of two or 

more eastward-tapering wedges, with dimensions measured in 

kilometers, probably composed of both Franciscan rocks and 

locally, to the east, Great Valley rocks thrust eastward into (that 

is, intruding) rocks of the forearc basin. Geophysical evidence 

also shows that the basement of the Great Valley sequence 

dips relatively smoothly west from the middle of the valley 

at least as far as the Franciscan-Great Valley contact, where 

the basement lies commonly at 10- to 15-km depth. Thus, 

the previously inferred subduction zone associated with the 

accretionary Franciscan Complex must lie still farther west, 

because tectonic wedging has transported Franciscan rocks 

eastward over the Great Valley basement, which necessarily 

lies east of the subduction zone.

A characteristic feature of the Franciscan-Great Valley 

contact is a magnetic anomaly, 15 to 30 km wide, that extends 

nearly continuously for 600 km along the contact (Griscom, 

1966, 1983; Griscom and Jachens, 1990). The anomaly is 

caused by a huge subhorizontal slab composed of the magnetic 

Coast Range ophiolite (15–30 km wide, max 4 km thick) that 

extends along the contact, mostly concealed at shallow (<5 km) 

depth and generally lying beneath the much smaller exposed 

and fragmented part of the Coast Range ophiolite. The western 

part of this slab always is structurally embedded within the 

Franciscan Complex (unit KJf, pl. 3), whereas the east edge 

of the slab approximately contacts the Coast Range ophiolite 

(unit Jou) at the base of the Great Valley sequence. We have 

modeled this slab (fig. 10) from magnetic data obtained near 

Coalinga (Griscom and Jachens, 1990) and interpret it to be 

caused by two successive wedges of Franciscan Complex that 

advanced eastward along or near the top of the Great Valley 

basement, the slab having been peeled up from the Great 

Valley basement by the advancing lower wedge. We infer that 

the 600-km length of this magnetic slab may have resulted 

from the wedge beheading a preexisting elongate parallel-tobasin structure, such as a faulted antiform, on the floor of the 

Great Valley basin and that eastward movement of the two 

wedges was subsequently halted by rapid shallowing of the 

basin to the east. Regionally, the slab is cut off obliquely at 

its southeast end (lat approx 35°30' N.) by the San Andreas 

fault, which has sliced off a major part of the Franciscan-Great 

Valley contact at the south end of the Great Valley.

In the Loma Prieta region (see pl. 3A), the block structure 

between the San Andreas fault and magnetic boundary 1 (fig. 

8) is remarkably similar to the structure of the magnetic slab 

described in the previous paragraphs. Structural similarities 

include the facts that (1) a substantial magnetic anomaly is 

associated with the block, caused by a large, thick, west-dipping magnetic slab; (2) the western part of the slab (fig. 8) is 

surrounded by the Franciscan Complex, yet its northeast edge 

contacts the Coast Range ophiolite which is basement for the 

Great Valley sequence; (3) the slab has dimensions (in cross 

section) similar to those of the slab shown in figure 10; and 

(4) the geologic and structural associations of the two slabs 

are similar. In figure 10, we show which part of our 1990 cross 

section is similar to figure 8 by schematically adding the San 

Andreas and Silver Creek faults at the locations corresponding 

to the structures shown in figure 8. The schematic topographic 

surface shown in figure 8 is also added to figure 10 to demonstrate that our corresponding area near Loma Prieta has been 

tilted westward and antiformally folded, in good agreement 

with conclusions reached in the field.

We conclude that magnetic boundary 1 is a wedge fault lying 

along the base of the thick, west-dipping magnetic slab and 

marking the top of a wedge composed of Franciscan Complex 

beneath the slab. A second wedge composed of Franciscan 

Complex (fig. 8) lies above this magnetic slab and beneath 

the Great Valley sequence. Magnetic boundary 1 in plan view 

to the southeast extends into the area of Franciscan Complex 

because here the magnetic slab is structurally lower, that is, 

embedded in the Franciscan rocks, such that the acute angle 

between magnetic boundary 1 and the exposed base of the 

Great Valley sequence identifies the tip of this upper wedge. 

This feature (the acute angle) we consider to be one of the 

most important geophysical and geologic clues establishing the existence of wedge structures in this area. Similarly, 

magnetic boundary 1A can then be explained as another, less 

important wedge boundary (see magnetic slab outcropping at 

km 23 on magnetic profile along line D–D', fig. 7B) that may 

have resulted from a minor shift in the trace of the wedge 

fault of magnetic boundary 1, and so the acute angle between 

magnetic boundaries 1 and 1A is interpreted to mark the 

side edge of a wedge, assuming that the wedges were thrust 

northeastward. Another wedge boundary is located at and 

northeast of magnetic boundary 3, where a now-horizontal 

wedge fault is interpreted to lie at the base of the horizontal 

sheet of serpentinite (unit Jou, pl. 3) north of magnetic boundary 3 (see fig. 8B).

Our geophysical method is, of course, able to define only 

wedges that have magnetic serpentinite along their boundaries, 

and so other wedge faults may remain unrecognized, including those within the Great Valley sequence. Other geologic 

boundaries within the Franciscan Complex are major shear 

zones or terrane boundaries that probably developed during 

formation of the Mesozoic subduction complex or during 

right-lateral faulting in the early Tertiary (McLaughlin and 

others, 1988a), although the positions of wedges may be 

partly controlled by preexisting terrane boundaries. We assign 

magnetic boundary 2 to this category of earlier boundaries 
E69

formed primarily before wedging (although a postwedging 

change in dip occurs here), and magnetic boundary 3 primarily 

to postwedging offset of the wedge fault.

The geophysical evidence for tectonic wedges in this part 

of the study area (see pl. 3C) is of great interest because, so 

far, most previously identified wedge structures elsewhere 

are concealed beneath the surface and have been inferred 

primarily from seismic data. We conclude here that the tilted 

rocks southwest of magnetic boundary 2 contain a series of 

major wedges and wedge boundaries and that these structures 

are substantially exposed at the surface, offering a splendid 

opportunity for observation and study. Because the relative 

thrust motion on wedge boundaries is in opposite directions 

at the top or bottom of a wedge (fig. 10), mapping of smallscale structures may help determine whether a specific fault 

forms the top or bottom of a wedge.

Most of the large amount of wedging inferred to have taken 

place in the study area (see pl. 3A) must predate the profound 

early Miocene erosion (McLaughlin and others, 1991) that 

exposed the wedged rocks. Early Miocene subsidence and 

deposition then covered the unconformity with flat-lying 

Miocene marine sedimentary strata. Initial movement on the 

San Andreas fault in earliest Miocene time (summarized by 

Stanley, 1987) began faulting these Great Valley and Franciscan rocks away from their original position on the west 

side of the southern Great Valley. Somewhat more recently, 

a strand of the San Andreas fault must here for a while have 

branched off eastward of its present position, probably along 

the Calaveras and Silver Creek faults, and moved this block 

of Great Valley and Franciscan rocks to their present position 

between the Silver Creek and San Andreas faults.

The continuities of magnetic anomalies, curvilinear magnetic boundaries, and geologic units seem to preclude that 

any major linear strike-slip faults of the San Andreas system 

may have cut across the area between the San Andreas and 

Silver Creek faults. Even the Sargent fault (see pl. 3B), whose 

length and linearity suggest a possible exception, appears to 

cut across magnetic gradients on both sides of the hole in the 

magnetic slab southwest of magnetic boundary 1 without 

laterally offsetting the magnetic boundary by more than 3 

km. Furthermore, although we are aware of relatively recent 

thrust faulting, reactivation of wedge faults, and possible 

transpression within the two blocks to the southeast of magnetic boundary 2 (see McLaughlin and Clark, this chapter), 

presumably in response to local stresses created by the San 

Andreas fault, we infer these movements to be relatively minor 

Figure 10.—Geologic cross section across contact between the Franciscan Complex (KJf) and the Great Valley sequence 

(KJs) at lat approx 36°10 ́ N. Two wedges of Franciscan Complex have successively intruded eastward along basement of the 

Coast Range ophiolite and thus isolated an intervening slab of ophiolite. Part of this contact is now represented by geologic 

relations between the San Andreas fault (SAF) and the Silver Creek fault (SCF) in the San Jose area (see pl. 3). Adapted from 

magnetic model of Griscom and Jachens (1990, fig. 5.3). Solid lines, faults; dashed lines, sedimentary contacts. Short heavy 

bars (labeled in parentheses) denote positions of SAF and SCF shown in magnetic model along line E–E' (fig. 8B). Shape 

of dotted line, which denotes present-day erosional surface (topography) along line E–E', indicates that corresponding area 

near San Jose (which has relatively flat topography at this scale) has been tilted westward and anticlinally folded. Vertical 

exaggeration, ×2. gb, gabbro; gr, granitic and older metamorphic rocks; Jou, Jurassic ultramafic rocks and serpentinite of 

the Coast Range ophiolite; KJf, Cretaceous and Jurassic Franciscan complex; KJs, Cretaceous and Jurassic sedimentary 

rocks; Ts, Tertiary sedimentary rocks.

DISTANCE, IN KILOMETERS

10

0

DEPTH,

IN KILOMETERS

0 10 20 30 40 50

20

KJf

Jou

Jou

Jou

KJf

Ts

KJs

(SAF)

(SCF)

San Andreas fault

West

gr

East

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E70 EARTHQUAKE OCCURRENCE

in comparison with the amount of wedge faulting that took 

place before Miocene time.

Finally, we call attention to a curious problem observed 

along magnetic boundary 1 and its lesser extension, magnetic boundary 1A. Although the magnetic anomalies and 

models indicate that an essentially continuous magnetic mass, 

presumably of serpentinite, is present along the southwest 

side of these magnetic boundaries, the geologic map (pl. 

3A) shows only discontinuous irregular exposures of this 

magnetic rock, none especially wide in comparison with the 

model dimensions. In fact, the southeastward extension of 

magnetic boundary 1 displays no associated serpentinite at 

all, even though the magnetic feature is exceptionally strong 

there and the magnetic gradients indicate a large, continuous magnetic mass within about 0.2 km of the surface. The 

question therefore arises, why is so much of the serpentinite 

only a short distance below the present irregular erosional 

surface, a surface that moves downward with erosion over 

time, such that we would expect a continuous belt of these 

magnetic rocks to have been exposed after a relatively small 

amount of erosion? We speculate that when a dipping layer of 

low-density, well-sheared serpentinite is exposed by erosion 

at the surface, the upper few hundred meters may, in some 

relatively unconfined environments, expand and become even 

less dense, leading to instability. This instability, driven by 

the surrounding denser country rock, results in extrusion of 

the serpentinite to the surface in a relatively short time and 

its subsequent erosion. In this model, the voids left by the 

extruded serpentinite would simply collapse. Thus, by means 

of such an intermittent dynamic process, at any moment in the 

erosional history only parts of a serpentinite belt might actually crop out at the surface, even though the other unexposed 

parts lie only a short distance below the surface.

CRUSTAL STRUCTURE AND SEISMICITY

As discussed in this section, the structural details provided 

by potential-field interpretations can aid the interpretation of 

seismicity. We compare the crustal model calculated during 

this study with the distribution of aftershocks in an attempt 

to determine which faults moved during the earthquake. 

The comparison is based mainly on pattern recognition 

(separating the aftershocks into subsets that seem to define 

coherent, approximately planar distributions) and the spatial 

coincidence of the distribution of hypocenters with inferred 

subsurface faults or other block boundaries. As a result of this 

comparison, we argue that most aftershocks are associated 

with three main faults: the Zayante, San Andreas, and Sargent 

faults. Within the context of this discussion, the Zayante and 

San Andreas faults are the geologic boundaries that separate 

rocks of the Logan block from Salinian plutonic rocks to 

the southwest, and from Franciscan and related rocks to the 

northeast, respectively. Furthermore, in many places, the 

aftershocks that we associate with a particular fault occupy a 

relatively wide (max 2 km) zone and evidently include events 

both on the fault plane and in the volume immediately surrounding the fault plane, a volume such as the San Andreas 

fault zone that studies of exhumed major faults show to be 

highly deformed.

The distribution of aftershocks for the period October 18, 

1989, through 1991 is plotted on a base that includes major 

faults (fig. 11). Figures 12 through 16 show various modeled 

cross sections presented previously in this paper, along with 

the locations of hypocenters of aftershocks that occurred 

within 2.5 km of each cross section. Only well-located aftershocks (horizontal uncertainty, <1.5 km; vertical uncertainty, 

<2.5 km) of M>1.0 recorded by eight or more stations are 

shown. The uncertainties do not reflect possible systematic 

errors introduced by limitations in the velocity model used to 

locate the earthquakes. A comparison between the aftershock 

data set mapped in figure 11 and a smaller set located by using 

a three-dimensional velocity model (D.M. Eberhart-Phillips, 

written commun., 1993) shows that the main effect of using a 

more realistic crustal velocity model is to systematically shift 

the hypocenters about 1 km southwestward, normal to the San 

Andreas fault. This shift is most pronounced for aftershocks 

located northeast of the San Andreas fault or beneath the La 

Honda Basin.

Aftershocks that we identify as associated with the Zayante 

fault are shown on our map inside an elliptical region (Z, fig. 

11). The epicenters for this group of aftershocks plot as far as 

5 km southwest of the surface trace of the Zayante fault but 

close to the inferred southwest limit of the gabbro of Logan, a 

limit that we believe marks the position of the southwest-dipping Zayante fault at midcrustal depths (for example, figs. 4B, 

7). Hypocenters superposed on the magnetic profile along line 

D–D', together with a schematic representation of the Zayante 

fault from the adjacent isostatic-residual-gravity profile along 

line R–R' (fig. 1) where the fault geometry is tightly constrained 

by the gravity data, are plotted in figure 12B. Aftershocks that 

we identify as associated with the Sargent fault zone are shown 

inside two elliptical regions (S, fig. 11). The Sargent fault has, 

at most, only a very weak geophysical expression, and so we 

were unable to infer its subsurface configuration on the basis 

of gravity or magnetic anomalies. Therefore, our association of 

these aftershocks with the Sargent fault zone is based simply 

on the distribution of hypocenters in relation to the mapped 

fault trace (figs. 13, 14). By this analysis, the Sargent fault dips 

steeply southwest and was active along only part of its length. 

Interestingly, the gap between the two clusters of aftershocks 

corresponds to a marked change in the Sargent fault trace, 

from long, straight, seismically active segments at the ends to a 

seismically quiet central segment where the fault is geomorphically less distinct and appears to be irregular in strike (fig. 11).

Region SA (fig. 11) includes those aftershocks that we tentatively associate with the San Andreas fault. In the southern 

part of this region, the hypocenters define a band that dips 

steeply southwest and projects to the surface close the San 

Andreas fault trace (fig. 15). The apparent southwestward dip 
E71

of the aftershock zone could be an artifact resulting from the 

use of an incorrect velocity model to locate the hypocenters 

(W.L. Ellsworth, written commun., 1992). However, the facts 

that the magnetic model (fig. 15) indicates a similar dip to the 

fault-bounded northeast face of the gabbro of Logan (well 

constrained only in the top 3 km), and that hypocenters in the 

southeastern part of the area of figure 11 located with a threedimensional velocity model suggest a similar dip (Michael 

and Eberhart-Phillips, 1991) and show little change from the 

locations shown in figure 11 (D.M. Eberhart-Phillips, written 

commun., 1993), argue against such an interpretation. Farther 

northwest along the fault, the distribution of epicenters widens, 

with epicenters plotting on both sides of the San Andreas fault 

trace. Those epicenters that plot southwest of the fault mainly 

represent deep-focus earthquakes that define a southwestdipping planar zone which intersects the vertical projection 

of the San Andreas fault trace at 8- to 10-km depth, approximately the inferred base of the gabbro of Logan (fig. 13).

As discussed in the subsection above entitled “Block 

Boundaries Southwest of the San Andreas Fault,” magnetic 

data indicate that magnetic sedimentary rocks of the Purisima Formation in the Logan block extend in the shallow 

subsurface as far as 1 km northeastward of the San Andreas 

fault trace (see pl. 3A; fig. 8). Therefore, the shallowest 

part of the San Andreas fault must here dip northeast, even 

though aftershocks indicate that the main section of the fault 

dips steeply southwest. A consequence of this shallow northeastward dip is that locally the upper part of the main San 

Figure 11.—Loma Prieta region, Calif., showing distribution of aftershocks (pluses) and locations of major faults, and 

profiles discussed in figures 12 through 16.

0 10 20 KILOMETERS

37°30’

37°15’

37°00’

36°45’

122°15’ 122°00’ 121°45’ 121°30’

SARGENT FAULT

ZAYANTE FAULT

Q

Q'

E

E'

S

S'

D

D'

D''

Northeast limit

of gabbro of Logan block

Minimum southwest

limit of gabbro of Logan block

?

?

?

?

SAN ANDREAS FAULT

PACIFIC OCEAN

T

T'

Probable southwest

limit of gabbro of Logan block

Z

S

S

SA

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E72 EARTHQUAKE OCCURRENCE

Andreas fault surface must lie northeast of the fault trace, in 

a position that could reasonably explain some of the shallow 

aftershocks northeast of the fault trace as seismicity associated with the main section of the San Andreas fault (fig. 13).

Of the rest of the aftershocks unaccounted for in the above 

discussion, a few could reflect movement on the poorly 

defined north boundary of the Logan block (fig 11), whereas 

others northeast of the San Andreas fault in the northern part 

of the map area (fig. 11) probably are associated with a suite 

of southwest-dipping thrust faults (for example, the Monte 

Vista, the northern section of the Berrocal, the Shannon, and 

related faults; Sorg and McLaughlin, 1975; see McLaughlin 

and Clark, this chapter), although our modeling does not extend 

far enough northward to relate block boundaries to seismicity. 

The aftershocks located northeast of the San Andreas fault 

about 5 km northwest of the magnetic profile along line D–D'

(fig. 12A), and the north-dipping zone of aftershocks that 

lies 2 to 5 km southeast of the southwest end of the magnetic 

Figure 12.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) 

along line D'–D'' (pl. 3) across the Logan block and area northeast of the San Andreas fault. Pluses along top of 

magnetic model, locations of magnetic observations; pluses in body of magnetic model, aftershocks that occurred 

within 2.5 km of line D'–D''. Numbers on units are magnetizations (in electromagnetic units per cubic centimeter). 

Circled numbers, magnetic boundaries. No vertical exaggeration. gb, gabbro; gr, granitic and older metamorphic 

rocks; Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range ophiolite; KJf, Cretaceous and Jurassic 

Franciscan complex; KJs, Cretaceous and Jurassic sedimentary rocks; Tp, Pliocene Purisima Formation; Ts, 

Tertiary sedimentary rocks. -200 0 200 MAGNETIC ANOMALY, IN NANOTESLAS 10 5 0 DEPTH, IN KILOMETERS 20 15

Tp

Jou

0.0004

0.0006 0.0005

KJf

KJf

KJf

gr Ts

gr?

gb 0.002

0.00025

Shannon fault

Lime Kiln fault

Berrocal fault

Aldercroft fault

San Andreas fault

Zayante fault

1 1A 2 3

Jou KJf

Ts+KJs

?

?

?

D' D''

A

B

0 5 10 15 20 25 30 35

DISTANCE, IN KILOMETERS
E73

profile along line E–E' (fig. 13B), do not correlate with any of 

the structures defined by our gravity and magnetic modeling.

Finally, we point out an apparent relation between the distribution of aftershocks and the crustal structure for which 

we have no good explanation. Analysis of the distribution of 

aftershocks with depth relative to the inferred crustal structure 

of the Logan block (figs. 12–16) shows that the gabbroic basement of this block is mostly devoid of seismicity and that the 

major southwest-dipping planar distribution of aftershocks 

effectively passes just beneath the lower northeast edge of the 

block. Although seismicity on the San Andreas fault must lie 

at the eastern margin of the gabbroic basement, we believe that 

Figure 13.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) along line 

E–E' (pl. 3) across the Logan block and area northeast of the San Andreas fault. Pluses along top of magnetic model, 

locations of magnetic observations; pluses in body of magnetic model, aftershocks that occurred within 2.5 km of line 

E–E'. Numbers on units are magnetizations (in electromagnetic units per cubic centimeter). No vertical exaggeration. 

gb, gabbro; gr, granitic and older metamorphic rocks; Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range 

ophiolite; KJf, Cretaceous and Jurassic Franciscan complex; KJs, Cretaceous and Jurassic sedimentary rocks; Tp, Pliocene 

Purisima Formation; Ts, Tertiary sedimentary rocks. Circled numbers, magnetic boundaries. -300 -100 100 300 MAGNETIC ANOMALY, IN NANOTESLAS 10 5 0 DEPTH, IN KILOMETER

S

20 15

Zayante fault

San Andreas fault

Shannon fault

Silver Creek 

fault

? ?

gr

Ts

Ts+KJs

KJf

KJf

Jou

0.0002

0.0003 0.0002

Tp

gb 0.002

0.0004

0.0004

0.0017

0.0015

Jou

gr?

KJf

Sargent f.

E E'

A

B

1 2 3

0 5 10 15 20 25 30 35

DISTANCE, IN KILOMETERS

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
E74 EARTHQUAKE OCCURRENCE

the aftershock patterns shown in figure 13 (where the zone of 

aftershocks appears to change dip at the depth of the gabbroic 

basement) and in figure 16 (where the aftershocks are distributed throughout the volume of crust below the gabbroic basement from approx 16-km depth upward to the inferred base of 

the gabbro, but not into the gabbro except at its lower northeast 

corner) indicate that the basement block exerted some control 

over the seismicity. (Recall that the northeast edge of the basement of the Logan block, which coincides with the San Andreas 

fault at depth, is poorly constrained by magnetic data along the 

reach where this body abuts the highly magnetic slab northeast 

of the San Andreas fault and southwest of magnetic boundary 

1.) A possibly related observation (see pl. 3B) is that the north 

end of the constraining “Loma Prieta” bend in the San Andreas 

fault coincides with the northernmost contact of the basement 

of the gabbro of Logan with the San Andreas fault and that this 

basement abuts the fault along about 70 percent of the length of 

this bend. Although it is difficult to understand how a gabbroic 

basement block that in places is only a few kilometers thick 

could exert a major influence on the local behavior of the San 

Andreas fault, the crust of the Logan block below the gabbro 

may differ sufficiently from the lower crust in the surrounding 

blocks to account for the seeming influence of the gabbro.

CONCLUSIONS

Our analysis of gravity and magnetic anomalies, coupled 

with firm geologic constraints, produces a structural and compositional image of the midcrust and upper crust in the Loma 

Prieta region that helps define the major structural elements in 

the upper 10 km of the crust, provides clues and constraints on 

the tectonic development of the region, and suggests a framework for associating most aftershocks with known major faults. 

This crustal model is limited both by being focused on those 

features that are expressed mainly in the gravity and magnetic 

Figure 14.—Distribution of aftershocks (pluses) that occurred within 2.5 km of line T–T'. Note cluster of aftershocks 

beneath trace of the Sargent fault.

 0 5 10 15 20

 DISTANCE, IN KILOMETERS 20 15 10 5 0

T T'

DEPTH, 

IN KILOMETERS

San Andreas fault

Sargent fault
E75

fields and by being confined to approximately the upper 10 

km of the crust by the resolution limits of these techniques. In 

many places, movements on major faults, that in the past have 

controlled the tectonic evolution of the region and that today 

are responsible for the local seismicity, have juxtaposed rocks 

with different densities and (or) magnetic properties. The gravity and magnetic anomalies produced by these juxtapositions 

have served as the basis for our analysis.

The crust in the Loma Prieta region is divided into seven 

major basement blocks on the basis of characteristic geology, bounding structures, gravity and magnetic signatures, 

and internal structural style. Many of these blocks probably 

represent tectonostratigraphic terranes. From southwest to 

northeast, the blocks are defined as follows.

1. The Ben Lomond block, bounded on the southwest by the 

northeast-dipping San Gregorio fault and on the northeast 

by the southwest-dipping Zayante fault, is floored mostly by 

Cretaceous plutonic rocks that are exposed in the northern 

Gabilan Range and the Ben Lomond pluton (see pls. 3B, 

3C).

2. The Logan block, bounded by the Zayante fault on the 

southwest, by the San Andreas fault on the northeast, by the 

Figure 15.—Magnetic profile (A) (circles, measured data; curve, calculated values) and magnetic model (B) along line S–S' (pl. 3) across 

the Logan block between the Zayante and San Andreas faults. Pluses along top of model, locations of magnetic observations; pluses in body 

of model, aftershocks that occurred within 2.5 km of line S–S'. Numbers on units are magnetizations (in electromagnetic units per cubic 

centimeter). No vertical exaggeration. gb, gabbro: Jou, Jurassic ultramafic rocks and serpentinite of the Coast Range ophiolite. Circled 1, 

magnetic boundary 1.

GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE

-200

0 200

MAGNETIC ANOMALY,

IN NANOTESLAS

10

5

0

DEPTH,

IN KILOMETERS

gb 0.0011

gb

0.0008

0.0005

0.0007

Jou

(schematic only)

San Andreas fault

1

S S'

A

B

 0 5 10 15 20 25 30 35

DISTANCE, IN KILOMETERS
E76 EARTHQUAKE OCCURRENCE

San Gregorio fault at its extreme northwest end, and by an 

unknown, completely concealed structure, probably a fault, 

on the north, is characterized by a thick section of Tertiary 

sedimentary rocks (La Honda Basin) resting on a gabbroic 

basement with affinities to rocks exposed at the south end 

of the Sierra Nevada batholith. Plutonic rocks of the Ben 

Lomond block were thrust northeastward over rocks of the 

Logan block along the Zayante fault, with total convergence 

of at least 10 km in places. We propose that the Zayante fault 

is a reactivated fault possibly left over from the suturing 

event which attached the basement of the Logan block to the 

North American Continent, or, alternatively, is a pre-Tertiary 

fault which juxtaposed the gabbro of Logan against more 

felsic crystalline rocks by large-scale, dominantly dip slip 

movement. This proto-Zayante fault may have as its counterpart the south-dipping contact (as interpreted from magnetic 

data) between the gabbro of Eagle Rest Peak and overlying 

granitic and high-grade metamorphic rocks of the southern 

Sierra Nevada. Although geologic evidence indicates that 

most movement on the Zayante fault had ended by early 

Miocene time, the spatial coincidence between some aftershocks and the inferred subsurface position of the Zayante 

fault suggests some ongoing movement. Most aftershocks 

define a southwest-dipping, crudely planar rupture zone 

that probably reflects the position of the main section of the 

San Andreas fault at depth. In contrast, magnetic anomalies 

Figure 16.—Isostatic-residual-gravity profile (A) (circles, measured data sampled from gridded gravity map; curve, calculated values) and 

density model (B) along line Q–Q' (pl. 3) across the Logan block between the Zayante and San Andreas faults. Pluses along top of density 

model, locations where gravity was calculated; pluses in body of density model, aftershocks that occurred within 2.5 km of line Q–Q'. 

Numbers on units are densities (in grams per cubic centimeter). No vertical exaggeration. gb, gabbro; gr, granitic and older metamorphic 

rocks; KJf, Cretaceous and Jurassic Franciscan complex; Qal, Quaternary alluvium; Ts, Tertiary sedimentary rocks. -40 -20 0 20 40 GRAVITY ANOMALY, IN MILLIGALS 15 10 5 0 DEPTH, IN KILOMETERS

Ts 2.37

Ts 2.48

 Ts 2.55

Ts 2.61

gr 2.67 Ts 2.64

gb 2.91

KJf 2.68

Qal+Ts 2.17

(schematic only)

Qal 2.04

2.87

gr? 2.67

San Andreas fault

gr? 2.67

Q Q'

A

B

?Zayante fault

0 5 10 15 20 25 30 35 40

DISTANCE, IN KILOMETERS
E77

over younger sedimentary rocks of the Logan block cut by 

the San Andreas fault indicate that locally the uppermost 

part of the fault must dip gently northeast. These two facts 

indicate that, in places, a cross section normal to the fault 

would show a curved San Andreas fault surface, with 

the deepest part of the seismically active fault projecting 

from southwest of the trace toward the surface at a point 

northeast of the trace, and that only in the uppermost part 

of the cross section would the fault surface bend backward 

to the southwest to emerge at the mapped fault trace. We 

speculate that the San Andreas fault has this shape in the 

central part of the Loma Prieta “bend” and that some of the 

shallow aftershocks which appear to be located northeast 

of the fault actually are associated with the main section 

of the San Andreas fault at shallow depth.

3. A largely unknown block north of the Logan block and west 

of the San Andreas fault may be floored by Cretaceous 

plutonic rocks similar to those of the Montara pluton at 

lat 37°30' N.

4. A block immediately northeast of the San Andreas fault 

has its northeast boundary defined primarily by the northeast border of a strong magnetic anomaly. We interpret 

the block to be made up of large, coherent packages of 

Franciscan rocks, Great Valley sequence, and overlying 

younger strata and to be bounded at its base by a thick, 

southwest-dipping slab of serpentinite of the Coast Range 

ophiolite. Because the normal, undeformed structural 

position of the Coast Range ophiolite is at the base of 

the Great Valley sequence, and because of similarities 

between the structure of this block and that of the Franciscan-Great Valley contact on the east side of the Coast 

Ranges, we interpret the presence of Franciscan rocks 

between Coast Range ophiolite and Great Valley sequence 

as evidence for past tectonic wedging by Franciscan 

rocks into the Coast Range-Great Valley contact. This 

block, after wedging, was transported northward from the 

southern Great Valley by strike-slip movement on the San 

Andreas fault system, which for part of the time included 

the Calaveras and Silver Creek faults. The Sargent fault 

cuts vertically through this block and into the block below. 

Although many aftershocks appear to be associated with 

movement on the Sargent fault, its failure to noticeably 

offset magnetic gradients caused by the underlying Coast 

Range ophiolite suggests that cumulative strike-slip offset 

on this fault totals no more than a few kilometers.

5. A block between magnetic boundaries 1 and 2 is composed 

of Franciscan rocks of the Permanente and Bald MountainEl Sombroso terranes. The northeast limit of the block 

correlates with a shear zone of melange that also marks a 

change in structural style, from southwest-dipping units on 

the southwest to regionally flat lying units on the northeast. 

This block is denser than most Franciscan rocks, and it 

causes a gravity high.

6. A block between magnetic boundaries 2 and 3 is composed 

of Franciscan volcanic and sedimentary rocks with abundant small, isolated masses of serpentinite. This block is 

defined mostly on the basis of its characteristic magnetic 

signature, generated by serpentinite. The block is composed 

of regionally flat lying units, an inference based on the 

presence of thin serpentinite masses and the observation 

that the large masses of dense volcanic rocks fail to cause 

gravity highs.

7. A block immediately southwest of the southwest-dipping 

Silver Creek fault is composed predominantly of magnetic 

serpentinite with mafic components of the Coast Range 

ophiolite, together with sedimentary rocks of the Franciscan Complex and some sedimentary rocks of the Great 

Valley sequence. Magnetic modeling indicates that the 

block contains a nearly continuous, subhorizontal sheet of 

serpentinite which extends across the Santa Clara Valley 

beneath a thin cover of alluvium. The sheet may once have 

extended southwestward of magnetic boundary 3 to magnetic 

boundary 2 but has since been faulted at magnetic boundary 3 and eroded, leaving only thin patches of serpentinite 

southwest of the fault. The contact between this serpentinite 

sheet and underlying Franciscan rocks is interpreted to be 

a wedge fault.

Comparison of this crustal model with the distribution of 

aftershocks suggests that most aftershocks were associated 

with movement on three major faults—the Zayante, San 

Andreas, and Sargent faults—and that the gabbroic body 

which composes the basement of the Logan block is mostly 

devoid of aftershocks, in marked contrast to the surrounding 

rocks. This gabbroic body appears to have played a role in 

controlling the distribution of aftershocks, although the precise 

mechanism is unknown.

ACKNOWLEDGMENTS

This study benefited greatly from extensive discussions with 

Bob McLaughlin about the geology of the Santa Cruz Mountains northeast of the San Andreas fault, and from numerous 

other discussions about the general geology of the study area 

with Clark Blake, Joe Clark, Bob Coleman, Carl Wentworth, 

and many other of our USGS colleagues. Dave Oppenheimer 

generously supplied the seismicity catalog covering the aftershock sequence, along with the software needed to manipulate 

it. We thank Bill Ellsworth, Bob McLaughlin, Dave Ponce, and 

Carl Wentworth for their reviews of the manuscript.

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GEOPHYSICAL AND GEOLOGIC SETTING OF THE EARTHQUAKE
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CONTENTS

————— Page

Abstract----------------------------------------------------------------------- E81

Introduction------------------------------------------------------------------ 81

Magnetotelluric coverage--------------------------------------------------- 82

Magnetotelluric models----------------------------------------------------- 82

 Loma Prieta profile---------------------------------------------------- 82

 Gabilan Range profile------------------------------------------------- 83

Discussion of models-------------------------------------------------------- 83 

 Loma Prieta profile---------------------------------------------------- 83 

 Gabilan Range profile------------------------------------------------- 89 

Conclusion-------------------------------------------------------------------- 91 

References cited------------------------------------------------------------- 91 

Appendixes:

 1. Magnetotelluric methodology------------------------------------- 92 

 2. Magnetotelluric coverage------------------------------------------ 93 

 3. Magnetotelluric models-------------------------------------------- 93 

 Loma Prieta profile---------------------------------------------- 93 

 Gabilan Range profile------------------------------------------- 96 

 —————

ABSTRACT

A magnetotelluric-sounding profile in the Loma Prieta 

region was completed before the earthquake. An electricalresistivity model calculated for the magnetotelluric profile 

indicates (1) a large, high-resistivity structure within 2 km 

of the surface between the San Andreas and Sargent faults 

at seismogenic depths that may have played an important 

role in stress buildup; and (2) a region of anomalously low 

resistivity at midcrustal depths near the San Andreas fault 

that may have been caused by fluids in the deeper parts of 

the fault-gouge zone, which could have reduced the shear 

stress required for sliding and thus promoted the M=7.1 

earthquake. The anomalously low resistivity region was 

a key and required element of the model, even though the 

model was insensitive to dip, shape, and distribution. The 

structure interpreted from this profile will provide constraints 

on future geologic interpretations of the Loma Prieta region. 

High-resistivity units southwest of the San Andreas fault 

may represent Salinian-block granitoids of Mesozoic age. 

High-resistivity units, below 4-km depth, northeast of the 

San Andreas fault could be related to the Franciscan complex. Low-resistivity rocks between the Zayante and Sargent 

faults in the uppermost 4 to 7 km of crust are interpreted 

from drill-hole data as mostly Tertiary marine sedimentary 

rocks exposed in tightly folded structures in the Loma Prieta 

region. Low-resistivity units, above 4-km depth, northeast 

of the San Andreas fault are believed to represent Mesozoic 

units of the Franciscan assemblage. The resistivity model 

agrees well with the three-dimensional seismic-velocity 

model of the Loma Prieta region derived by using raypaths 

from aftershocks recorded by the U.S. Geological Survey 

seismic network. The resistivity model is also consistent 

with gravity and magnetic models of the Loma Prieta region. 

The San Andreas fault and adjoining subparallel Sargent and 

Zayante faults appear to juxtapose diverse deep geologic 

structures spanning the seismogenic zone in the Loma Prieta 

region, suggesting that local geologic structures may have 

played a role in triggering the earthquake and may control 

the overall seismicity on the San Andreas and Sargent 

faults. Another magnetotelluric-sounding profile about 35 

km southeast, traversing the northwest end of the Gabilan 

Range, shows a distinctly different structure. Low-resistivity 

units there appear to be more spatially uniform and limited 

to the uppermost 4 km of crust. This type of geoelectrical 

structure may indicate stable fault slip on the central section 

of the San Andreas fault.

INTRODUCTION

The Loma Prieta earthquake associated with the San 

Andreas fault was unusual in its depth of focus (18 km) and 

the thrust component evident in first motions. Because the 

Loma Prieta region is complex, it needs interpretative input 

from every applicable field of study if we are ever to really 

understand its structure, dynamics, and tectonic history. Few 

published accounts of geoelectrical measurements across 

the San Andreas fault exist (Mazzella, 1976; Phillips and 

Kuckes, 1983; Mackie and others, 1988; Park and others, 

1991). The absence of geoelectrical results is partly explainable by the difficulty of acquiring data in the region. Deeppenetrating electrical soundings measure extremely small 

magnetic and electrical fields and are prone to interference 

from such cultural features as powerlines, television and 

radio transmitters, and pipelines.

Two magnetotelluric (MT) profiles have been used to 

characterize the deep-crustal electrical structure of the San 

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION

By Brian D. Rodriguez, Victor F. Labson, William D. Stanley, and Jay A. Sampson, U.S. Geological Survey

E81
E82 EARTHQUAKE OCCURRENCE

Andreas fault in the Loma Prieta region and northern Gabilan Range as part of the U.S. Geological Survey (USGS)’s 

National Geologic Mapping and Deep Continental Studies 

programs. The MT method is used to determine the electrical-resistivity structure of the subsurface (see app. 1). 

Two MT profiles that define the subsurface resistivity to 

below 10-km depth indicate the geometry of faulting that is 

complementary to seismic and other geophysical techniques. 

In combination with geologic information and the results 

of other geophysical methods, lithologic distributions may 

be inferred. We hope that our interpretation of the MT data 

will not only lead to a more accurate understanding of the 

faulting mechanism but also assist other workers in their 

efforts to understand how the 1989 Loma Prieta earthquake 

actually fits into the larger scope of studies on the San 

Andreas fault.

MAGNETOTELLURIC COVERAGE

One north-by-northeast profile of MT soundings (Loma 

Prieta profile, fig. 1) near the Loma Prieta region was completed in May 1989. Later, MT coverage on this profile was 

expanded, and a northeast profile (Gabilan Range profile, 

fig. 1) across the northwest end of the Gabilan Range was 

completed in May 1990. Stations were then added to the Loma 

Prieta profile in June 1992 to provide more MT coverage in 

geologically complex areas (see app. 2).

The Loma Prieta profile (fig. 1) was positioned along the 

south end of the locked 1906 segment of the San Andreas 

fault, and the Gabilan Range profile was positioned along 

the transition from the locked segment to stable fault slip on 

the central section of the San Andreas fault near San Juan 

Bautista. A main objective of the study was to investigate 

deep structures associated with seismic-stress release so as 

to aid in understanding earthquake mechanisms along the 

San Andreas fault.

MAGNETOTELLURIC MODELS

LOMA PRIETA PROFILE

The 14 MT soundings on the Loma Prieta profile (fig. 1) 

were modeled with an interactive, finite-element algorithm 

to calculate the tensor resistivities and phases for twodimensional structures (Wannamaker and others, 1987). The 

finite-element grid used in this model consisted of 263 by 40 

variable-dimension cells extending to several hundred kilometers horizontally and 70 km vertically to minimize edge 

effects. The seawater of the Pacific Ocean was simulated at 

the southwest end of the model. We manually perturbed the 

starting two-dimensional model (see app. 3) more than 100 

times; the best-fit model is shown in figure 2.

The gross structure of our model came from fitting the 

transverse-magnetic data (see app. 1). Wannamaker and 

others (1987) showed that finite strike-length structure is 

better resolved with the transverse-magnetic mode and that 

the transverse-electric mode is highly insensitive to gross 

resistivity structure. We noted, however, that we needed to add 

conductance in the depth range 4–16 km near the San Andreas 

fault so as to adequately fit the observed transverse-electric 

data in the low to midband frequencies. Thus, the dip, shape, 

and precise distribution of this added conductance were irresolvable, even though its presence was a major requirement of 

the model. We determined that, in its absence, we had no other 

solution to adequately fit the observed transverse-electric data 

(see app. 3). The deeper conductive zones could conceivably 

have numerous variations and still be consistent with the 

observed data, although our model requires some degree of 

lateral inhomogeneity in resistivity in the uppermost 16 km 

of crust across the Loma Prieta region (see app. 3).

In the uppermost 4 km of crust, resistivities ranged from 

3 to 1,000 Ω-m. The conductive material beneath these stations could vary somewhat in thickness, distribution, and 

depth from what we consider to be our best discrete twodimensional resistivity model, though probably not without 

adversely affecting the computed fit to the observed data. 

The thickness, distribution, and depth of the more resistive 

material, however, are poorly constrained and probably could 

vary somewhat. For example, the thickness of the 100 Ω-m 

material northeast of station 3 (fig. 1) could be increased to at 

least 19-km depth; however, the observed data do not require 

this increase in our models.

The key features of the Loma Prieta MT model are as follows: 

(1) high resistivities (approx 100–1,000 Ω-m) above 

 2-km depth beneath stations 3 and 16 (fig. 1); 

(2) high resistivities (approx 1,000 Ω-m) at 4- to 7-km 

 depth over much of the prof ile southwest of 

 the San Andreas fault and at 10-km depth northeast 

 of the San Andreas fault; 

(3) low resistivities of 3 to 30 Ω-m above 4-km depth 

 southwest of the San Andreas fault and of 1 to 

 100 Ω-m above 10-km depth northeast of the 

 San Andreas fault; and 

(4) low-resistivity (approx 3 Ω-m) bodies in the depth 

 range 4–19 km near the San Andreas fault, with 

 their dip, shape, and precise distribution irresolvable.

The resolution of the resistivity boundaries used for the 

model is somewhat subjective. If different resistivities are 

used, then boundary positions and layer depths would have 

to be adjusted to achieve similar fits to the observed data. The 

extreme case would be to use a model with a “continuous” 

resistivity gradient from low to high resistivities. The resolution of the resistivity boundaries also is partly a function of our 

model grid-mesh design. We have attempted to keep the model 

simple. Our final model is, in our opinion, the best discrete 

two-dimensional model that fits the observed data.
E83

GABILAN RANGE PROFILE

The six MT soundings on the Gabilan Range profile (fig. 

1) were modeled with the same interactive, two-dimensional 

finite-element algorithm as the one used on the Loma Prieta 

profile. The finite-element grid used in this model consisted of 

239 by 48 variable-dimension cells extending to several hundred kilometers horizontally and vertically; the best-fit model 

is shown in figure 3. This model indicates large contrasts in 

resistivity in the uppermost 4 km of crust across this region.

The key features of the Gabilan Range MT model are as 

follows: 

(1) a high-resistivity (1,000 Ω-m) basement at about 

 4-km depth, 

(2) a central low-resistivity (3–10 Ω-m) anomaly to about 

 4- to 7-km depth beneath stations 1, 5, and 6 (fig. 1), 

 and 

(3) an adjacent low-resistivity (1–3 Ω-m) anomaly south- 

 west of station 4 to about 4-km depth.

Like the Loma Prieta MT model, the resolution of the 

resistivity boundaries used for the Gabilan Range MT model 

is somewhat subjective. We again have attempted to keep the 

model simple. Our final model is, in our opinion, the best 

discrete two-dimensional model that fits the observed data.

DISCUSSION OF MODELS

LOMA PRIETA PROFILE

Electrical resistivity can be reduced by fracturing, conductive pore fluids, and alteration minerals. Unaltered, unfractured crystalline rocks normally are highly resistive; marine 

sedimentary rocks and clay-rich alluvium normally are highly 

Figure 1.—Index map of the Loma Prieta region, Calif., showing locations of two magnetotelluric (MT) 

profiles used in resistivity modeling. Numbered dots, stations where MT soundings were made; star, epicenter 

of October 17, 1989, main shock. Y=0 line overlaying Loma Prieta MT profile is projection used for threedimensional velocity-inversion vertical model slice (Eberhart-Phillips and others, 1990) at Y=0 km; Y=35 

line overlaying Gabilan Range MT profile is projection used for Gabilan Range model 35 km southeast of 

Y=0 line. Irregular line with end points is profile of geologic cross section shown in figure 8.

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION
E84 EARTHQUAKE OCCURRENCE

conductive; and metamorphic rocks, including gneiss, schist 

(nongraphitic), and marble, are moderately to highly resistive. 

In general, low resistivity, low density, and low velocity are 

correlative because they all are associated with increasing 

porosity and water content. Exceptions to this behavior in 

the Loma Prieta region may be high-velocity, low-resistivity 

serpentinized mafic rocks.

The three-dimensional velocity model of the Loma Prieta 

region, derived using raypaths from aftershocks recorded by the 

U.S. Geological Survey seismic network (Eberhart-Phillips and 

others, 1990), includes a low-velocity (3.2–5.9 km/s) wedge 

of material between the Zayante and San Andreas faults that 

extends below 8-km depth, and a high-velocity (6.3–6.4 km/s) 

zone at depth between the San Andreas and Sargent faults. 

Some striking similarities in the velocity and resistivity models 

are visible in figure 4. The high-resistivity anomaly in the 

uppermost 10 km of crust beneath the San Andreas and Sargent 

faults coincides with high velocities, whereas the low-resistivity anomaly in the uppermost 8 km of crust between stations 

10 and 5 (fig. 1) coincides with a low-velocity zone. We also 

note that a low-resistivity branch extends into the much higher 

velocity part of the velocity model. The thick, 10- to 100-Ωm-resistivity units northeast of the Sargent fault correspond to 

velocities of less than 6 km/s. Readers of Eberhart-Phillips and 

others (1990) will notice minor differences between our current 

and previously published resistivity models. Reprocessing of 

the original six stations, the inclusion of nine additional stations (station 14 replacing station 7), and improved fits to the 

observed data account for these minor differences.

Although the high-resistivity units southwest of the San 

Andreas fault are believed to represent Salinian-block granitoids of Mesozoic age (Clark and Rietman, 1973), they could 

be any other granitic or metamorphic-rock units. Similar 

resistivities have been measured in granitic rocks worldwide, 

including granitoids in the Gabilan Range (southeast of the 

Loma Prieta region) by Phillips and Kuckes (1983). The 

high-resistivity unit in our model that occurs between the San 

Andreas and Berrocal faults could be related to a highly metamorphosed section of Franciscan complex (Brabb and Hanna, 

1981; McLaughlin and others, 1988); this unit appears to be 

limited in strike length relative to the rest of the geoelectrical 

structure (see app. 2).

Figure 2.—Loma Prieta magnetotelluric (MT) profile (see fig. 1 for locations). A, Two-dimensional finite-element MT model. Numbers 

denote stations shown in figure 1. B, Resistivity data. C, Phase data. Observed data: crosses, transverse electric (TE); circles, transverse 

magnetic (TM). Calculated data: solid curves, TE; dashed curves, TM.

A
E85

We interpret the 3- to 30-Ω-m-resistivity rocks between 

the Zayante and Sargent faults in the uppermost 4 to 7 km of 

crust as mostly Tertiary marine sedimentary rocks, which are 

exposed in tightly folded structures in the Loma Prieta region 

(McLaughlin and others, 1988). Jachens and Griscom (this 

chapter) state that the total Tertiary marine section is about 6 

to 7.5 km thick between the Zayante and San Andreas faults 

beneath our Loma Prieta profile. The Loma Prieta MT model 

shows a conductive (3–30 Ω-m) section, about 4 to 7 km thick; 

however, this section extends only between station 14 (fig. 1) 

and the San Andreas fault (fig. 2). Extending this conductive 

section to the Zayante fault increases the transverse-magnetic 

curve misfit significantly at station 14; however, if we use a 

resistivity of 100 Ω-m in the depth range 4–7 km, extending 

this conductive section to the Zayante fault increases the 

transverse-magnetic curve misfit only slightly at station 14.

Although the conductive material that extends to 19-km 

depth may be partly explainable with altered basement rocks 

and fault gouge, we find it geologically difficult to accept 

3-Ω-m-resistivity material at these depths. However, serpenFigure 2.—Continued

B

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION
E86 EARTHQUAKE OCCURRENCE

tinization of mafic rocks reduces their resistivities to the range 

5–40 Ω-m (Parkhomenko, 1967). Also, Stesky and Brace 

(1973) reported laboratory samples of serpentinite to have a 

steady 3-Ω-m resistivity when subjected to pressures of 0 to 

6 kbars and that the resistivity of a serpentinized gabbro is 

100 to 1,000 times lower than that of a serpentine-free gabbro. 

C.S. Cox (in Stesky and Brace, 1973, p. 7619) inferred a 

resistivity 10 times that of seawater for rocks within 30 km 

of the sea bottom. A 10- to 13-km-thick section of 10-Ω-mresistivity rocks mapped under the east edge of the Gabilan 

Range (southwest of the San Andreas fault) by Phillips and 

Kuckes (1983), using controlled-source electromagnetic 

soundings, was interpreted by them to be either fault gouge 

or serpentinized oceanic crust. Serpentinite, however, is not 

known to exist anywhere in the Salinian block (Jachens and 

Griscom, this chapter). Therefore, resistivities in these deep 

rocks must be almost entirely related to fracturing and fluid 

content (see app. 1).

The 10-, 30-, 100-, and 1,000-Ω-m-resistivity units beneath 

stations 1, 4, and 8 on the Loma Prieta profile (fig. 1) are 

Figure 2.—Continued

C
E87

believed to represent Mesozoic Franciscan rocks, which 

consist of a highly sheared assemblage of sandstone and 

shale, with lesser amounts of chert, limestone, greenstone, 

serpentine, and interbedded basalt flows (Alt and Hyndman, 

1975; Cole and others, 1989). Mazzella (1976) measured 

similar resistivities for these units in the Hollister area (fig. 

1). Velocities of less than 6.0 km/s coincident with this 100-

Ω-m-resistivity section (fig. 4) are compatible with interpreted 

velocities for the Franciscan Formation in similar threedimensional inversions and refraction surveys at Coalinga, 

Calif. (Eberhart-Phillips, 1989; Walter, 1990). The southwest 

edge of the 100-Ω-m-resistivity Franciscan wedge from 4- 

to 10-km depth approximately correlates with the position 

of the Berrocal fault, if its fault plane dips southwest. The 

interpreted wedge also appears to be limited in strike length 

relative to the rest of the geoelectrical structure (see app. 2), 

which we could not determine without additional data for our 

three-dimensional MT modeling. The northeast edge of the 

10-Ω-m-resistivity unit under station 4 (fig. 1) correlates with 

the position of the Chesbro fault. The Santa Teresa fault is not 

apparent in the resistivity data at these depths.

A plot of aftershocks (main-shock section Y=0; Dietz and 

Ellsworth, 1990) along the Loma Prieta MT model cross 

section (fig. 5) shows a southwest-dipping main-shock plane 

below 10-km depth that approximately coincides with the 

deeper parts of the 3-Ω-m-resistivity material below stations 

12 and 14 (fig. 1), but above that depth the seismicity is more 

diffuse and shifts northeastward, with seismicity patterns that 

are spatially associated with the San Andreas and Sargent 

faults. A distinct vertical pattern from 8- to 11-km depth also 

appears to be associated with the Sargent fault.

Given the small contrasts in the gravity and magnetic models 

of Jachens and Griscom (this chapter), we found our resistivity 

model to be consistent with the gross structure of their models 

near our MT profile.

Although the following information is given as an aid in 

correlating the various geologic units with the Loma Prieta 

MT model, the reader should remember that the drill-hole data 

have been projected twice, first onto the geologic cross section 

and then onto the Loma Prieta MT model (fig. 2), which is a 

two-dimensional model of finite-strike-length structures (see 

Jachens and Griscom, this chapter). Thus, the two moderately 

deep (1–2 km) drill holes between stations 10 and 14 (fig. 1) 

and surface geologic data (McLaughlin and others, 1988) 

were used to help geologically constrain the uppermost 2 

km of crust in the Loma Prieta MT model (fig. 6) during our 

modeling efforts. On the basis of the geologic interpretation 

based on the drill-hole data (McLaughlin and others, 1988) 

near stations 10 and 14, the following correlation of resistivity with lithology is generally assumed in our interpretation 

of the MT data: 

 Resistivity Unit (X-m)

Tertiary shale, mudstone, siltstone, 

 and sandstone---------------------------------------------- 3–30

Cretaceous shale, sandstone, and 

 conglomerate---------------------------------------------- 10–100

Jurassic Coast Range ophiolite------------------------------ 3–30

Figure 3.—Gabilan Range magnetotelluric (MT) profile (see fig. 1 for locations). A, Two-dimensional 

finite-element MT model. Numbers denote stations shown in figure 1. B, Resistivity data. C, Phase data. 

Same symbols as in figure 2.

A

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION
E88 EARTHQUAKE OCCURRENCE

Franciscan complex------------------------------------------- 30–1,000

Granitic rocks (quartz diorite)------------------------------- 30–1,000

The resistivities of all these rock types overlap, although 

some rock types appear to favor the low end of the resistivity spectrum (Tertiary and Jurassic), whereas others appear 

to favor the high end (Franciscan and granitic). The reader 

is, therefore, cautioned that our MT model boundaries are 

somewhat insensitive to precise lithologic boundaries.

We find it peculiar that the Salinian-block granitoids 

beneath station 10 (fig. 1) would have resistivities as low as 

30 Ω-m; however, the close proximity of the Zayante fault 

may be responsible for the low resistivity here because of 

fault gouge, altered rocks, and an increase in fluid content in 

the fault zone. This low-resistivity zone also reduces the MT 

model resolution of determining both the depth and lateral 

boundaries of resistive granitoids near the fault. A perplexFigure 3.—Continued.

B

Figure 3.—Continued

C
E89

ing anomaly is the low-resistivity body that apparently dips 

northeast beneath stations 12 and 13, where the surficial 

geology indicates that the subsurface structure should dip 

southwest. Either the geologic extrapolation of McLaughlin 

and others (1988) is invalid here, or the data could indicate 

the uncertainty in determining the lithologic boundaries with 

our resistivity measurements. We also find it peculiar that the 

Butano sandstone northeast of station 11 has a lower resistivity than the same unit southwest of station 11. The close 

proximity of the San Andreas fault zone northeast of station 

11 may be responsible for the lower resistivity because of 

fault gouge, altered rocks, and an increase in fluid content 

within the fault zone; or the data could also indicate the 

uncertainty in determining the lithologic boundaries with 

our resistivity measurements. The close proximity of the 

Sargent fault zone may be responsible for the wide variation 

in resistivities in the Franciscan complex because of fault 

gouge, altered rocks, and an increase in fluid content within 

the fault zone.

The interpreted underlying structures southwest of the 

San Andreas fault evidently have different strain responses 

from the metamorphic rocks interpreted to occur between 

the San Andreas and Sargent faults. This contrast in strain 

response could have played a major role in stress buildup of 

the locked zone before the 1989 Loma Prieta earthquake, and 

it seems to have been instrumental in post-main-shock stress 

release controlling the aftershock distribution. A buildup of 

overpressured fluids in the deeper parts of the fault-gouge 

zones could have reduced the shear stress required for sliding 

to trigger the earthquake (Blanpied and others, 1992).

GABILAN RANGE PROFILE

Profiles across the San Andreas fault at Loma Prieta and in 

the Gabilan Range (fig. 1) show a distinctly different structure, 

especially for the San Andreas fault zone. Low-resistivity 

units appear to be more spatially uniform and limited to the 

uppermost 4 km of crust. This type of geoelectrical structure 

may indicate stable fault slip on the central section of the 

San Andreas fault. The Zayante, Chesbro, and Santa Teresa 

faults were not detected on the Gabilan Range profile probably because the profile did not cross them, a result consistent 

with surface geologic mapping (fig. 1). The 7-km-thick, lowresistivity (10 Ω-m) zone beneath station 6 (fig. 1) appears to 

correlate with a projection of the Sargent fault. The Calaveras 

fault, however, is not apparent in the resistivity data at these 

depths. The 10-km-thick, 100-Ω-m-resistivity Franciscan 

wedge on the Loma Prieta profile is only about 4 km thick 

on the Gabilan Range profile.

On the Gabilan Range profile (fig. 1), the 1- to 100-Ω-mresistivity units in the uppermost 4 km of crust southwest of 

Figure 4.—Loma Prieta magnetotelluric profile (fig. 2; see fig. 1 for locations), with representative seismic 

velocities (in kilometers per second) from three-dimensional velocity-inversion model slice (Eberhart-Phillips 

and others, 1990) superimposed.

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION
E90 EARTHQUAKE OCCURRENCE

Figure 5.—Loma Prieta magnetotelluric profile (fig. 2; see fig. 1 for locations), with locations of 1989 

main shock (star) and aftershocks (circles; size relative to magnitude) within 5 km of Y=0 line.

Figure 6.—Loma Prieta magnetotelluric profile (fig. 2; see fig. 1 for locations), with geologic 

cross section of McLaughlin and others (1988) superimposed. Irregular line at top of profile is 

topography; thick dashed lines are inferred faults. Units: F, Franciscan complex; G, granitic rocks 

(quartz diorite); J, Jurassic Coast Range ophiolite; K, Cretaceous sandstone, shale, and conglomerate 

of Sierra Azul; T, Tertiary marine sandstone, shale, and mottled marine mudstone and sandstone 

of Mount Chual; Tb, Tertiary sandstone and siltstone (Butano sandstone); Tla, Tertiary mudstone, 

siltstone, sandstone (Lambert shale); Tme, Tertiary marine shale and sandstone (Highland Way); 

Tp, Tertiary siltstone and sandstone (Purisima formation); Tsl, Tertiary shale and sandstone (San 

Lorenzo formation); Tsr, Tertiary mudstone and sandstone (Rices Mudstone member); Tv, Tertiary 

sandstone (Vaqueros sandstone).
E91

the San Andreas fault are believed to represent Tertiary marine 

sedimentary rocks. The 1- to 10-Ω-m-resistivity units in the 

uppermost 4 km of crust northeast of the San Andreas fault 

are believed to represent Mesozoic Great Valley sedimentary 

rocks (Irwin, 1990). The 1,000-Ω-m-resistivity material may 

represent Salinian-block granitoids of Mesozoic age, Coast 

Range ophiolite, and Mesozoic units of the Franciscan assemblage in some combination, irresolvable because of masking 

by the thick (4 km) conductive (1- to 10-Ω-m-resistivity) 

overburden.

CONCLUSION

Our resistivity modeling suggests that details of the local 

deep structure can be mapped by using magnetotelluric 

surveys to assist in understanding the dynamics of the San 

Andreas fault system. The interpreted structure from our two 

profiles will provide constraints on future geologic interpretations of the Loma Prieta region. Resistivities in these deep 

rocks are almost entirely related to fracturing and fluid content; 

thus, the deep geoelectrical structure adds significant new 

information toward understanding seismic-stress release in 

the region. The Loma Prieta MT model indicates (1) a large, 

high-resistivity body within the upper 2 km of crust between 

the San Andreas and Sargent faults at seismogenic depths that 

may have played an important role in seismic-stress buildup; 

and (2) a low-resistivity anomaly at midcrustal depths near 

the San Andreas fault that may have been caused by a buildup 

of overpressured fluids in the deeper parts of the fault-gouge 

zone, which reduced the shear stress required for sliding and, 

subsequently, triggered the earthquake. The presence of this 

low-resistivity anomaly was a key and required element of 

our model, even though the model was insensitive to the dip, 

shape, and distribution of the anomaly. High-resistivity units 

southwest of the San Andreas fault may represent Salinianblock granitoids of Mesozoic age. High-resistivity units below 

4-km depth northeast of the San Andreas fault could be related 

to the Franciscan complex. Low-resistivity rocks between the 

Zayante and Sargent faults in the uppermost 4 to 7 km of crust 

are interpreted from drill-hole data as mostly Tertiary marine 

sedimentary rocks exposed in tightly folded structures in the 

Loma Prieta region. Low-resistivity units above 4-km depth 

northeast of the San Andreas fault are believed to represent 

Mesozoic units of the Franciscan assemblage. The resistivity 

model agrees well with the three-dimensional seismic-velocity 

model, revealing distinctive velocity and resistivity variations 

in the Loma Prieta region associated with seismicity and faultrupture patterns. The resistivity model is also consistent with 

gravity and magnetic models of the Loma Prieta region. The 

San Andreas fault and adjoining subparallel Sargent and Zayante faults appear to juxtapose diverse deep geologic structures 

spanning the seismogenic zone in the Loma Prieta region, 

suggesting that local geologic structures may have played a 

role in triggering the earthquake and may control the overall 

seismicity on the San Andreas and Sargent faults. Another 

MT-sounding profile about 35 km southeast, traversing the 

northwest end of the Gabilan Range, shows a distinctly different structure. Low-resistivity units there appear to be more 

spatially uniform and limited to the uppermost 4 km of crust. 

This type of geoelectrical structure may indicate stable fault 

slip on the central section of the San Andreas fault.

REFERENCES CITED

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California: Missoula, Mont., Mountain Press, 244 p.

Brabb, E.E., and Hanna, W.F., 1981, Maps showing aeromagnetic 

anomalies, faults, earthquake epicenters, and igneous rocks 

in the southern San Francisco Bay region, California: U.S. 

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358, no. 6387, p. 574–576.

Clark, J.C., and Rietman, J.D., 1973, Oligocene stratigraphy, tectonics, and paleogeography southwest of the San Andreas 

fault, Santa Cruz Mountains and Gabilan Range, California 

Coast Ranges: U.S. Geological Survey Professional Paper 

783, 18 p.

Cole, W.F., Cotton, W.R., and Fowler, W.L., 1989, Engineering geology of the San Francisco Bay region, in Engineering geology 

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Prieta, California, earthquake and its aftershocks; geometry 

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Eberhart-Phillips, D.M., 1989, Investigations of crustal structure and 

active tectonic processes in the Coast Ranges, central California: Stanford, Calif., Stanford University, Ph.D. thesis, 216 p.

Eberhart-Phillips, D.M., Labson, V.F., Stanley, W.D., Michael, A.J., 

and Rodriguez, B.D., 1990, Preliminary velocity and resistivity models of the Loma Prieta earthquake region: Geophysical 

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no. 1, p. 53–68.

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61–80.

Kaufman, A.A., and Keller, G.V., 1981, The magnetotelluric sounding method (Methods in Geochemistry and Geophysics, no. 

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Mackie, R.L., Bennett, B.R., and Madden, T.R., 1988, Long-period 

magnetotelluric measurements near the central California 

coast; a land-locked view of the conductivity structure under 

the Pacific Ocean: Royal Astronomical Society Geophysical 

Journal, v. 95, no. 1, p. 181–194.

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Mazzella, A.T., 1976, Deep resistivity study across the San Andreas 

fault zone: Berkeley, University of California, Ph.D. thesis, 

145 p.

McLaughlin, R.J., Clark, J.C., and Brabb, E.E., 1988, Geologic map 

and structure sections of the Loma Prieta 71⁄2' quadrangle, 

Santa Clara and Santa Cruz Counties, California: U.S. Geological Survey Open-File Map 88–752, 31 p., scale 1:24,000, 

2 sheets.

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flow tuff from Yucca Mountain, Nevada: Journal of Geophysical Research, v. 97, no. B5, p. 6823–6841.

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Magnetotelluric evidence for crustal suture zones bounding 

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Plenum, 314 p.

Phillips, W.J., and Kuckes, A.F., 1983, Electrical conductivity structure of the San Andreas fault in central California: Journal of 

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Smith, J.T. and Booker, J.R., 1988, Magnetotelluric inversion for minimum structure: Geophysics, v. 53, no. 12, p. 1565–1576.

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B3, p. 3905–3922.

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Wiley & Sons, 414 p.

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Rymer, M.J., and Ellsworth, W.L., eds., The Coalinga, California, earthquake of May 2, 1983: U.S. Geological Survey 

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no. 1, p. 60–74.

Wannamaker, P.E., Stodt, J.A., and Rijo, Luis, 1987, PW2D; finite 

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of two-dimensional earth resistivity structure: Salt Lake City, 

University of Utah, Research Institute, 40 p.

APPENDIX 1. 

MAGNETOTELLURIC METHODOLOGY

The magnetotelluric (MT) method (Kaufman and Keller, 

1981; Vozoff, 1991) is an electromagnetic-frequencysounding method utilizing natural variations of the Earth’s 

magnetic and electric fields originating from geomagnetic 

micropulsations over the frequency range 0.0001–1 Hz and 

from worldwide lightning activity over the frequency range 

1–10,000 Hz. These natural electric and magnetic fields are 

recorded in two orthogonal, horizontal directions (and the 

vertical magnetic field is sometimes recorded as well) at 

the Earth’s surface with a computer-based data acquisition 

and processing system. The resulting time-series signals are 

transformed into the frequency domain and used in a crossspectral analysis to derive earth tensor resistivities and phases. 

Generally, a rotated coordinate system that corresponds to 

resistivity measured along electrical strike, called the transverse-electric direction (TE), and normal to strike, called the 

transverse-magnetic direction (TM), is used. Phase values for 

the complex earth impedances used to calculate resistivities 

are also computed and can be used in modeling. In two- and 

three-dimensional modeling, models are derived to fit both 

the TE and TM resistivities.

Measurement of earth resistivities can be an effective means 

of mapping geologic structures because of the wide variation 

in the electrical resistivity of rocks. The resistivity of geologic 

units largely depends on their fluid content, alteration, fracturing, carbon and other conductive-mineral contents, and 

temperature (Keller and Frischknecht, 1966). Formation fluids 

play a large role in reducing resistivities below those of the 

normally high resistivity rock matrix, especially when formation salinities and porosity are high. Resistivities can also be 

lowered by the presence of highly conductive clay minerals, 

graphitic carbon, and metallic mineralization. Volcanic rocks 

and fault zones commonly contain conductive authigenic 

minerals and gouge with resistivities 10 to 100 times lower 

than those of the surrounding rocks (Nelson and Anderson, 

1992). Increasing temperatures cause higher ionic mobility 

and mineral-activation energy, reducing rock resistivities 

significantly. The MT method has distinct advantages over 

other geophysical methods in the study of complex geologic 

environments because of the wide variation in rock resistivities 

and the inherent tensor nature of the technique, which makes 

it sensitive to vertical structures. Furthermore, the method 

is capable of establishing whether the subsurface structures 

being measured are effectively one, two, or three dimensional 

in terms of their response to electromagnetic fields.

Originating from the naturally occurring sources that the MT 

method uses, time-varying magnetic fields induce current flow 

in the conductive Earth. The Earth’s resistivity is determined 

from the relation of these time-varying magnetic and electric 

fields (electrical impedance). Because lower frequencies penetrate deeper, measurement of a spectrum of frequencies allows 

the creation of an electrical sounding, which is a measure 

of electrical resistivity versus frequency. In a plane-layered 

Earth, the MT sounding could be inverted to produce a plot of 

resistivity versus depth. However, the Earth is seldom plane 

layered, and the MT method responds to lateral variations in 

resistivity. For a two-dimensional Earth—that is, an Earth 
E93

that varies only with depth and in a direction perpendicular to 

strike—profiles of MT soundings may be inverted to produce 

electrical cross sections. Although much of the surrounding 

structure is finite and complex (see Jachens and Griscom, 

this chapter), the linearity of some geologic structures in the 

Loma Prieta region makes two-dimensional modeling useful 

for approximating the local geology. At present three-dimensional modeling is inadequately crude and prohibitively time 

consuming, and so we used two-dimensional modeling to 

construct the interpretative cross sections presented here.

APPENDIX 2. 

MAGNETOTELLURIC COVERAGE

Station spacing (fig. 1) ranged from 0.4 to 5 km. Frequencies sampled were from 0.001 to 300 Hz, using single-station 

recordings of both orthogonal horizontal components of the 

electric and magnetic fields. These data were measured with 

USGS-designed instrumentation and processed in real time 

with the computer that controls the MT system. The recorded 

time-series data were Fourier-analyzed to determine a twodimensional resistivity and phase tensor at each station, which 

were then freely rotated to maximum- and minimum-resistivity directions so that propagation modes for the signals were 

decoupled into transverse-electric and transverse-magnetic 

modes that are equivalent to measurement of the Earth’s 

resistivity with the electric field directed along and across 

geologic strike, respectively. This data-analysis method uses 

the crosspower method of Gamble and others (1979), which 

produces estimates of apparent resistivity that are unbiased by 

noise but are subject to large random errors when coherencies 

are less than 0.9. Therefore, the observed data were manually 

smoothed and digitized in an effort to reduce systematic and 

random errors and coherent noise. The observed data were 

also corrected for static shift from near-surface effects, such 

as thin localized conductors.

Plots of tensor resistivity and phase for the manually 

smoothed representations of the Loma Prieta data are shown 

in figures 2A and 2B, respectively, along with the model 

curves. Both resistivity and phase values are included in the 

modeling process. These data have been interpreted by using 

two-dimensional finite-element forward modeling, which 

resulted in an estimated variation in resistivity from 3 to 1,000 

Ω-m for a cross section extending from about 0.1- to more 

than 10-km depth.

If the data are obtained over structures that are three dimensional, the skew of the tensor will be significant, whereas for a 

truly two dimensional Earth the skew will be zero. Most of the 

data from the Loma Prieta survey had low skews, 0.1 to 0.3; 

but near stations 3 and 8 (fig. 1), the skews for some frequencies ranged from 0.3 to 1. Associated with the low skews was a 

northwestward principal-axis rotation of the resistivity tensors 

(fig. 7). Associated with the high skews at stations 3 and 8 is a 

departure from the typical northwestward rotation of the tensor 

principal direction to an azimuth of 000° (north). The north 

principal direction suggests a three-dimensional geoelectrical 

structure at these stations, indicating that the underlying resistivity structure may be limited in strike length. The anomalous 

principal tensor direction of station 15 was north by northwest, 

yet had low skews. Most of the data from the Gabilan Range 

survey had low skews (typically, 0.1–0.5) except station 2, 

which had powerline interference and whose skews ranged 

from 0.1 to 1. All of these stations exhibited a northwestward 

principal-axis rotation of the resistivity tensors (fig. 7).

APPENDIX 3. 

MAGNETOTELLURIC MODELS

The following discussion provides a detailed account of our 

modeling approach not presented in the main body of text, 

for those readers interested in our starting models, assumptions, boundary conditions, biases, and alternative models.

LOMA PRIETA PROFILE

Initially, using the combined 1989, 1990, and 1992 data 

sets, we inverted the data with the two-dimensional inversion 

algorithm of Smith and Booker (1991), using a simple layer 

(4 km thick, 30-Ω-m resistivity) over a half-space (1,000 

Ω-m) as the starting model. Their inversion method is called 

the rapid-relaxation inverse (RRI) because of similarities 

to classical relaxation solutions and because of its speed. 

Another advantage to the RRI method arises from the use 

of a minimum-structure criterion (Smith and Booker, 1988). 

The algorithm uses a selected number of iterations to achieve 

a fit to the observed data, and then applies another selected 

number of steps in which the model complexity is reduced 

while maintaining an overall goodness of fit. These subsequent 

steps substantially reduce nonessential model geometries in 

the final solution. The final inversion model (fig. 8) generally 

fitted the transverse-magnetic (TM) data better than the transverse-electric (TE) data (figs. 8B and 8C, respectively). The 

gross structure of our model came from fitting the TM data. 

Wannamaker and others (1987) showed that finite structure 

is better resolved with the TM mode and that the TE mode is 

highly insensitive to gross resistivity structure. We then used 

this minimum-resistivity-structure inversion as the starting 

model for our two-dimensional forward-modeling efforts to 

improve the computed fits to the observed TE data.

Although the inversion model gives an approximation to the 

minimum-resistivity structure, it does not adequately take into 

account the effect of the Pacific Ocean because of a limitation in the algorithm. We modeled the seawater of the Pacific 

Ocean by using a resistivity of 0.3 Ω-m (Stacey, 1977; Sheriff, 

1991). The seawater thickness was varied from 20-m depth in 

Monterey Bay (U.S. Geological Survey, 1980) to 4-km depth at 

a distance of 200 km from the California coast (U.S. GeologiDEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION
E94 EARTHQUAKE OCCURRENCE

cal Survey, 1981). We modeled the oceanic crust beneath the 

Pacific Ocean by using a resistivity of 1,000 Ω-m at about 4-km 

depth. The conductive seawater had the effect of increasing the 

resistivity values on the TM curve by at least 10 percent of a 

log cycle at 0.01 Hz for stations as close as 2.5 km and at 0.003 

Hz for stations as far away as 28.1 km and of decreasing the 

resistivity values on the TE curve by at least 10 percent at 0.01 

Hz for stations as close as 2.5 km. Our forward modeling, we 

believe, adequately accounts for these effects.

In the minimum-resistivity-structure model (fig. 8), we see 

a resistivity low at about 5-km depth between stations 5 and 

13 (fig. 1), a resistivity high beneath stations 3 and 16, and a 

broader resistivity low northeast of station 3. The model curves, 

overall, make a poor to fair fit to the observed data. At most 

stations, the TM curve was fitted satisfactorily, but the TE 

curve needed to be decreased at the low to midband frequencies. The only approach that seemed to achieve this effect was 

to put lower resistivities in the depth range 4–16 km near the 

San Andreas fault. Although the model was insensitive to this 

resistivity configuration (fig. 9), within that depth range the 

TE curves had to be decreased at stations 6, 9, and 10 through 

14 (figs. 10, 11). This model sensitivity is consistent with the 

resolution inherent in the MT method, which decreases with 

penetration. Thus, both horizontal and vertical boundaries at 

these depths were imprecisely resolvable.

We used 3 Ω-m as our best choice for the conductive material at these depths. We also tried 1 Ω-m, 10 Ω-m, and higher 

resistivities, but the computed fit was generally worse. The 

1-Ω-m resistivity tended to decrease the resistivities on the 

TE curve too much at low to midband frequencies, and the 

10-Ω-m resistivity tended to increase the resistivity values 

on the TE curve too much at midband frequencies. We tried a 

single-body conductor beneath stations 6 and 11 through 13 

(fig. 1), reducing its thickness to compensate for the decrease 

in resistance, but we lost the anisotropy that is needed to fit 

the splitting TE and TM curves at station 12. Increasing its 

thickness does not cause the computed curves to split, and only 

worsens the misfit to the TE curves at stations 6, 11, 15, and 

Figure 7—Index map of the Loma Prieta region, Calif., showing electrical-resistivity strike directions 

(double-headed arrows) at stations (numbered dots) where magnetotelluric soundings were made. Same 

symbols as in figure 1.
E95

Figure 8.—Loma Prieta magnetotelluric (MT) inversion model (see fig. 1 for location) with 

contours of resistivity (in ohm-meters). A, Two-dimensional MT inversion model. Numbers 

denote stations shown in figure 1. B, Resistivity data. C, Phase data. Same symbols as in 

figure 2.

16 by decreasing the resistivities too much. Modeling vertical 

3-Ω-m conductors closer to the surface expression of the San 

Andreas fault decreased the resistivities on the TE curves too 

much at midband frequencies at stations 15 and 16.

The resistivity high beneath stations 3 and 16 (fig. 1) is within 

2 km of the surface. Modeling this geoelectrical structure at 

greater depths decreased the resistivities on the TM curve at 

low to midband frequencies at stations 3 and 16. We noted, 

however, that we could move the northeast edge of the resistive interface southwest as much as 0.5 km without adversely 

misfitting the observed data. The 100-Ω-m-resistivity material 

northeast of station 3 could be increased in thickness to at least 

19-km depth; however, the observed data do not require such 

an increase in our models.

We manually perturbed the starting two-dimensional model 

more than 100 times. The best-fit model is shown in figure 

2. Fits to the observed data were generally within measurement errors in the data. Significant misfits included the TE 

data at station 8 (fig. 1) at low to midband frequencies and 

the TM data at station 11 at low frequencies. Wannamaker 

and others (1984) showed in three-dimensional MT modeling that approximating three-dimensional structure beneath 

a sounding with two-dimensional modeling is best achieved 

by fitting the TM curve even at the expense of fitting the 

TE curve. This result is what our model achieves at station 

8, where the data denoted an underlying three-dimensional 

structure (see app. 2). The misfit to the TM curve at low 

frequencies for station 11 most likely indicates a problem 

with the data at this station—that is, a data-acquisition error 

occurred at this station.

In summary, we found that we had to decrease the resistivities in the depth range 4–16 km near the San Andreas fault 

to bring the TE curve down at low to midband frequencies. 

The dip, shape, and precise distribution of this low-resistivity 

body were irresolvable, even though its presence was a major 

requirement of our model. We found that, in its absence, no 

other solution adequately fitted the observed TE data. The 

low-resistivity zones at deeper levels could conceivably have 

numerous variations and still be consistent with the observed 

data, although our model requires some degree of lateral 

inhomogeneity in resistivity within the uppermost 16 km of 

crust across the Loma Prieta epicentral region.

A

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION
E96 EARTHQUAKE OCCURRENCE

GABILAN RANGE PROFILE

Initially, we inverted the data with the two-dimensional 

inversion algorithm, using a simple layer (4 km thick, 

30-Ω-m resistivity) over a half-space (1,000 Ω-m) as the 

starting model. The final inversion model generally fitted the 

data. We then used this minimum-structure inversion as the 

starting model in our two-dimensional forward-modeling 

efforts to improve the computed fits to the observed data. 

We determined that we needed to add a low-resistivity body 

southwest of station 4 (fig. 1) in the uppermost 4 km of 

crust to better fit the structure required by both modes of the 

Figure 8.—Continued

B
E97

observed data, deepen the central low-resistivity anomaly 

beneath stations 5 and 6 to better fit the TE data, and add 

low-resistivity material in the uppermost 4 km of crust 

beneath station 1 to better fit the TM data. The dip, shape, 

and precise distribution of this low-resistivity material were 

Figure 8.—Continued

C

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION

partly irresolvable because of the wide spatial sampling 

of the stations and because of the thick near-surface lowresistivity material. We manually perturbed the starting 

two-dimensional model more than 100 times. The best-fit 

model is shown in figure 3.
E98 EARTHQUAKE OCCURRENCE

Figure 9.—Loma Prieta magnetotelluric (MT) profile (fig. 2; see fig. 1 for locations), with alternative 

variation of low-resistivity case of two-dimensional finite-element MT model. A, Two-dimensional 

finite-element MT model. Numbers denote stations shown in figure 1. B, Resistivity data. C, Phase 

data. Same symbols as in figure 2.

B

A
E99

C

Figure 9.—Continued

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION
E100 EARTHQUAKE OCCURRENCE

Figure 10.—Loma Prieta magnetotelluric (MT) profile (fig. 2; see fig. 1 for locations), with moderate depth extension of lowresistivity case of two-dimensional finite-element MT model. A, Two-dimensional finite-element MT model. Numbers denote 

stations shown in figure 1. B, Resistivity data. C, Phase data. Same symbols as in figure 2.

A
E101

Figure 10.—Continued

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION

B
E102 EARTHQUAKE OCCURRENCE

Figure 10.—Continued

C
E103

Figure 11.—Loma Prieta magnetotelluric (MT) profile (fig. 2; see fig. 1 for locations), with shallow depth extension of low-resistivity case 

of two-dimensional finite-element MT model. A, Two-dimensional finite-element MT model. Numbers denote stations shown in figure 1. 

B, Resistivity data. C, Phase data. Same symbols as in figure 2.

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION

A
E104 EARTHQUAKE OCCURRENCE

Figure 11.—Continued

B
E105

Figure 11.—Continued

DEEP GEOELECTRICAL STRUCTURE OF THE LOMA PRIETA REGION

C
E106 EARTHQUAKE OCCURRENCE
E107

CONTENTS

 

 Page

Abstract--------------------------------------------------------------------- E107

Introduction---------------------------------------------------------------- 107 

Geologic setting----------------------------------------------------------- 108 

Data acquisition and processing----------------------------------------- 108 

Data description----------------------------------------------------------- 110 

Forward modeling--------------------------------------------------------- 111 

Velocity models------------------------------------------------------------ 114 

 Loma Prieta transect------------------------------------------------ 115 

 Año Nuevo transect------------------------------------------------ 116 

Limitations of the velocity models-------------------------------------- 118 

 Absence of reversals----------------------------------------------- 118 

 Faint PiP reflections----------------------------------------------- 120 

Discussion------------------------------------------------------------------ 121 

Summary------------------------------------------------------------------- 122 

Acknowledgments-------------------------------------------------------- 122 

References cited----------------------------------------------------------- 124 

—————

ABSTRACT

Wide-angle seismic-reflection/refraction data obtained in 

the vicinity of the epicenter of the 1989 Loma Prieta earthquake provide evidence for a slab of oceanic crust extending 

from the deep-sea floor to the San Andreas fault. During 

marine seismic-reflection profiling of the north-central California margin in May 1990, two onshore recording arrays were 

deployed transverse to and straddling the San Andreas fault in 

the Santa Cruz Mountains and the Diablo Range. Wide-angle 

reflections recorded by these arrays define the top and bottom 

of oceanic crust as far eastward as the San Andreas fault. The 

wide-angle refraction data obtained by the arrays provide 

evidence that the oceanic crust, apparently connected to and 

inferred to represent Pacific plate, is overlain by a seawardtapering wedge of rocks with velocities of 6.0 to 6.6±0.1 km/s. 

We interpret this wedge as composed of slivers of San Simeon 

and Salinian (and, possibly, other) rocks. The wide-angle 

data are consistent with a landward dip on the oceanic crust 

of 8°–11°. The velocity model inferred from these data suggests that the hypocenter of the earthquake was located near 

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN

WEST OF THE SAN ANDREAS FAULT

FROM ONSHORE-OFFSHORE WIDE-ANGLE SEISMIC RECORDINGS

By Thomas M. Brocher, Michael J. Moses, and Stephen D. Lewis, U.S. Geological Survey

the base of the interpreted wedge of Salinian rocks beneath 

the continental margin but above the oceanic crust. On the 

basis of this geometry, we infer that the top of oceanic crust 

represents a detachment surface accommodating oceanward 

movement of North America. This oceanward movement 

mirrors the eastward-directed tectonic wedging observed 

beneath the Diablo Range and the western part of the Great 

Valley. An observed concave-upward geometry of the Moho 

in the vicinity of the San Andreas fault should strongly focus 

wide-angle reflections from the Moho that travel along the 

fault from earthquake sources on or near the fault.

INTRODUCTION

Knowledge of the structure of the midcrust to lower crust 

offshore central California is fundamental to our understanding 

of the tectonic processes associated with, and the earthquake 

hazards posed by, ongoing transpression of the California 

margin. Since early studies of the continental margin off 

California (Uchupi and Emery, 1963; Curray, 1966), workers 

have recognized that an Andean-type continental magmaticarc/trench system existed during the Mesozoic throughout the 

region of present central California. That system was disrupted 

by Cenozoic strike-slip faulting, which included a period of 

transtension (Miocene) and transpression (Pliocene and Quaternary) (for example, Dickinson, 1970; Ernst, 1971; Saleeby, 

1986; McCulloch, 1987). Major anticlinal-fold structures in 

the offshore Santa Maria Basin have been dated at Pliocene 

(3–5 m.y.), suggesting that compressional forces related to 

Pacific plate-motion changes began at that time (McCulloch, 

1987; Clark and others, 1991). Present transpression across 

the margin has been inferred from geodetic, focal-mechanism, 

stress-measurement, and other data (Zoback and others, 

1987). Although the ongoing transpressional tectonics favors 

thrust and strike-slip faults (for example, Namson and Davis, 

1988), the existence and geometry of these thrust faults in 

north-central California, and their relation to the San Andreas 

fault, are poorly known because most of the faults are buried 

(Wentworth and others, 1984). In the vicinity of Morro Bay, 

south of Monterey, seismic-reflection/refraction profiling 

E107
E108 EARTHQUAKE OCCURRENCE

by the Pacific Gas and Electric Co. (PG&E) and EDGE (a 

consortium of university investigators) suggests that the 

outer margin and shelf are underlain at midcrustal levels by 

oceanic crust, possibly of the Farallon plate, and that the top 

of this oceanic crust may act or have acted as a subhorizontal 

detachment surface (Ewing and Talwani, 1991).

The 1989 Loma Prieta earthquake served both as a reminder 

of the active transpression occurring along the San Andreas 

fault and as a catalyst for framework studies aimed at better 

understanding the crustal structure in the vicinity of the 

fault (U.S. Geological Survey staff, 1990). Large-amplitude 

reflections from the midcrust to the lower crust and the top 

of the upper mantle may have increased the ground motions 

generated by the earthquake at some distant sites (Somerville 

and Yoshimura, 1990) and provide additional motivation for 

the delineation of subhorizontal structures in the midcrust to 

lower crust. We report here the primary results of an onshoreoffshore seismic-reflection/refraction investigation of the 

crustal structure of the California continental margin in the 

Loma Prieta region. We present seismic evidence for oceanic 

crust underlying the margin and extending continuously from 

the present deep-sea floor as far eastward as the San Andreas 

fault. We describe evidence favoring the presence of a sliver 

of Salinia terrane overlying the oceanic crust and propose 

a speculative tectonic model based on the observed crustal 

structure.

GEOLOGIC SETTING

The seismic profiles presented here nearly coincide with the 

C2 transect of central California (fig. 1; Saleeby, 1986; Fuis 

and Mooney, 1990). We therefore refer the reader to these 

studies for a detailed summary of geologic relations and provide here only a brief review of the main geologic structures 

crossed by the seismic survey.

As described by Fuis and Mooney (1990), the central 

California margin consists of an inactive, early Tertiary 

accretionary prism that is overlain by Oligocene to Holocene sedimentary rocks (Uchupi and Emery, 1963; Silver 

and others, 1971; Mullins and Nagel, 1981; McCulloch, 

1987). On the basis of dredging results and geophysical data, 

parts of the margin have been interpreted to be underlain by 

Franciscan complex (Silver and others, 1971; Mullins and 

Nagel, 1981). The margin is believed to be floored by oceanic 

crust dated at 20 to 26 m.y., on the basis of an identification of magnetic anomalies on the Pacific plate (Atwater 

and Menard, 1970; Atwater, 1989; Fernandez and Hey, 

1991; Lonsdale, 1991). Near the coast, the inactive Tertiary 

accretionary prism is juxtaposed against the Salinian block, 

which consists of granitic and high-temperature-metamorphic rocks overlain by Upper Cretaceous, Paleogene, and 

Neogene sedimentary rocks. At the latitude of our study 

(fig. 1), the Salinian block is bounded on the west by the San 

Gregorio-Hosgri fault and on the east by the San Andreas 

fault (U.S. Geological Survey and California Division of 

Mines and Geology, 1966; Jennings and others, 1977; Ross, 

1978) and here includes mafic basement adjacent to the San 

Andreas fault.

Franciscan rocks lie between the San Andreas fault and a 

subsurface boundary beneath the San Joaquin Valley and in 

the Diablo Range (U.S. Geological Survey and California 

Division of Mines and Geology, 1966; Jennings and others, 

1977; Ross, 1978). According to Wentworth and Zoback 

(1990, p. 44):

The Franciscan assemblage, * * * of Late Jurassic to Late 

Cretaceous age, is an accretionary prism that consists principally of oceanic sedimentary and lesser volcanic rocks. Some 

tectonostratigraphic units in the Franciscan consist of wellbedded, albeit variously folded, rocks, whereas others are a 

melange of various kinds and sizes of blocks in a matrix of 

sheared mudstone (Blake and others, 1984). Most of the rocks 

have varyingly undergone high-pressure, low-temperature 

metamorphism to blueschist grade. The structural base of 

the Franciscan is nowhere exposed at the surface.

In the study area (fig. 1), two quite different types of 

Franciscan rocks are involved: in the Santa Cruz Mountains, 

graywacke, melange, and much greenstone, with a shallow 

high-velocity “basement” (pumpellyite to blueschist grade); 

and in the Diablo Range, metasandstone and melange (all 

blueschist, locally containing jadeite), with a thickness based 

on seismic-refraction velocities of 12 to 15 km.

Fuis and Mooney’s (1990) compilation of existing seismic 

refraction and reflection studies along the C2 transect is 

modified in figure 2. Seismic-refraction profiling by Shor 

and others (1971) found oceanic crust of the Pacific plate 

just to the west of the base of the continental slope in deep 

water. Onshore seismic-refraction profiles, which were 

acquired along the strike of the San Andreas fault to the 

south of Loma Prieta, indicate that the Moho lies at 24- to 

26-km depth just west of the San Andreas fault and that the 

crust may be about 30 km thick in the Diablo Range (Walter 

and Mooney, 1982; Blümling and Prodehl, 1983; Blümling 

and others, 1985).

DATA ACQUISITION AND PROCESSING

In May 1990, the U.S. Geological Survey (USGS) conducted 

a marine multichannel-seismic-reflection investigation of 

the central California margin (Lewis, 1990). The 10-element 

airgun array used by the research vessel S.P. Lee for the 

seismic-reflection profiling, totaling 40 L, was simultaneously recorded by two temporary onshore arrays of recorders 

deployed transverse to and straddling the San Andreas fault 

in the Santa Cruz Mountains and the Diablo Range (fig. 1B; 

Brocher and others, 1992). The Loma Prieta array consisted 

of 13 stations stretching from the coast landward 94 km to the 

San Joaquin Valley. A shorter array, located 25 km northeast 

of Loma Prieta, consisted of six stations stretching landward 

about 70 km from Año Nuevo to Cedar Mountain in the Diablo 
E109

Study area

Morro Bay

Continental 

 Shelf

 and slope

California

0 100 200 

KILOMETERS

Line 32

Line 38

AN2

AN5

LP10

LP4

LP7

HAYWARD FAULT

SAN ANDREAS FAULT

CALAVERAS FAULT

SAN GREGORIO 

FAULT

40

122 00' o

38 00' o

 San

Francisco

 Bay

USGS 5-day Recorder

LP12

LP1

Año Nuevo

 transect

Loma

Prieta

transect

C2 Transect

Base of

 continental

 slope

Monterey

0 10 20 30

KILOMETERS

PACIFIC 

 OCEAN

121 30' o

122 30' o 123 00' o

37 30' o

37 00' o

36 30' o

Fig. 1B

BASIX

Figure 1.—Loma Prieta region, Calif. A, California, showing locations of study area and base of 

the Continental Shelf and slope (dashed line). B, Central California, showing locations of marine 

seismic-reflection lines 32 and 38, temporary recording stations (triangles), selected faults, 1991 

BASIX lines (Holbrook and others, 1996), and C2 transect of Saleeby (1986). Star, epicenter of 

October 17, 1989, M=7.1 Loma Prieta, Calif., earthquake.

A

B

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E110 EARTHQUAKE OCCURRENCE

Range. These temporary arrays were deployed to record two 

parallel 61-km-long seismic-reflection lines (32, 38, fig. 1B) 

from source-receiver offsets ranging from 11 to 165 km. The 

goals of the wide-angle recording were to constrain the midcrustal to lower-crustal structure in the Loma Prieta region 

and to look for variations in this structure along the strike of 

the San Andreas fault.

The wide-angle data were acquired with analog threecomponent, 5-day seismic recorders, as described by Criley 

and Eaton (1978). The 5-day recorders were deployed along 

access roads and were turned on before acquiring the marine 

seismic-reflection lines (Brocher and others, 1992). At four 

stations (AN2, LP1, LP9, LP11, fig. 1B), seismic signals from 

a vertical-component seismometer were telemetered to an 

adjacent 5-day recorder for data acquisition.

The digitization of the 5-day analog tapes, the conversion 

of the digitized data into SEG–Y format, and the subsequent 

processing of the wide-angle record sections were described 

by Brocher and others (1992). Record sections reduced by a 

velocity of 6 km/s were generated by this procedure.

DATA DESCRIPTION

Marine seismic-reflection lines 32 and 38 (fig. 1B) image 

similar structural elements on the central California Continental Shelf and slope (fig. 3). Seaward of Santa Cruz and the 

Outer Santa Cruz Basin high on the Continental Shelf, these 

lines image a thick slope basin locally containing more than 

2,000 m of sedimentary deposits, probably of Oligocene to 

Holocene age. The slope deposits thin over, and are deformed 

by growth of, an unnamed anticline at the base of the continental slope. Sedimentary deposits of the Monterey deep-sea fan 

thin against, onlap, and appear to be uplifted by this unnamed 

anticline. Moderately dipping reflections from within the core 

of the anticline probably represent splay thrust faults that dip 

landward. Low-frequency and diffractive reflections beneath 

the Neogene sedimentary deposits, which we infer to represent 

the top of subducted oceanic crust of the Pacific plate, can be 

traced about 20 km landward of the base of the slope.

All of the wide-angle records obtained along the Loma 

Prieta and Año Nuevo transects were presented by Brocher 

and others (1992). These records indicate five crustal and 

upper-mantle arrivals, as shown in figure 4. The first of these 

arrivals represent refractions from the upper crust (Pg), which 

differ noticeably on records from the two transects. On the 

Loma Prieta transect, Pg arrivals have apparent velocities of 

4 to 5 km/s and grade smoothly with arrivals with velocities 

closer to 6 km/s at greater ranges. On the Año Nuevo transect, Pg arrivals have an apparent velocity close to 6 km/s 

and are abruptly truncated near model coordinate 30 km. 

Modeling suggests that a group of secondary arrivals at the 

Figure 2.—Compilation of seismic refraction and reflection studies along C2 transect (modified from Fuis and Mooney, 1990). 

The San Simeon terrane west of the San Gregorio-Hosgri fault system is inferred to be composed of the Cretaceous Franciscan 

assemblage. The Salinian block in the Santa Cruz Mountains is composed mostly of upper Cretaceous granitic, high-temperaturemetamorphic, and Tertiary sedimentary rocks. The Franciscan assemblage in the Santa Clara Valley is composed of the Permanente 

terrane, which is overlain by Tertiary sedimentary rocks and outliers of Coast Range ophiolite and Great Valley sequence. Shaded 

areas indicate uncertainties in locations of midcrustal and Moho discontinuities. No vertical exaggeration.

Inactive early Tertiary accretionary wedge:

 Late Tertiary sedimentary rocks Franciscan 

assemblage?

Early Tertiary

 sedimentary rocks

Late Tertiary

oceanic crust

 Moho

(estimated zone of

possible locations)

6 . 3 5 - 6 . 7 

Franciscan assemblage

SAN ANDREAS

 FAULT CONTINENTAL SLOPE

 SAN GREGORIO-

 HOSGRI FAULT

 SYSTEM

 CALAVERAS

FAULT SYSTEM

Coast- line

SANTA CRUZ MOUNTAINS SANTA CLARA VALLEY

DIABLO

RANGE

Franciscan 

assemblage

5.45

6.3-6.4?

6.0

6 . 5 5

7.6

8.2

5.8

6.7

8 . 0

6 . 8

6 . 8

6 . 4

6 . 3 5

6 . 1 5

6 . 0

3.6/4.6

8 . 0 7.5

6.3

5.7

5 . 2 5 3.34

2.7

 Moho

(estimated zone of

possible locations)

4.5 4.5 5.0

3.6

4.5

6.8

5.4?

5.3?

SW NE

Salinian block

0 20 40

0

20

40

DEPTH, IN KILOMETERS

San Simeon terrane 

10

30

10 30

KILOMETERS
E111

farthest source-receiver ranges represent refractions from a 

slab with velocities of 6.6 to 7.1 km/s. In the nomenclature 

of seismic-refraction profiling of the oceanic crust, these 

arrivals represent P3 refractions from the gabbroic layer, 

although to our knowledge this nomenclature has not previously been applied to continental margins. A third group 

of arrivals represents wide-angle reflections (PiP) from the 

midcrust. On the Loma Prieta transect, these arrivals are fairly 

weak and generally observed only at small source-receiver 

offsets and on the highest quality profiles; these wide-angle 

reflections have much larger amplitudes on the Año Nuevo 

transect. A fourth group of arrivals represents large-amplitude reflections from the top of the upper mantle, the Moho, 

and is designated as PmP. PmP arrivals are curved strongly 

downward on the reduced record sections, indicating a landward-dipping Moho or a seaward-thinning crust (fig. 4). Pn

arrivals, which represent refractions from the upper mantle, 

are generally faint and have complex traveltime curves due 

to the complex structure.

The low signal-to-noise ratio of the recorded data and the 

reverberant-airgun-array source result in a greater uncertainty 

of the traveltime observations than is normally reported 

for wide-angle seismic experiments. Here, we believe that 

absolute traveltime uncertainties, though widely varying, are 

generally as great as 0.1 to 0.2 s, owing to the reverberant 

waveform and low signal levels. Relative traveltime uncertainty, however, is only 0.01 to 0.05 s. Thus, inferred phase 

velocities are believed to be more accurate than inferred 

layer thicknesses.

FORWARD MODELING

We forward-modeled observed P-wave traveltimes and 

amplitudes by using the ray theory of �ervený and others 

(1977) and the computer algorithms of McMechan and 

Mooney (1980) and Luetgert (1988) for two-dimensional 

velocity-depth structures. The velocity model consists of 

layers with linear velocity gradients and velocity discontinuities of varying magnitude between layers. The traveltime 

modeling proceeded by first fitting direct arrivals and then 

the refractions recorded nearest each receiver. These arrivals 

Santa

Cruz

high

Top of subducted

oceanic crust

Unnamed

anticline

Monterey Fan

B

Monterey 

 Fan

A

Unnamed

anticline Slope basin

Top of subducted

oceanic crust

Santa

Cruz

high

Outer

Santa 

Cruz

Basin

0 20 40

0

4

8

0

4

8

TWO-WAY TRAVEL TIME, IN SECOND

S

SW NE

DISTANCE, IN KILOMETERS

10 30 50 60

Slope basin

Figure 3.—Line-drawing interpretations of seismic-reflection lines 32 (A) and 38 (B) (see fig. 1B for locations), showing major Neogene 

basins and other structural elements revealed by these profiles. Seismic data have been migrated; for 2-km/s velocity appropriate for 

Neogene sedimentary deposits (shaded), sections are plotted with a vertical exaggeration of about 1.78.

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E112 EARTHQUAKE OCCURRENCE

Figure 4.—Synthetic (A) and observed (B) record sections and ray diagram (C) for receiver at station LP4 on the Loma Prieta transect 

(see fig. 1B for location). Receiver was located at model coordinate 92 km. Synthetic seismograms in figure 4A were calculated from 

velocity model shown in figure 5B, using ray-theoretical method of McMechan and Mooney (1980); synthetic seismograms are plotted 

in true-relative-amplitude format, using a correction for geometric spreading of (Range)0.5. Record section shown in figure 4B is plotted 

showing true relative amplitudes, using a correction for geometric spreading of (Range)0.7. Velocity rosette on right indicates apparent 

phase velocities of arrivals. Calculated traveltimes along the Loma Prieta transect are superposed on record section and are queried where 

predicted branch does not correspond to clear arrivals in data. Arrivals include refractions from upper crust (Pg) and upper-mantle (Pn), 

and midcrustal (PiP) and upper mantle (PmP) reflections, all keyed to velocity model through layer numbers of model shown in figure 

4C. Pg arrivals are refractions from layer 6, PmP is a postcritical reflection from top of layer 9, PiP is a postcritical reflection from top 

of layer 8, P3 is a refraction from layer 8, and Pn is a refraction from layer 9. Critical points are indicated by dots along tops of layers 8 

and 9. Velocity model in figure 4C is plotted without vertical exaggeration.

0 20 40 60

DISTANCE, IN KILOMETERS

80

C

B

10

20

0

DEPTH, IN KILOMETERS

T-X/6, IN SECONDS

2

4

San Gregorio

fault

A

T-X/6, IN SECONDS

2

4

San Andreas

 fault

SW NE
E113

are used to determine the velocities and thicknesses of the 

uppermost layers of the model near each receiver, which, 

when interpolated between receivers, yield a two-dimensional 

velocity model for the entire receiver array. Subsequent arrival 

times are then fitted, always working downward through the 

model, to ensure that all previous, earlier arrivals are matched 

by the model. Velocity discontinuities between layers are 

introduced to replicate observed reflection arrivals. Because 

the velocity models are two dimensional, we assumed that 

the seismic-reflection lines and recording geometry could 

be approximated by a straight line. From figure 1B, this 

assumption is reasonable. Final velocity models obtained by 

this analysis for seismic-reflection lines 32 and 38 are shown 

in figure 5.

Observed Pg traveltimes were generally matched to within 

0.10 s by forward modeling. We estimate from trial-and-error 

perturbations of model parameters that P-wave velocities 

are resolved to within ±0.1 km/s by this forward-modeling 

procedure, particularly in the center of the velocity model, 

and that the depths of interfaces are resolved to within 5 to 10 

percent. The horizontal resolution varied along the model. On 

the seaward end, the shot spacing was 50 m, and the shortest 

Figure 5.—Velocity models inferred from forward modeling of P-wave traveltimes and amplitudes for 

the Año Nuevo transect along seismic-reflection line 32 and Año Nuevo (AN) deployment (A), and 

for the Loma Prieta transect along seismic-reflection line 38 and Loma Prieta (LP) deployment (B). 

Isovelocity contours between 4 and 8.2 km/s are plotted every 0.2 km/s. Dots, temporary recording 

stations (see fig. 1B for locations); dashed lines, layer boundaries. Layers: 1, ocean water, with a velocity 

of 1.5 km/s; 2, Neogene sedimentary deposits, with a velocity of 2 km/s; 3, Neogene sedimentary 

deposits, with a velocity of 3 to 4 km/s. Note low-velocity zone (LVZ) between layers 5 and 6 near 

model coordinates 40 km on Año Nuevo transect. Vertical exaggeration, 2:1.

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E114 EARTHQUAKE OCCURRENCE

wavelengths of observed P waves were about 400 m, suggesting a horizontal resolution of less than 1 to 2 km for the offshore 

region. Onshore, the receiver spacing averaged 8 km, suggesting that the horizontal resolution in upper-crustal structure 

onshore is closer to 8 to 10 km. Ray diagrams, showing P-wave 

raypaths through the model, indicate that the resolution of the 

procedure varies along the model, depending on the density of 

raypaths. In particular, the landward end of the model lacks ray 

coverage, owing to the absence of sources there, thus degrading 

the resolution of our forward-modeling procedure. For a more 

detailed description of model parameterization and error estimates in this type of forward modeling, the reader is referred 

to Hill and others (1985) and Fuis and others (1991). Recent 

comparisons between velocity models obtained from forward 

modeling and inverse methods of traveltime analysis suggest 

that the first-order features obtained from forward modeling 

are robust (Brocher and others, 1991). Synthetic seismograms 

were calculated from the velocity models by using asymptoticray theory (McMechan and Mooney, 1980), as implemented 

at the USGS (Luetgert, 1988).

VELOCITY MODELS

Velocity models were constructed from forward modeling of 

traveltimes and amplitudes for both the Loma Prieta and Año 

Nuevo transects, which are defined by seismic-reflection lines 

38 and 32, respectively (fig. 1B). The initial velocity models 

used for forward modeling were derived from several data 

sets, including multichannel-seismic-reflection lines 32 and 

38 (fig. 3), earthquake tomography of the shallow crust in the 

vicinity of the San Andreas and Calaveras faults, and onshore 

seismic-refraction studies (fig. 2).

Seismic-reflection lines 32 and 38 provide important constraints on the sea-floor bathymetry and geometry of Neogene 

basins along the line of sources. Sea-floor topography and 

Neogene sedimentary-basin thicknesses in terms of two-way 

traveltime were obtained from a time-migrated version of 

seismic-reflection lines 32 and 38 (fig. 3); to convert these 

estimates to depths, we assumed a seawater velocity of 1.5 

km/s and a velocity in the Neogene sedimentary deposits of 

2 km/s. The seismic-reflection lines allowed us to determine 

the geometry of the Outer Santa Cruz Basin, the Santa Cruz 

high near the seaward edge of the Continental Shelf southwest 

of the Outer Santa Cruz Basin, a slope basin lying seaward 

of the Santa Cruz high, the unnamed anticline, and, farthest 

seaward, sedimentary deposits of the Monterey deep-sea fan 

(fig. 3). The unnamed anticline on seismic-reflection lines 32 

and 38 is a major structural feature, displaying nearly 2 km of 

relief on the base of the Neogene sedimentary deposits. Mafic 

volcanic rocks dredged from the Santa Cruz high suggest that 

it may be cored either by Franciscan basement or Tertiary 

accretionary rocks (Silver and others, 1971; Mullins and 

Nagel, 1981). Diffractive reflections from the top of oceanic 

crust can be traced with confidence about 20 km landward of 

the relic trench on these seismic-reflection lines (fig. 3). The 

presence of reflections from the oceanic Moho in these data, 

however, is equivocal.

We forward-modeled vertical-incidence two-way traveltimes predicted by our model of the sea-floor bathymetry and 

Neogene sedimentary basin thickness for model coordinates 0 

to 60 km. The model was adjusted until the calculated two-way 

traveltimes matched the observed two-way traveltimes on the 

seismic-reflection profiles to within 0.05 s. Similar modeling 

of the reflection from the top of oceanic crust was performed 

from model coordinates 0 to 20 km (where this reflection is 

observed), and the model was adjusted until the calculated twoway traveltimes matched the observed two-way traveltimes on 

the seismic-reflection profiles to within 0.075 s.

The initial model for the shallow (0–15-km depth) crustal 

structure onshore (model coordinates 70–165 km) was taken 

from a tomographic study of the San Andreas and Calaveras 

faults using microearthquakes (Michael, 1988; Eberhart-Phillips and others, 1990a, b). This structure displays pronounced 

velocity contrasts across the San Andreas and Calaveras faults, 

as well as across low-velocity zones associated with these 

faults. Michael’s (1988) and Eberhart-Phillips and others’ 

(1990a, b) structures had to be significantly modified during 

our forward modeling to improve the fit to the observed traveltimes. Generally speaking, the modifications reduced the 

complexity of the velocity structure and minimized the size 

of lower-velocity basinlike structures. These modifications 

included the addition of a layer of surficial sedimentary deposits, as much as several hundred meters thick, with a velocity 

of 2.0 km/s. Estimates of crustal thickness east of the San 

Andreas fault and in the Diablo Range are those of Walter and 

Mooney (1982), Blümling and Prodehl (1983), and Blümling 

and others (1985).

In the absence of offshore receivers, we made several 

assumptions in constructing the initial velocity model beneath 

the continental margin. First, we projected the 4-, 5-, and 

6-km/s contours of Eberhart-Phillips and others (1990a, b) 

across the Continental Shelf and slope by assuming that these 

contours are subparallel to the base of the Neogene sedimentary deposits. Second, on the basis of our interpretation 

of the seismic-reflection data and in agreement with models 

for the C2 transect (Fuis and Mooney, 1990), we assumed 

that oceanic crust underlies the entire Continental Shelf and 

slope. Given our lack of control on the structure to the east 

of the San Andreas fault, and for convenience during forward 

modeling, we assumed that the oceanic crust is continuous 

across the fault. Third, following Shor and others (1971), we 

assumed that the oceanic crust is 6 km thick and that it can 

be adequately modeled by using the same structure as that 

determined by Tréhu (1991) for subducted oceanic crust in the 

vicinity of Morro Bay. Tréhu divided the oceanic crust into two 

layers with different thicknesses and velocity gradients. Thus, 

oceanic layer 2, consisting of extrusive volcanic rocks and 

sheeted dikes, is modeled as 1 km thick, with a linear-velocity gradient of 5.5 to 6.4 km/s; and oceanic layer 3, consisting 
E115

of gabbro, is modeled as 5 km thick, with a linear-velocity 

gradient of 6.6 to 7.1 km/s (fig. 5). Fourth, we assumed that 

the upper-mantle velocity is 8 km/s, in agreement with the Pn 

velocity determined by using Calnet data (Oppenheimer and 

Eaton, 1984) and with the upper-mantle velocity reported by 

Walter and Mooney (1982).

Velocity models inferred for the Loma Prieta and Año Nuevo 

transects differ significantly within the upper crust, although 

the overall structural elements are similar (fig. 5). The three 

major structural features resolved by the data include (1) a 

slab of oceanic crust with velocities of 5.5 to 7.1 km/s that 

dips landward and can be traced as far eastward as the San 

Andreas fault and possibly farther, (2) a midcrust with velocities of 6.0 to 6.5 km/s, and (3) lower average crustal velocities 

east than west of the San Andreas fault. The major difference 

in the velocity models for the two transects is the presence of 

upper-crustal velocities higher than 6.5 km/s west of the San 

Gregorio fault on the Año Nuevo transect but which are absent 

on the Loma Prieta transect to the south.

Before providing a detailed description of the velocity model 

obtained from each recorder array, we briefly summarize the 

evidence for the three major features of these models. The 

first feature, an east-dipping slab of oceanic crust, is primarily constrained by wide-angle reflections (PiP) observed from 

both the top and bottom of a 6.6- to 7.1-km/s layer. In addition, 

at greater source-receiver offsets, large-amplitude secondary 

arrivals are modeled as refractions from oceanic layer 3 that 

arrive earlier than refractions from within the midcrust with 

velocities of 6.0 to 6.5 km/s. The second feature, a low vertical-velocity gradient in the midcrust west of the San Andreas 

fault, is needed to match the large source-receiver ranges to 

which Pg arrivals are observed. If larger velocity gradients are 

used, the ranges to which these arrivals are observed cannot be 

matched, owing to the more rapid turning of rays with depth. 

The third feature, lower average crustal velocities east of the 

San Andreas fault, is consistent with velocity models obtained 

by inverting microearthquake traveltimes (Michael, 1988; 

Eberhart-Phillips and others, 1990a, b). These inversion results 

along the Loma Prieta transect indicate that velocities within 

the Franciscan Complex to the east of the San Andreas fault are 

lower than in the Salinian block to the west, as also determined 

by seismic-refraction profiling (Walter and Mooney, 1982). 

In the next section, we review in greater detail the constraints 

on velocity models for both the Loma Prieta and Año Nuevo 

transects.

LOMA PRIETA TRANSECT

The wide-angle seismic data recorded southwest of the San 

Andreas fault (stas. LP1–LP5, fig. 1B), between Loma Prieta 

and the coast, provide the best available information on the 

geometry of the oceanic crust and on upper-crustal velocities southwest of the San Andreas fault. Because these five 

stations provided nearly identical record sections, spanning 

a distance of only 28 km, we describe the modeling of these 

profiles as an ensemble rather than individually. An example 

of our modeling of these profiles is shown in figure 4. First 

arrivals observed from airgun shots nearest the receiver at 

stations LP1 to LP3 are explained by refractions within layer 

5 (fig. 4C). At greater source-receiver offsets, large-amplitude 

Pg arrivals are modeled by using raypaths that sample layer 6 

to about 10-km depth. First arrivals observed at greater ranges 

are more than 1 s too early to be explainable as refractions 

from layer 6 but are well matched by P3 and Pn refractions 

from layers 8 and 9. Undulations in observed traveltimes on 

the southwest ends of the profiles between model coordinates 

0 and 30 km are modeled by a pronounced structural relief of 

velocity contours across the unnamed anticline. The unusual 

moveout of reflected PmP arrivals on these profiles requires 

a seaward-thinning crust because a flat or substantially more 

gently dipping Moho beneath the continental margin does not 

predict the observed strong curvature of PmP arrivals (fig. 4). 

In turn, seaward thinning of the crust further constrains velocity gradients within the midcrust; deeply diving rays produced 

by the large velocity gradients in layer 6 would intersect the 

shoaling Moho, thereby truncating Pg arrivals prematurely. 

PmP arrivals and a faint wide-angle PiP reflection from the 

midcrust observed as far landward as the coastline (fig. 4B) are 

consistent with the 5-km modeled thickness of layer 8.

The records from stations LP6 to LP8 (fig. 1B), located 

between the San Andreas and Calaveras faults, are consistent 

with the presence of a small low-velocity body between these 

faults at model coordinates 105 to 120 km (fig. 6). First-arrival 

times require a large time delay modeled by replacing velocities of 4 to 5 km/s in layers 4 and 5 with velocities of 3 to 4 

km/s: no arrivals that turn within this low-velocity region are 

observed. Observed Pg first arrivals with apparent velocities of 

4.4 km/s are closely matched by refractions from layer 6. Largeamplitude secondary arrivals from model coordinates 0 to 30 

km with phase velocities higher than 6 km/s are modeled as P3

refractions from layer 8. The predicted amplitudes of reflected 

PiP arrivals are much weaker than the predicted amplitudes 

of PmP arrivals. The Pn arrivals predicted by refractions from 

layer 9 are very faint and not clearly observed.

Arrivals recorded at stations LP10, LP12, and LP13 (fig. 

1B), located east of the Calaveras fault, provide only limited 

constraints on the crustal structure near the San Andreas fault. 

The data recorded at stations LP9 and LP11 were of significantly lower quality than at the others (Brocher and others, 

1992) and so were not used to constrain the velocity model. 

Large-amplitude Pg arrivals with apparent velocities of 4 to 

6 km/s are modeled by refractions from layer 6 (fig. 7C). The 

calculated Pg raypaths (1) match the maximum source-receiver 

range to which Pg arrivals are observed (approx. 130 km), 

(2) sample the crust down to about 15-km depth beneath the 

coastline and San Gregorio fault, and (3) support the existence 

of a weak velocity gradient within the midcrust to lower crust. 

Weak Pn arrivals observed on the record from station LP12 

can be modeled by refractions that sample the Moho to the 

east of the San Andreas fault (fig. 7). Strong secondary arrivals 

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E116 EARTHQUAKE OCCURRENCE

between model coordinates 40 and 60 km are modeled as wideangle reflections and refractions from layer 8. We interpret 

the large velocity gradient in layer 8 as evidence that a slab 

of oceanic crust extends as far eastward as the San Andreas 

fault, although alternative explanations, such as magmatic 

underplating of the North American Continent, could also 

produce the modeled velocity gradient.

AÑO NUEVO TRANSECT

Recordings of seismic-reflection line 32 at stations AN1 to 

AN5 (fig. 1B) were used to construct a velocity model along the 

Año Nuevo transect. The data recorded at Station AN6 were of 

significantly lower quality than at the others (Brocher and others, 

1992) and so were not used to constrain the velocity model.

Figure 6.—Synthetic (A) and observed (B) record sections and ray diagram (C) for receiver at station LP6 on the Loma Prieta transect 

(see fig. 1B for locations). Receiver was located at model coordinate 108 km. Same format as in figure 4. Note large time separation 

between Pn, P3, and Pg branches at largest source-receiver offsets.

0 20 40 60

DISTANCE, IN KILOMETERS

80

�

�

10

20

0

DEPTH, IN KILOMETERS T-X/6, IN SECONDS

2

4

San Gregorio 

fault

�

T-X/6, IN SECONDS

2

4

San Andreas 

 fault

Fig. 6

100

SW NE
E117

Stations AN1 to AN3 (fig. 1B) all recorded wide-angle data 

southwest of the San Andreas fault in the Salinian block. The 

following discussion focuses on the high-quality data from 

station AN1 plotted in figure 8. Records from stations AN1 

to AN3 require an upper-crustal velocity structure noticeably 

different from that along the Loma Prieta transect. Pg arrivals 

along the Año Nuevo profile show evidence of a rapid decay 

in amplitude, indicating a conspicuous low-velocity zone (fig. 

8). These Pg arrivals are modeled as refractions within layer 

5 with a velocity of 5.3 to 6.6 km/s and a thickness of about 

7 km. We model the Pg arrivals as cut off at greater ranges by 

a low-velocity zone in layer 6 beginning at 7-km depth. At this 

low-velocity zone, velocities of 6.6 km/s overlie a region with 

velocities of 6.3 to 6.5 km/s between model coordinates 40 to 

0 40 80 120

DISTANCE, IN KILOMETERS

160

C

B

10

20

30

0

DEPTH, IN KILOMETERS T-X/6, IN SECONDS

2

4

Moho

PiP

San Gregorio 

fault

San Andreas

fault

Calaveras

fault

A

T-X/6, IN SECONDS

2

4

Figure 7.—Synthetic (A) and observed (B) record sections and ray diagram (C) for receiver at station LP12 on the Loma Prieta transect 

(see fig. 1B for locations). Receiver was located at model coordinate 156 km. Same format as in figure 4. Pg arrivals are modeled by 

refractions from layer 6 that turn upward in vicinity of the San Gregorio fault. Large-amplitude secondary arrivals are modeled by 

reflections and refractions from layers 7 and 8 in region between the San Gregorio and San Andreas faults. A weak Pn arrival is predicted 

by velocity model.

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E118 EARTHQUAKE OCCURRENCE

60 km. This narrow low-velocity zone is bounded to the east 

by the San Gregorio fault (fig. 5). At greater source-receiver 

ranges, large-amplitude first arrivals are modeled as wideangle reflections from the top of layer 8 and define this layer 

as extending eastward of the San Gregorio fault (fig. 8). It is 

difficult to trace these wide-angle reflections to smaller offsets 

at precritical angles of incidence. First arrivals at increasing 

ranges are modeled as P3 and Pn refractions from layers 8 and 

9. PmP arrivals are observed only at wide angles and cannot 

easily be traced to smaller offsets; the synthetic seismograms 

indicate that both PiP and PmP reflections will have small 

amplitudes at small offsets.

The wide-angle data recorded at stations AN4 and AN5 

(fig. 1B), both located in the Diablo Range to the northeast 

of the Calaveras fault, constrain structures in the vicinity of 

the San Andreas fault. The data recorded at these stations, 

separated by only 10 km, are nearly identical, and so we focus 

our discussion on station AN4 (fig. 9). The large-amplitude 

first arrivals on the profile from station AN4 were modeled 

as Pg refractions from layer 6 and, as on the Loma Prieta 

transect, provide evidence for a weak velocity gradient within 

the midcrust. Faint but clearly observed Pn arrivals (fig. 9) 

are consistent with the inferred Moho geometry and assumed 

upper-mantle velocity. A spatially limited set of secondary 

arrivals observed at about model coordinate 40 km are explainable by P3 refractions from layer 8. Faint PiP reflections are 

well matched by reflections from the top of layer 8 within 

10 km of the San Andreas fault. The strongly dipping, largeamplitude reflected arrivals at the nearest source-receiver 

ranges (model coordinates 40–60 km, fig. 9) are modeled as 

PmP reflections from the top of layer 9. These PmP reflections 

sample the Moho in the immediate vicinity of the San Andreas 

fault and, in combination with midcrustal PiP reflections, are 

consistent with the presence of oceanic crust underneath the 

entire Continental Shelf and slope.

LIMITATIONS OF 

THE VELOCITY MODELS

Uncertainties in the velocity models for the Loma Prieta 

and Año Nuevo transects are introduced primarily by (1) the 

absence of reversed raypaths caused by the absence of receivers on the sea floor and sources on land, and (2) the weakness 

of PiP reflections from the top of the 6.6- to 7.1-km/s layer. 

In the next subsections, we discuss the implications of both 

uncertainties.

ABSENCE OF REVERSAL

The unreversed raypaths make it impossible to distinguish 

between layer velocity and dip and to uniquely locate velocity 

anomalies along raypaths. A thicker low-velocity (2.0 km/s) 

layer beneath stations LP6 to LP8 (fig. 1B), for example, could 

be used to model time delays introduced by the basinlike 

structure between the San Andreas and Calaveras faults on the 

Loma Prieta transect. Because of the constraints introduced 

by the multichannel-seismic-reflection lines, we do not view 

this absence of reversal as a major limitation in the upper 

crust in the offshore region, where structures are reasonably 

well known and layer dips appear to be low; however, the 

uncertainties introduced by the absence of reversal are far 

more significant in the midcrust to lower crust. An inherent 

tradeoff, for example, exists between the velocities within the 

midcrust beneath the Continental Shelf and slope, and those 

within the unnamed anticline and the 6.6- to 7.1-km/s layer. 

Decreasing velocities within the midcrust require increasing 

velocities in the unnamed anticline and oceanic crust, and 

conversely. Furthermore, velocities within the midcrust are 

unlikely to be significantly lower than those shown (6.0 and 

6.5±0.1 km/s) because this difference would require velocities 

within the 6.6- to 7.1-km/s layer to be increased above 7 km/s 

in the middle of this layer.

The dip of and depth to the Moho, and the average velocity 

within the crust above the Moho, are well determined from 

PmP reflections. PmP and midcrustal reflections constrain an 

inherent tradeoff between the dip and velocity of the oceanic 

crust. Decreasing the dip of the Moho requires increasing the 

velocities of the oceanic crust. PmP reflection times limit the 

range of permissible Moho dips to 8°–11°, in turn limiting the 

range of permissible velocities within the oceanic crust.

The absence of source points near the east end of the Loma 

Prieta transect introduces a major uncertainty into the presence 

or absence of significant midcrustal to lower-crustal differences across the San Andreas fault. The presence and position 

of a near-vertical velocity boundary across the San Andreas 

fault is poorly determined from the available onshore-offshore 

data. Because of the weak velocity gradient in the midcrust, 

the near-vertical 6-km/s isovelocity contour near the San 

Andreas fault on the Loma Prieta transect (bottom of fig. 5) 

can be moved horizontally several kilometers in either direction without noticeable effect on the calculated traveltimes.

Finally, the available onshore-offshore data simply do not 

allow us to determine the geometry of the Moho east of the 

San Andreas fault. The inferred 26-km crustal thickness in 

the vicinity of the San Andreas fault, however, agrees closely 

with that reported from seismic-refraction profiling (Walter 

and Mooney, 1982). The crustal thickness west of the San 

Andreas fault is thought to be reliable and accurate to within 

2 to 3 km. The implied presence of oceanic crust east of the 

fault is undetermined and is shown only to suggest that the 

wide-angle data are consistent with, but do not require, the 

presence of this layer. This apparent continuity of oceanic crust 

is an intriguing aspect of our model that deserves to be tested 

by additional seismic profiling. Interpretation of explosion 

data recorded onshore confirms that the top of oceanic crust 

may be traceable as far landward as the San Andreas fault 

(Catchings and Kohler, 1996), and suggests that a reflection 

from the top of oceanic crust may extend eastward beneath 

the fault.
E119

0 20 40 60

DISTANCE, IN KILOMETERS

C

B

10

20

0

DEPTH, IN KILOMETERS

T-X/6, IN SECONDS

2

4

San

Gregorio

fault

A

T-X/6, IN SECONDS

2

4

SW NE

Figure 8.—Synthetic (A) and observed (B) record sections and ray diagram (C) for receiver at station AN1 on the Año Nuevo 

transect (see fig. 1B for locations). Receiver was located at model coordinate 70 km. Same format as in figure 4. Pg arrivals are 

modeled as refractions from layer 5 that are truncated at depth by a low-velocity zone. Synthetic seismograms for these Pg arrivals 

are much too large, suggesting a weaker vertical-velocity gradient in layer 5 than that used to model data. Strong PiP reflections 

are modeled by precritical reflections from tops of layers 7 and 8. Large-amplitude arrivals at greater ranges are modeled by 

postcritical reflections and refractions from layer 9.

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E120 EARTHQUAKE OCCURRENCE

FAINT PiP REFLECTIONS

The PiP reflection is the most compelling argument in 

favor of the existence of a 6.6- to 7.1-km/s layer beneath the 

Continental Shelf and slope. If these reflections are ignored, 

first arrivals can be fitted to the conspicuous wide-angle PmP

reflection from a landward-dipping Moho with the same dip 

as that shown in figure 5 but located about 3 km higher than 

shown in figure 5. This Moho depth assumes that the 6.0- to 

6.4-km/s-velocity midcrust extends downward to the Moho. For 

many of the records at the farthest ranges, P3, Pn, and even Pg

arrivals show little time separation, and so we argue that they do 

not provide the most compelling argument in favor of a 6.6- to 

7.1-km/s layer. For stations at and east of the San Andreas fault, 

the calculated traveltimes indicate a significant time separation 

between these three arrivals (fig. 6); however, the quality of 

the data from these stations is generally insufficient to allow 

unequivocal identification of all three of these branches.

0 40 80 100

DISTANCE, IN KILOMETERS

120

C

B

10

20

0

DEPTH, IN KILOMETERS T-X/6, IN SECONDS

2

4

San Andreas

fault Calaveras

fault

A

T-X/6, IN SECONDS

2

4

20 60

San Gregorio fault

SW NE

Figure 9.—Synthetic (A) and observed (B) record sections and ray diagram (C) for receiver at station AN4 on the Año 

Nuevo transect (see fig. 1B for locations). Receiver was located at model coordinate 126 km. Same format as in figure 4. 

Pg arrivals are modeled by refractions from layer 6. Weak PiP reflections are explained by precritical reflections from tops 

of layers 7 and 8 used to model oceanic crust. PmP arrivals are modeled as reflections from top of layer 9 to west of the 

San Andreas fault in vicinity of the San Gregorio fault.
E121

The records from stations LP1 to LP5 and from stations 

AN1 to AN3 (fig. 1B) provide the best evidence for the PiP

reflection. Synthetic seismograms for the models shown in 

figure 5 suggest that PiP and PmP reflections have similar 

amplitudes at identical ranges (figs. 4, 6, 8) and that locally 

at wide angles the amplitudes of PiP reflections exceed those 

of PmP reflections. The predicted amplitudes of PiP and PmP

reflections at precritical incidence are much smaller than 

those of Pg arrivals. Both of these predicted amplitudes agree 

closely with the observed amplitudes.

The case for the presence and depth of oceanic crust between 

model coordinates 0 and 20 km is probably best made by seismic-reflection lines 32 and 38 (fig. 3). The depth of oceanic 

crust near model coordinate 10 km can be moved, on the basis 

of the wide-angle data themselves, a few kilometers up or 

down. The vertical seismic-reflection data, however, constrain 

the depth of oceanic crust to within 1 km.

DISCUSSION

The primary elements of the velocity structures inferred 

from forward modeling of P-wave-arrival traveltimes on 

onshore-offshore wide-angle data are (1) a landward-thickening wedge of rocks with velocities of 6.0 to 6.5 km/s and 

a low velocity gradient west of the San Andreas fault, corresponding to layer 6 (figs. 4–9); (2) a landward-dipping layer 

with velocities of 6.6 to 7.1 km/s extending as far eastward 

as the San Andreas fault and, possibly, farther (layer 8), also 

imaged by seismic tomography (Hole and others, 2000); (3) 

a landward decrease in average crustal velocities east of the 

San Andreas fault, in agreement with seismic-tomography 

models (for example, Hole and others, 2000); and (4) a crustal 

thickness of about 26 to 28 km beneath the San Andreas fault. 

These major structural elements, as well as the amplitudeversus-range characteristics of the arrivals, have important 

ramifications for the modeling of strong ground motions 

generated by earthquakes along the San Andreas fault and 

along subsidiary faults in the offshore region.

A well-constrained result of our modeling is a major difference in the upper-crustal structure west of the San Gregorio 

fault observed along the Loma Prieta and Año Nuevo transects. Velocities in layers 5 and 6 (figs. 4–9) west of the San 

Gregorio fault are significantly higher (max 1 km/s) along the 

Año Nuevo transect than along the Loma Prieta transect (fig. 

5). One possible explanation for this difference in velocity is 

that at the latitude of Año Nuevo, the shelf and upper slope 

are underlain by rocks of the Salinian block, whereas at the 

latitude of Loma Prieta, the shelf and upper slope are underlain by rocks of the San Simeon terrane, which is composed 

of Franciscan rocks.

Our preferred interpretation of the near-uniform velocities (from 6.0 to 6.5±0.1 km/s) within the structural wedge 

beneath the continental margin between the San Gregorio and 

San Andreas faults is that they represent a sliver of Salinian 

basement. These velocities are reasonable for quartz monzonite of the Salinian block but apparently too high for typical 

metamorphosed graywacke of the Franciscan Complex (Lin 

and Wang, 1980). Likewise, bearing in mind that these velocities are minimums for the wedge, they are also too high to 

be representative of Tertiary rocks of the accretionary wedge 

(Stewart and Peselnick, 1981). We note, however, that previous 

workers have placed the west boundary of the Salinian block 

at the Sur-Nacimiento fault, which on our transect is offset 

along the San Gregorio-Hosgri fault (Silver and others, 1971; 

Mullins and Nagel, 1981). The lower (5.9–6.2 km/s) velocities 

in the block west of the San Gregorio fault on the Loma Prieta 

transect are more consistent with Franciscan rocks, as proposed 

by these workers.

Possibly the most surprising result of our experiment is the 

evidence for a 5-km-thick slab with velocities of 6.8 to 7.1 

km/s beneath the Continental Shelf and slope extending as 

far landward as the San Andreas fault. We interpret this slab 

(layers 7, 8, figs. 4–9) as oceanic crust, owing to its thickness 

and velocity structure. Alternatively, it may represent a layer 

of material magmatically underplated to the base of the crust 

during subduction of the ridgecrest (Holbrook and others, 

1996). We prefer the first interpretation, however, because of 

the coincidence requiring that the layer’s thickness and velocity 

structure produced by magmatic underplating should resemble 

that of oceanic crust so strongly.

The presence of a slab of oceanic crust beneath the central 

California margin appears to conflict with the lithosphericslab-window hypothesis (Dickinson and Snyder, 1979). That 

hypothesis, however, neither predicts nor requires a slab 

window west of the San Andreas fault, because it predicts a 

slab window only east of the fault. Our evidence, in conjunction 

with evidence for subducted oceanic crust from deep seismic 

profiling by PG&E and EDGE off central California (Ewing 

and Talwani, 1991; Tréhu, 1991), and permissive evidence for 

oceanic crust from the BASIX line near San Francisco (fig. 1B; 

Holbrook and others, 1996), suggests that the outer continental 

margin of California is underlain by oceanic crust from Morro 

Bay to as far north as San Francisco, a distance of more than 

250 km. This relation is important because it demonstrates 

that the observed oceanic crust underneath central California 

cannot be explained simply as a local feature associated with 

subduction of the Monterey fragment of the Farallon plate, 

which is found to the south of Santa Cruz (Ewing and Talwani, 

1991). The wide-angle PmP reflections observed during our 

experiment are consistent with a landward dip of the oceanic 

crust of about 8°–11°.

We cannot, however, determine the identity of the oceanic 

crust from the seismic data alone. Nonetheless, because the oceanic crust imaged at wide angles appears to be continuous with 

the Pacific plate imaged by multichannel-seismic-reflection 

data on the west end of the profile, the oceanic crust beneath 

the continental margin probably represents crust of the Pacific 

plate. Page (1991) and Page and Brocher (1993) proposed that 

at least part of this oceanic crust was thrust beneath North 

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E122 EARTHQUAKE OCCURRENCE

America during Pliocene-Quaternary plate reorganization and 

extension in the Basin and Range Province. They calculated 

that at least 15 to 25 km of convergence between the Pacific 

plate and North America occurred during this interval.

We suggest that the Salinian block and Franciscan rocks are 

in thrust contact with the underlying slab of oceanic crust. 

We postulate that this thrust accommodates overthrusting of 

the oceanic crust by North America (fig. 10), as suggested on 

the basis of observed incipient overriding of the sedimentary 

deposits in the trench by the toe of the continental margin 

(McCulloch, 1987; Page, 1991; Page and Brocher, 1993). 

This tectonic wedge is the structural mirror image of the 

tectonic wedge observed east of the San Andreas fault, composed of Franciscan rocks, above which rocks of the Great 

Valley sequence have been uplifted (Wentworth and others, 

1984; Wentworth and Zoback, 1989). We support the model 

of Eaton and Rymer (1990) that on either side of the San 

Andreas fault, tectonic wedges accommodate fault-normal 

compression in the crust by outward thrusting of large, rigid 

blocks along low-angle decollement surfaces. Onlap relations 

of the young sedimentary deposits on the unnamed anticline 

along seismic-reflection lines 32 and 38 (figs. 1B, 3) provide 

evidence for active thrusting within the unnamed anticline, 

supporting the inference that fault-normal compression is 

accommodated by low-angle thrusts. The inferred mirror 

symmetry of tectonic wedges across the San Andreas fault 

is consistent with the observation of fold-and-thrust belts on 

both sides on the fault that have opposite senses of vergence 

(Namson and Davis, 1988). Indeed, the opposite senses of 

thrusting on either side of the fault require material flow 

toward the fault, consistent with the proposed tectonic-wedge 

model of Eaton and Rymer (1990).

Comparison of the velocity structure derived from forward 

modeling with the hypocentral and aftershock data (U.S. 

Geological Survey staff, 1990) suggests that the Loma Prieta 

rupture initiated near the base of this structural wedge. The 

hypocenter apparently ruptured the highest velocity rocks 

near the base of the Salinian block at about 18-km depth (fig. 

5). This observation agrees with previous spatial correlations 

of high-velocity anomalies and earthquake rupture zones 

(Michael and Eberhart-Phillips, 1991).

The wave-propagation characteristics of the wide-angle 

data and complex crustal structure revealed by our analysis 

of these data have important implications for the modeling of 

strong ground motions produced by earthquakes along the San 

Andreas and subsidiary faults. The concave-upward geometry 

of the Moho strongly focuses seismic amplitudes and thus will 

locally enhance ground motions in ways that can be anticipated 

by knowing the likely locations of future earthquakes (fig. 11). 

Upper-mantle PmP reflections are conspicuous and generally 

of greater amplitude than Pg arrivals, and so understanding 

the propagation of PmP arrivals is vital for understanding 

strong ground motions (Somerville and Yoshimura, 1990; 

Catchings and Kohler, 1996). (PiP amplitudes are generally 

smaller than PmP arrivals but locally can be quite strong.) The 

complex Moho geometry may help explain the localization of 

strong ground motion observed in northern San Francisco and 

Oakland during the 1989 Loma Prieta earthquake (Somerville 

and Yoshimura, 1990) and in subsequent seismic-refraction 

studies (Catchings and Kohler, 1996).

SUMMARY

We have used onshore recordings of airgun profiling offshore 

Santa Cruz and Año Nuevo to develop a model for the major 

structural elements beneath the continental margin west of the 

San Andreas fault in the Loma Prieta region. Our most important result is that wide-angle seismic-reflection data provide 

evidence for a slab of oceanic crust extending landward beneath 

the continental margin as far eastward as the San Andreas fault. 

These data require a landward-dipping oceanic crust and Moho 

dipping 8°–11° E. We interpret this oceanic crust as a slab of 

Pacific plate that was overthrust by North America during recent 

oceanward thrusting probably related to crustal extension in 

the Basin and Range Province. The data are consistent with 

the presence of a block of Salinian rocks, with velocities consistent with either granite or quartz monzonite, above the slab 

of oceanic crust, between the San Gregorio and San Andreas 

faults. This block forms part of a structural wedge composed 

of a sliver of Salinian rocks, the Tertiary accretionary prism, 

and, possibly, a piece of Franciscan terrane. Our interpretation 

is consistent with a mirror image across the San Andreas fault 

with respect to the existence and geometry of tectonic wedges 

and fold-and-thrust belts. Oceanic crust is also observed beneath 

the continental margin to the south near Morro Bay, and so a 

tectonic wedge west of the San Andreas may extend along the 

fault for hundreds of kilometers. The concave-upward geometry 

of the Moho in the vicinity of the San Andreas fault should 

strongly focus those wide-angle reflections from the Moho 

that travel along the fault from earthquake sources on the San 

Andreas and subsidiary faults.

ACKNOWLEDGMENTS

This research was supported by the National Earthquake 

Hazards Reduction Program and the Offshore Geologic 

Framework Program. Numerous private landowners kindly 

gave us permission to locate 5-day recorders on their property. 

Kevin Shea of the East Bay Regional Park District and Chief 

Lucky of the California Department of Forestry provided 

access to property under their jurisdiction. John Coakley 

and others of the USGS in Menlo Park, Calif., deployed the 

5-day recorders. Donna Eberhart-Phillips projected the threedimensional velocity models from earthquake tomography 

onto our sections for incorporation into our two-dimensional 

velocity model. Gary Fuis, Jill McCarthy, Ben Page, and Carl 

Wentworth provided useful reviews of the manuscript. This 

paper is dedicated to the memory of Ben Page.
E123

0 20 40

KILOMETERS

Tectonic

 wedge

Franciscan

 wedge

Ongoing subduction 

or pseudosubduction

of oceanic crust caused 

by component of faultnormal compression

SW NE

San Andreas

 fault

Moho

Oceanic crust

? ?

?

10 30

Moho

Focusing 

of rays

S. F.

Bay

San

 Andreas

 fault

Surface

Hypocenter

Rays

Figure 10.—Tectonic model of velocity structure, showing westward-directed tectonic wedge west of the San Andreas 

fault, similar to that proposed by Eaton and Rymer (1990). This wedge is inferred to be in thrust contact with underlying 

oceanic crust, which is undergoing subduction or pseudosubduction (Page, 1991; Page and Brocher, 1993) beneath North 

America. Also shown is eastward-directed Franciscan wedge east of the San Andreas fault, as described by Wentworth and 

others (1984) and Wentworth and Zoback (1989). This wedge uplifts rocks of the overlying Great Valley sequence. Eastward 

extent of oceanic crust, east of the San Andreas fault, is unknown. Heavy solid lines, seismogenic parts of crustal wedges; 

heavy dashed line, aseismic ductile/brittle transition. Vertical projection of the San Andreas fault downward through oceanic 

crust is unknown and purely speculative. No vertical exaggeration.

Figure 11.—Schematic perspective view of the Moho in central coastal California, showing earthquake-generated rays 

propagating along the San Andreas fault. These rays, which initially diverge away from hypocenter of 1989 Loma Prieta 

earthquake, are reflected by curved Moho and become focused close to the San Andreas fault. Ground motions will thus 

be amplified near the San Andreas fault because of curved Moho geometry beneath fault.

EVIDENCE FOR OCEANIC CRUST BENEATH THE CONTINENTAL MARGIN
E124 EARTHQUAKE OCCURRENCE

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E126 EARTHQUAKE OCCURRENCE
E127

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

By Eylon Shalev, Duke University; 

and 

Jonathan M. Lees, University of North Carolina, Chapel Hill

CONTENTS

 

 Page

Abstract--------------------------------------------------------------------- 127

Introduction---------------------------------------------------------------- 127 

Geologic setting----------------------------------------------------------- 127 

Data selection-------------------------------------------------------------- 129 

Method---------------------------------------------------------------------- 129 

 Model parametrization---------------------------------------------- 129 

 Nonlinear tomographic inversion--------------------------------- 130 

 Resolution and variance-------------------------------------------- 131 

Initial model---------------------------------------------------------------- 132 

Results---------------------------------------------------------------------- 133 

Discussion------------------------------------------------------------------ 136 

Conclusions---------------------------------------------------------------- 139 

Acknowledgments-------------------------------------------------------- 139 

References cited----------------------------------------------------------- 139 

—————

ABSTRACT

A high-resolution tomographic study, using cubic B-spline 

parametrization, provided a three-dimensional P-wave-velocity map of the Loma Prieta region. A total of 11,977 highquality raypaths from 844 aftershocks of the 1989 Loma Prieta 

earthquake were used in the inversion. The velocity model 

exhibits a low-velocity feature between the San Andreas and 

Zayante-Vergeles faults in the uppermost 10 km of the crust. 

This low-velocity feature is interpreted as a sedimentary unit 

exposed to the northwest and separated from the Salinian 

block by the Zayante-Vergeles fault. Below 10-km depth, no 

consistent change in velocity is observed between the Salinian and Franciscan blocks. A high correlation exists between 

aftershock activity and localized high-velocity anomalies 

southeast of the main shock. Whereas these anomalies may 

represent brittle rocks associated with a fault-zone asperity 

that failed after the main shock, they also suggest a body of 

serpentinite. The serpentinite exhibits high seismic velocities and is potentially less competent than the surrounding 

country rock, thus providing a sector along the fault more 

likely to be associated with many smaller earthquakes or 

creep behavior.

INTRODUCTION

The 1989 Loma Prieta earthquake has provided the seismological community with a unique opportunity to investigate 

the relation between seismic-velocity distributions, earthquake 

seismicity, and geologic structures along the San Andreas fault. 

The velocity distributions provide insight into the structural 

details of the subsurface, in addition to serving as the basis for 

future analyses of wave propagation in regions of major strikeslip faults. An integrated analysis, using seismic structures, 

geologic interpretation, and statistical analysis of seismicity, 

may ultimately provide a model for seismic-hazard estimation. In this paper, we present a detailed, three-dimensional 

tomographic analysis of P-wave-velocity variations, suggest 

several possible interpretations, and discuss the implications 

of these interpretations with regard to seismic hazard.

GEOLOGIC SETTING

The geology of the Loma Prieta region has been summarized 

in recent literature (Aydin and Page, 1984; Saleeby, 1986; 

Fuis and Mooney, 1990), and so we provide only a brief 

overview here. Central California crustal structure consists 

of parallel northwest-striking fault-bounded blocks (fig. 1). 

The San Andreas fault near Loma Prieta separates the Salinian block from an internally complex block that is composed 

mostly of rocks of the Franciscan assemblage, a late Mesozoic 

complex (Fuis and Mooney, 1990). The Salinian block is an 

orphaned granitic-basement terrane that was transported at 

least 500 km northwestward by the San Andreas fault system 

(Ross, 1978). The part of the Salinian block between the San 

Andreas and Zayante-Vergeles faults is considered different 

from the main part of the Salinian block. Ross postulated that 

this part is a gabbroic subdivision of the Salinian block. Clark 

and Rietman (1973) suggested a structure of sandstone and 

mudstone reaching 4 to 5 km deep. Between the San Andreas 

and Calaveras faults, Franciscan rocks of oceanic origin are 

overlain by a thin sliver of Coast Range ophiolite and Great 

Valley sequence (McLaughlin and others, 1988).

Seismic-refraction profiles of the Loma Prieta region were 

summarized by Fuis and Mooney (1990). In the Salinian block, 

the crust can be divided into layers, with a midcrustal interface 

at 10-km depth (Walter and Mooney, 1982). An alternative 

interpretation, also offered by Walter and Mooney, is a subdivision of the crust below 10-km depth into two sublayers. Ross 

and McCulloch (1979) postulated that the uppermost 10 km 

of crust of the Salinian block is detached from the basement 

THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION

E127
E128 EARTHQUAKE OCCURRENCE

and that the crust below 10-km depth might correspond to the 

Franciscan complex. Fuis and Mooney noted, on the basis of 

the seismic velocity below 10-km depth, that the midcrustal 

and lower-crustal rocks are most likely not Franciscan. A 

seismic-refraction profile across the San Andreas, Sargent, 

and Calaveras faults 15 km southeast of Loma Prieta (Mooney 

and Colburn, 1985) shows heterogeneous upper-crustal velocities in the uppermost 3 km of the crust. No velocity change 

across the San Andreas fault was noted in that study. A strong 

reflector at 8- to 9-km depth is visible, but the seismic velocity 

below this depth was unresolvable.

After the 1989 Loma Prieta earthquake, Dietz and Ellsworth (1990) used the main shock and 89 aftershocks, evenly 

distributed with a maximum depth of 20 km, to calculate a 

joint hypocenter-velocity inversion for the two sides of the 

San Andreas fault. Their two velocity models differ in the 

uppermost 9 km of the crust, where the velocities southwest 

of the fault are lower. Below 9-km depth, the two models are 

nearly identical. Velocities below 18-km depth were poorly 

resolved because of inadequate sampling.

Previous three-dimensional tomographic studies of the study 

area (fig. 1) were conducted by using aftershocks of the 1989 

0 5 10 KILOMETERS

36°

50"

37°

00"

37°

10"

37°

20"

37°

30"

122°10" 122°00" 121°50" 121°40" 121°30" 121°20"

SAN ANDREAS FAULT

SARGENT FAULT

Target area HAYWARD FAULT

CALAVERAS FAULT

PACIFIC OCEAN

S.F. 

 Bay

Santa Cruz

Main shock

Salinian block

Franciscan rocks

HCBM

HPRM

JBZM

JPLM

JTGM

ZAYANTE-VERGELES FAULT

Figure 1.—Loma Prieta region, Calif., showing locations of major faults and study area (rectangle). Circles, epicenters; triangles, 

recording stations; solid triangles, stations used in inversion. Five named stations had unusually large station corrections (asterisks, 

table 1).
E129

Loma Prieta earthquake (Eberhart-Phillips and others, 1990; 

Lees, 1990; Michael and Eberhart-Phillips, 1991; Lees and 

Shalev, 1992; see Eberhart-Phillips and Michael, this chapter). 

Those studies showed a marked variation in P-wave velocity 

across the San Andreas fault in the vicinity of the main shock 

from the surface to 10-km depth. The northeast side of the 

fault has seismic velocities nearly 1.0 km/s higher than the 

southwest side, with the highest contrast extending to 4- to 6-

km depth. This pattern changes below 10-km depth, where the 

two sides of the fault appear to be similar. The velocity contrast 

across the fault is noticeable mainly in the vicinity of the main 

shock and is not observed elsewhere along the fault.

DATA SELECTION

The data used in this study were selected from the U.S. 

Geological Survey (USGS) catalog of 2,775 events recorded 

from October through December 1989. All traveltime picks 

and observation weights were made by USGS scientists. We 

chose only those events that were located inside our target 

volume which had at least 10 readings from among the 27 

recording stations included in the study area (fig. 1). Events 

with root-mean-square (rms) residuals of greater than 0.15, 

horizontal errors of greater than 1.0 km, and vertical errors of 

greater than 2.0 km were excluded, as were all observations 

marked by the original observer as poor-quality picks. The 

remaining 844 events produced 11,977 high-quality raypaths 

suitable for tomographic inversion.

METHOD

MODEL PARAMETRIZATION

The slowness structure of the target volume was parametrized by using cubic B-spline functions (Firbas, 1987; 

Michelini and McEvilly, 1991; Lees and Shalev, 1992; Shalev, 

1993a), which have an implicit smoothness and no artificial 

discontinuities of velocity (as in block models) or velocity 

gradients (as in linear-interpolation models). The B-spline 

model is defined by a discrete number of spatially distributed 

points (nodes), and the slowness field at the points provides 

the coefficients for B-spline interpolation. A total of 64 points 

are needed to calculate the slowness at any given position 

inside the target volume. One advantage of using cubic Bsplines is the ability to represent a complex structure with 

only a few nodes (Firbas, 1987); one disadvantage is that the 

locality of B-spline functions may entail underparametrization of the model in places where the raypaths are highly 

concentrated. To address this problem, we used a large concentration of nodes in the vicinity of the San Andreas fault, 

with a smaller node spacing than in the rest of the target 

volume. Furthermore, closer spacing in the vicinity of the 

fault allows a cubic B-spline structure to display rapid gradient changes (Shalev, 1993b). We used 18×22×12 nodes in xcoordinates (perpendicular to the fault), y-coordinates (along 

the fault), and depth coordinates, respectively, for a total 

of 4,752 nodes, to model the 45×65×25-km target volume 

(fig. 2). The coordinate system was rotated so that the San 

Andreas fault trends primarily in the y-coordinate direction 

(fig. 1). Regular node spacing in the x (perpendicular)- and 

Figure 2.—Distribution of nodes (dots) in study area (fig.1): map 

view (A) and across-strike cross section (B). Lines A–A' through 

D–D' are traces of cross sections shown in figure 10.

36°50"

37°00"

37°10"

121°40"

121°30"

N

A A′

B B′

C C ′

D D′

0

5

10

15

20

25

0 5 10 KILOMETERS

DEPTH, IN KILOMETERS 

B

A

THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION
E130 EARTHQUAKE OCCURRENCE

y (strike)-coordinates was 3 km, with 1.5-km spacing on the 

x-coordinate near the fault.

Rays in the three-dimensional model were traced by using 

a modified version of the ray-bending algorithm of Um and 

Thurber (1987). Using their algorithm, we found it important 

that the same model parametrization be used for both ray tracing and tomographic inversion. If we use a velocity structure 

for ray tracing and a slowness structure for tomographic inversion, an additional nonlinearity is introduced (Shalev, 1993a). 

Therefore, we modified the original ray-bending algorithm to 

work with slowness.

NONLINEAR TOMOGRAPHIC INVERSION

P-wave-arrival data were used in a joint hypocenter, station-correction, and three-dimensional slowness inversion. We 

added station-correction parameters to the linearized equation 

of Thurber (1983) relating the traveltime residual of the kth 

ray, rk, to the model parameters:

 

 (1)

where ∆t

s

 is the perturbation to the station correction (positive 

station correction=low velocity near the station or at higher 

elevation); ∆te, ∆xe, ∆ye, and ∆ze are the perturbations to the 

hypocenter parameter; ∂t/∂xe, ∂t/∂ye, and ∂t/∂ze are the partial 

derivatives of the traveltime with respect to hypocenter coordinates; N is the number of slowness parameters; and ∆an is the 

perturbation to the slowness parameter n. The partial derivative of the traveltime with respect to the slowness parameter

∂tk /∂an (Shalev, 1993a) is defined as

 (2)

where the integration is along the path of kth ray and the threedimensional cubic B-spline function bn(a) is defined as

 (3)

where a is the position vector; N1, N2, and N3 are the number 

of nodes (functions) in x-, y-, and z-coordinates; and the B

terms are the one-dimensional B-spline functions defined by 

de Boor (1978).

Station-correction parameters were added to the model space 

to automate the iterative, nonlinear inversion. Without a stationcorrection adjustment after each iteration, some of the data that 

could have been attributed to station-correction perturbations 

are mapped to structural perturbations directly beneath the stations. Because cubic B-spline structures are less local than block 

or linear-interpolation structures, those structural perturbations 

could propagate downward through the model.

To minimize the adverse affects of mapping noise into the 

models, we employed a damped-least-square approach (Crosson, 1976; Van Der Sluis and Van Der Vorst, 1987). Equation 

1 includes three different types of model parameter: station 

correction, hypocenter location, and slowness. We found no 

theoretical justification for using the same damping parameter 

for each of these three types of model parameter. Following 

Spencer and Gubbins (1980), we chose three different damping parameters for each of the three types of model parameter. 

Each parameter class can be made more significant by lowering its damping, and less significant by increasing its damping. 

For example, setting the damping of the hypocenter location 

and the station correction to infinity leads to a damped-leastsquare slowness inversion. The tradeoff of this approach is 

the difficulty in determining the correct damping for each 

parameter class.

In addition to damping, we further constrained the slowness 

model by requiring the three-dimensional Laplacian of the 

model (that is, the slope of local gradients) to be small (Lees 

and Crosson, 1989; VanDecar, 1991), effectively smoothing the model. Without Laplacian smoothing, the slowness 

model can develop a random-noise appearance, alternating 

small areas of low and high velocity (Herman, 1980). This 

phenomenon is important primarily when using a large number 

of model parameters where individual parameters may be 

less constrained by the data. Because we are smoothing in 

three dimensions, the Laplacian constraint introduces three 

smoothing parameters. Although it is reasonable to assume 

that the two horizontal smoothing weights may be the same, 

choosing the same weight for the depth dimension appears 

to oversmooth the model vertically. The Laplacian is a local 

constraint, and so we expect that, at least on local scales, 

there may be more vertical heterogeneity than horizontal. 

We therefore differentiated between horizontal and vertical 

smoothing by using different parameters for each smoothing. 

This approach leads to a block-matrix equation with five different damping and smoothing parameters: station-correction 

damping, hypocenter damping, slowness damping, horizontal 

smoothing, and vertical smoothing.

Not all the observations have the same accuracy, and so some 

downweighting of data with high uncertainties is desirable. 

Following the standard approach (Crosson, 1976; Menke, 

1984), we multiplied each row of the matrix equation by the 

inverse of the estimated uncertainty of observation. The final 

weighted block-matrix equation is

 

 (4)

where W is the weighting matrix; A, B, and C are the partial-derivative matrices of station corrections, hypocenter 

parameters, and slowness parameters; I is the identity matrix; 

x1, x2, and x3 are the perturbations to the station corrections, 

r t t

t

x x ty

y

t

z

z

t

a k e

k

e

e

k

e

e

k

e

e

k

n

n

n

N

= + + + + +

=

∆ ∆ s ∆ ∆ ∆ ∆ ∑ ∂

∂

∂

∂

∂

∂

∂

∂ a

1

,

∂

∂ = ∫ t b dl k

n n

k

a ( ) a ,

ray

b B x B y B z n i j l

l

N

j

N

i

N

( ) a = ( ) ( ) ( ),

= = =

∑∑∑ 1 1 1

1 2 3

WA WB WC

I

I

I

L

Wr

λ

λ

λ

λ λ

1

2

3

4 5

1

2

3

0 0

0 0

0 0

0 0

0

0

0

( , ) 0

,

x

x

x

=
E131

hypocenter parameters, and slowness parameters, respectively; 

λ1, λ2, and λ3 are damping parameters; λ4 and λ5 are the horizontal and vertical smoothing parameters, respectively; and 

r the traveltime-residual vector. To illustrate the Laplacian 

operator L, consider one row of this operator associated with 

node ci,j,k in Cartesian coordinates:

 (5)

Objectively choosing five independent damping and 

smoothing parameters for a large inversion is a complex 

problem. To avoid a totally subjective choice (which may be 

severely biased), a penalty function is sought that represents 

the merit of choosing a particular combination of these five 

parameters. A simple χ2 test will not work because the minimum χ2 score corresponds to no damping. Instead, we chose 

a crossvalidation score (Golub and others, 1979; Inoue and 

others, 1990) as a penalty function. The crossvalidation function can be used to choose the best model among competing 

models (Davis, 1987), and can indicate how well the computed 

model may predict traveltimes in the model area. The data are 

randomly divided into two subsets, say, S1 and S2. Subset 

S1 is excluded from the inversion, and subset S2 is used to 

conduct a full nonlinear inversion for a given set of damping 

parameters λ. After the inversion, the crossvalidation score 

is computed from subset S1:

 

 (6)

where f(λ) is the crossvalidation score, N is the number of 

data in subset S1, di

o

is the observed traveltime, and di

e

is the 

estimated traveltime in the model associated with λ.

The crossvalidation score displayed a monotonic increase 

with respect to the horizontal smoothing parameter (λ4) and 

therefore failed to provide an estimate for this parameter. For 

this reason, we chose the horizontal smoothing parameter (λ4) 

to be the smallest value that appears to eliminate the random 

model fluctuations discussed above. The other four damping 

parameters that minimized f (λ) were found by conducting a 

multiparameter search with a genetic algorithm (Goldberg, 

1989; Sen and Stoffa, 1991; Stoffa and Sen, 1991; Shalev, 

1993a), which is an efficient method of optimization that 

searches a multiparameter space of unknown shape, seeking 

a minimum (or maximum), using a small number of model 

evaluations relative to such arbitrary searches as Monte Carlo 

techniques. We used binary strings of 25 bits to represent 

models with four parameters. With a population size of 30 

and 25 generations, an average genetic-algorithm run was 

completed in 10 days on a Sun SPARC–2 computer. The 

damping and smoothing that parameters that minimized the 

crossvalidation score for this data set were λ1=46, λ2=70, 

λ3=6, λ4=30, and λ5=11. The numerical values of the damping parameters depend on the scaling of the matrices in 

equation 4.

The large number of raypaths, model parameters, and damping parameters yielded a very large system of simultaneous 

equations (28,820×10,123 in our case), which were solved 

with the LSQR algorithm (Paige and Saunders, 1982; Nolet, 

1987). After solving these equations, we retraced rays from 

the relocated events in the new, three-dimensional structure, 

computed a new set of equations, and iterated. The iterations 

were carried out until the time-residual-variance reduction of 

the last step was less than 2 percent (typically, four iterations). 

Peterson and Davey (1991) suggested that crossvalidation 

methods are superior to change of variance as an indicator 

for convergence of a nonlinear inversion. Our method incorporates the two methods by finding, through a crossvalidation 

technique, the best set of dumping parameters that fits the data 

when the stopping criteria is the change in variance reduction. 

The total variance reduction from the best one-dimensional 

starting models was 49 percent.

RESOLUTION AND VARIANCE

Providing a measure of the reliability of inversion models 

entails the formation and analysis of resolution and covariance 

matrices (Aki and Richards, 1980). In this study, however, the 

large size of the matrix equation makes an explicit calculation of resolution and covariance prohibitive. Furthermore, 

since the nonlinear inversion is iterative, the standard matrix 

approach to resolution and error estimates is inappropriate. 

For this reason, we must estimate the resolution and variance by other, approximate means. One such approach is to 

examine the impulse response (point-spread function) of the 

matrix system for each individual slowness parameter as an 

estimate of the resolution for those parameters (Humphreys 

and Clayton, 1988). Because this procedure would require 

thousands of inversions, we compromised by considering 

a response to a checkerboard test pattern (Spakman and 

Nolet, 1988; Inoue and others, 1990) augmented by impulse 

responses for two individual slowness parameters at locations 

of interest in the model (fig. 3). A checkerboard test pattern 

with a 12-km horizontal wavelength was tested with the same 

source-receiver geometry and damping and parameters used 

in the real data inversion. We observed that in the vicinity of 

the San Andreas fault, where the ray coverage is high, the test 

pattern is reproduced with little distortion. Using this synthetic 

test, we found that the average event relocation was approximately 0.001 km, with an origin-time perturbation of 0.001 s 

and an average station correction of 0.004 s. These errors are 

smaller by a factor of 10 than the hypocenter relocations and 

station-delay corrections calculated in the real data inversion. 

These small values indicate that very little of the slownessperturbation signal is mapped into the parameters associated 

with hypocenter locations and station delays. The small size 

of these corrections is due primarily to the heavy damping 

THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION

λ4 1 1 4 1 1 c c c c ( i j − + , , k i + , ,j k ) + + λ ( i j , , − + k i, , j k )

+ + λ5 1 (c c i j , ,k i − + , ,j k 1 4 ) − + ( ) 4 2 λ λ5 ci j , ,k = 0

f N

d d i i

i

N

( ) � = [ ] � ( ) , � �

�

�

�

�

�

�

� =

�1 2

1

1

2

o e
E132 EARTHQUAKE OCCURRENCE

Figure 3.—Resolution-test results. Model is displayed as percentage of perturbation from reference slowness, where darker 

regions represent lower velocity and lighter regions represent 

higher velocity. Top row shows synthetic checkerboard model 

at 6- and 11-km depth. Center row shows results of nonlinear 

inversion for the same depths; scale for checkerboard perturbation is ±6 percent. Bottom row shows two impulse responses 

at 8-km depth; scale for impulse-response perturbation is ±4 

percent.

imposed on the hypocenter and station-correction parameters, 

providing an estimate of the coupling noise between the slowness parameters and the other model parameters. We conclude 

that only a small part of the signal in the hypocenter locations 

comes from coupling to the slowness parameters and that the 

rest is due to the signal required by the data.

Estimation of the model variance is accomplished by using 

a statistical resampling technique called the jackknife (Efron, 

1982). This technique has been used previously to estimate 

the slowness-parameter variance in linear inversions (Lees 

and Crosson, 1989). Extending the jackknife to nonlinear 

analysis is straightforward. The complete nonlinear inversion 

analysis, including all the intermediate linear approximations, 

is treated as a single inversion, and the variances of all the 

model parameters are computed as described by Lees and 

Crosson (1989). We partitioned the data into 30 random, 

nonoverlapping subsets and performed each “mini- inversion” 

with the same constraining parameters as the main inversion. 

This approach produced variance estimates for the slowness 

parameters (fig. 4), as well as for event relocations and station 

corrections. The jackknife approach is advantageous in that it 

makes no assumptions regarding the statistical distributions of 

errors in the data or model parameters. Furthermore, consistent estimates of variance are provided for all model variables, 

regardless of the number of nonlinear iterations.

INITIAL MODEL

Although the San Andreas fault represents a major tectonic 

boundary separating two distinct blocks, we used a onedimensional, vertically varying velocity model as a reference 

starting model because slowness contrast across the fault is not 

always evident. Some studies suggest a significant seismicvelocity variation across the fault (Dietz and Ellsworth, 1990), 

whereas others indicate little or no variation (Mooney and 

Colburn, 1985; Fuis and Mooney, 1990). Three-dimensional 

tomographic studies (Eberhart-Phillips and others, 1990; 

Lees, 1990; Michael and Eberhart-Phillips, 1991; Lees and 

Shalev, 1992) suggested a marked velocity contrast across 

the fault in some areas, but this contrast was inconsistent 

along the full length of the fault. For this reason, we selected 

a one-dimensional, vertically varying model, assuring that all 

horizontal variations of slowness are warranted by the earthquake data used in our study. The large number of raypaths 

in the vicinity of the San Andreas fault assures that the area 

of the fault is well sampled and has a high resolution. Lees 

and Shalev (1992) showed that when a large number of highquality data from the Loma Prieta aftershock sequence are 

used, nonlinear tomographic-inversion results appear to be 

robust with respect to the initial reference model. The initial 

model has, however, a greater influence on those parts of the 

model that are more sparsely sampled. In our study, the initial 

model was one dimensional, and the associated station corrections that minimized the residuals of the full data set used 

36°50"

37°00"

37°10"

6.08 km/s 5.39 km/s

-6.0 % Slowness 6.0

Depth: 6.0 km

N

6.67 km/s 5.91 km/s

-6.0 % Slowness 6.0

Depth: 11.0 km

N

36°50"

37°00"

37°10"

6.08 km/s 5.39 km/s

-6.0 % Slowness 6.0

Depth: 6.0 km

N

6.67 km/s 5.91 km/s

-6.0 % Slowness 6.0

Depth: 11.0 km

N

36°50"

37°00"

37°10"

Depth: 8.0 km -4.0 % Anomaly 4.0

N

121°40"121°30"

D -4.0 % Anomaly 4.0 epth: 8.0 km

N

121°30"

0 5 10 KILOMETERS 121°40"
E133

the USGS station correction (L.D. Dietz, written communs., 

1992). This one-dimensional model (fig. 5) lies approximately 

between the two models previously constructed by Dietz and 

Ellsworth (1990).

RESULTS

Final perturbations to the initial station corrections are plotted in figure 6 and listed in table 1. We note a regional gradient 

associated with these corrections that varies linearly across 

the fault, from positive on the Franciscan side to negative in 

the Salinian block (fig. 6A). This pattern, which remained in 

the final station corrections, represents a broad regional variation across the San Andreas fault. Five stations were found to 

have large negative perturbations (fig. 1; table 1). The original 

station corrections had a positive correction in the vicinity of 

these five stations, and the inversion results all but removed 

this feature (fig. 6C). The inversion thus removed structures 

in the initial model associated with the station corrections and 

distributed that structure into the three-dimensional velocity 

model. We found that this result affected only the shallow part 

of our three-dimensional inversion results. Inversion models 

both with and without the station corrections differed only in 

the uppermost few kilometers near these five stations. We note 

that the deeper structure (2–10-km depth) in this region also 

exhibited an anomalously low velocity that was unaffected by 

including the station corrections.

The overall pattern of the station corrections can be interpreted as a possible long-wavenumber slowness variation that 

was mapped into the station corrections. To test this interpretation, we performed a synthetic test whereby we selected a 

model with a different velocity (Dietz and Ellsworth, 1990) 

for each side of the San Andreas fault. We traced rays through 

this model by using the same source-receiver geometry with 

zero station corrections. We also added random noise with 

0.1-s standard deviation to the synthetic data. Starting from 

a one-dimensional velocity structure whose slowness lies 

between the two sides of the synthetic model, we ran two 

test inversions: one with the same damping parameters that 

we determined for the real inversion, and the other with a 

large damping on the slowness perturbation (to simulate 

overdamping of the slowness model and mapping of regional 

slowness variation into the station corrections). The results 

were a very small perturbation of the station corrections in 

the first test (fig. 7A) and a smooth regional gradient in the 

second test (fig. 7B).

The results of the damped synthetic run (fig. 7B) show, 

as expected, that higher velocities (Franciscan assemblage) 

mapped into negative station corrections, and lower velocities 

(Salinian block) mapped into positive station corrections, in 

contrast to the station corrections plotted in figure 6. Furthermore, the synthetic station corrections plotted in figure 

7B are 25 to 40 percent of those plotted in figure 6. From the 

synthetic test, we can say that the station-correction gradient 

in figure 6C is a result of the data and not some large-waveFigure 4.—Jackknife error estimates for different depths in study 

area (fig. 1). Size of estimated standard deviations is displayed as 

percentage of reference slowness, where darker regions represent 

higher standard deviation and lighter regions represent smaller 

standard deviation. Scale for each layer is from 0 to 8 percent.

36°50'

37°00'

37°10'

 0.0 % Slowness 8.0 

Depth: 11.0 km

N

121°40'

121°30'

0 5 10 KILOMETERS

 0.0 % Slowness 8.0 

Depth: 15.5 km

N

121°40'

121°30'

THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION

36°50�

37°00�

37°10�

 0.0 % Slowness 8.0 

Depth: 1.5 km

N

 0.0 % Slowness 8.0 

Depth: 4.0 km

N

36°50�

37°00�

37°10�

121°40�

12130�

 0.0 % Slowness 8.0 

Depth: 6.0 km

N

121°40�

121°30�

 0.0 % Slowness 8.0 

Depth: 8.0 km

N
E134 EARTHQUAKE OCCURRENCE

length mapping of the regional velocity variation, and so we 

can resolve regional velocity change without adverse effect 

on the station corrections (fig. 7C).

The optimal choice of damping for hypocenter relocations 

(λ1, eq. 4), using crossvalidation as a merit function, forced 

the hypocenters to be constrained close to their original 

locations (fig. 8). We found that the hypocenter-parameter 

perturbations averaged 0.03 km and 0.01 s. Relaxing this 

constraint, however, indicated that the structural features of 

our model are robust with respect to relocations. We attributed the small hypocenter corrections to the initial choice of 

events. Only raypaths from well-defined events, with small 

residual rms and hypocenter-location errors, were chosen for 

the inversion. Therefore, for small velocity perturbations, 

only small hypocenter corrections were expected. As a test, 

we ran the inversion with a large damping on the slowness 

and station corrections and no damping on the hypocenter 

corrections. This inversion allows the hypocenter to move as 

if the entire signal in the data set is the result of hypocenter 

mislocation. The average hypocenter perturbations were 0.24 

km and 0.04 s, and none of the hypocenters moved more 

than 1 km. Variance reduction for this test was 7 percent, 

indicating that only a small part of the signal is attributable 

to hypocenter-location errors. Thus, the original hypocenter 

locations appear to be well defined.

To display the results of the tomographic inversion, we 

present plots of slowness perturbation both in horizontal 

slices (fig. 9), and in vertical cross sections perpendicular to 

the fault (fig. 10). For ease of interpretation, the vertical cross 

sections also display absolute velocity contours and projected 

hypocenter locations of events lying less than 3 km from the 

cross section.

The most outstanding feature present in the horizontal slices 

(fig. 9) is the velocity contrast across the San Andreas fault in 

the uppermost 4 km of the model. Velocities in the Salinian 

block are lower than in the Franciscan assemblage, in agreement with other tomographic studies in the Loma Prieta region 

(Eberhart-Phillips and others, 1990; Lees, 1990; Michael and 

Figure 5.—Three velocity models for the Loma Prieta region (fig. 

1): 1 and 2; velocity models for two sides of the San Andreas fault 

(Dietz and Ellsworth, 1990), translated to B-splines parametrization; 

3, least-squares model used as initial model in this study.

Figure 6.—Contour plots of station corrections in study area (fig. 1). A, Initial station corrections before 

inversion. B, Calculated perturbations from joint inversion. C, Final station corrections, which are sum 

of initial corrections and calculated perturbations. The San Andreas fault is shown as irregular heavy 

line on each contour plot for reference. Contour interval in figure 6B differs from that used in figures 

6A and 6C.

4.0 5.0 6.0 7.0 8.0

30

25

20

15

10

 5

 0

 DEPTH, IN KILOMETER

S

EXPLANATION

 1. Franciscan assemblage

 2. Salinian block

 3. Least-square-model

VELOCITY, IN KILOMETERS PER SECOND

0

0

0

0

0

0.2

0.2

0.2

0.4 0.4

0.4

0.6

0.6

0

20

40

60

DISTANCE, IN KILOMETERS

0 20 40

-0.4

-0.2

-0.2

�

0

0

0

0

0.05

0.05

0 20 40

DISTANCE, IN KILOMETERS

-0.1

-0.1

-0.05

-0.05

-0.05 -0.05

�

0

0

0

0

0

0.2

0.2

0.2

0.4

0.4

0.4

0.6

0.6

0 20 40

-0.4

-0.2

-0.2

0 2

�
THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION E135

Table 1.—Station corrections for the 27 stations used in the inversion

[Original correction, starting station correction as received from the U.S. Geological Survey; inversion correction cumulative correction 

of nonlinear inversion; total correction sum of the two former corrections; jackknife error standard deviation of inversion corrections 

as computed by jackknife method. Asterisks, station corrections that changed the most (see fig. 1 for locations)]

Eberhart-Phillips, 1991; Lees and Shalev, 1992; Eberhart-Phillips and Michael, 1997). At 4-km depth, the low-velocity zone 

in the Salinian block diminishes in size, and large velocity 

variations are not evident across the San Andreas fault at 

distances greater than 15 km in either direction along the fault 

from the main-shock epicenter. The low-velocity zone, shown 

as a southwesterly dipping low-velocity zone above the line 

of hypocenters in cross sections B–B' and C–C' (fig. 10), lies 

mainly between the San Andreas and Zayante-Vergeles faults. 

It decreases in size down to 10-km depth, where it appears as 

a small body, 4 to 6 km wide. Below 10-km depth, no clear 

velocity contrast across the San Andreas fault is evident.

An impulse-response test for a point 8 km deep beneath the 

San Andreas fault on cross section C–C' (at the location of the 

greatest velocity gradient across the fault, fig. 10) indicates 

approximately 3 km of spatial resolution perpendicular to the 

fault and about 6 km of spatial resolution parallel to the fault 

(fig. 3). Error analysis (fig. 4) shows only one major area with 

a large (>5 percent of the slowness size) standard deviation: at 

4-km depth in the low-velocity zone between the San Andreas 

and Zayante-Vergeles faults. The perturbation there, however, 

is almost double the size of the standard deviation—14 and 8 

percent, respectively—and so we conclude that a low-velocity 

zone is required by the data. The standard deviation in the rest 

of the model is less than 5 percent of the model slowness.

Another conspicuous feature in figures 9 and 10 is the 

high-velocity prism directly beneath the San Andreas fault 

along the southeast 10 km of the study area (cross section 

A–A'). This high-velocity prism, which separates two blocks 

of lower velocity, is visible from 6- to 14-km depth (there 

is no ray coverage below 14-km depth in this region). The 

high-velocity prism has a high correlation with hypocenter 

locations (fig. 10). The dip of the hypocenter plane changes 

from 70° in cross section B–B' to 85° in cross section A–A', 

and the associated high-velocity prism appears to follow the 

same trend. An impulse-response test for a point at 8-km depth 

beneath the San Andreas fault in cross section A–A' (in the 

upper part of the high-velocity prism) indicates approximately 

4 km of spatial resolution perpendicular to the fault and about 

6 km of resolution parallel to the fault (fig. 3). We note that 

this high-velocity prism persists when different horizontal 

Laplacian parameters are used, in contrast to the shallower 

low-velocity feature discussed above, which apparently is 

sensitive to changes in the smoothing.

Station Latitude

(N.)

Longitude

(W.)

Elevation

(km)

Original

correction

(s)

Inversion

correction

(s)

Jackknife

error

(s)

Total

correction

(s)

CCOM 37�15.46� 121�40.35� 0.366 0.850 �0.001 0.006 0.849

CSCM 37�17.11� 121�46.35� .128 .500 .017 .009 .517

HCAM 37�01.52� 121�29.02� .332 .260 .009 .004 .269

HCBM 36�55.88� 121�39.63� .219 .130 �.085* .014 .045

HCOM 36�53.31� 121�42.34� .129 � � �.018 .012 �.018

HCRM 36�57.46� 121�35.01� .241 .010 .012 .012 .022

HGWM 37�01.02� 121�39.20� .133 � � �.006 .009 �.006

HPRM 36�57.19� 121�41.70� .094 .100 �.105* .017 �.005

HSPM 37�06.91� 121�30.94� .850 .370 �.005 .003 .365

JALM 37�09.50� 121�50.82� .244 �.150 .022 .012 �.128

JBCM 37�09.62� 122�01.57� .660 �.080 �.058 .013 �.138

JBLM 37�07.69� 122�10.08� .792 �.370 �.005 .007 �.375

JBZM 37�01.07� 121�49.15� .213 .220 �.132* .022 .088

JCBM 37�06.71� 121�41.33� .192 .100 .027 .006 .127

JECM 37�03.04� 121�48.56� .438 .040 �.056 .015 �.016

JHLM 37�06.54� 121�49.99� .908 �.050 �.008 .010 �.058

JLXM 37�12.11� 121�59.17� .244 �.080 �.013 .009 �.093

JPLM 36�58.62� 121�49.93� .158 .070 �.113* .016 �.043

JRGM 37�02.22� 121�57.87� .213 �.280 �.028 .008 �.308

JRRM 37�03.27� 121�43.61� .408 �.070 .071 .011 .001

JSCM 37�17.07� 122�07.42� .357 � � �.032 .016 �.032

JSGM 37�16.96� 122�03.00� .198 .260 �.041 .021 .301

JSMM 37�12.74� 122�10.06� .262 � � .000 .010 .000

JSSM 37�10.17� 121�55.84� .946 �.060 �.004 .012 �.064

JSTM 37�12.41� 121�47.84� .149 .130 .034 .011 .164

JTGM 37�01.71� 121�52.58� .253 � � ��.112* .011 �.112

JUCM 37�00.07� 122�02.91� .177 �.470 �.017 .006 �.487
E136 EARTHQUAKE OCCURRENCE

The zone adjacent to the main shock appears to have a 

slightly higher velocity than the surrounding rocks. We 

observe this high-velocity zone at 11- to 16-km depth, with 

the main shock to the southwest. This high-velocity zone is 

not directly beneath the low-velocity zone in the uppermost 

10 km of the crust but is shifted about 3 km northeastward.

DISCUSSION

We emphasize that the absolute size of perturbations (and, 

therefore, the velocity) is partly a function of the particular 

damping parameters chosen (Lees and Shalev, 1992). We are 

therefore reluctant to associate specific rock types with given 

velocity anomalies. However, the spatial distribution of lowand high-velocity zones appears to be robust with respect to 

the initial model, damping parameters, and methods of parametrization. Thus, we believe that it is appropriate to interpret 

structural details relating to subsurface geology.

A schematic representation of a geologic interpretation of 

our results is shown in figure 11. The area between the San 

Andreas and Zayante-Vergeles faults has an apparent lowvelocity anomaly down to 10-km depth, in the same area as an 

observed negative gravity anomaly (Clark and Rietman, 1973). 

The low velocity and negative gravity suggest that this section 

of the Salinian block consists of sedimentary rocks (Clark and 

Rietman, 1973). Below 10-km depth, no consistent difference 

between the two sides of the San Andreas fault is evident.

The high degree of correlation between seismicity and high 

velocity along the San Andreas fault has been noted previously (Lees, 1990; Lees and Malin, 1990; Nicholson and Lees, 

1992). Cross section A–A', on the southeast edge of the study 

area (fig. 10A), reveals a high-velocity prism (nearly 1-km/s 

anomaly) along the plane of aftershocks below 6-km depth. 

This high-velocity prism gradually becomes less noticeable 6 

km to the northwest in cross section B–B' (fig. 10B). Farther 

northwest, in cross section C–C' (fig. 10C), the main shock 

is on the edge of a high-velocity anomaly restricted to greater 

than 11-km depth. In cross section D–D', 18 km northeast 

of the main shock (fig. 10D), we see no correlation of highvelocity anomalies and seismicity. We also see no evidence 

for a low-velocity fault-gouge zone in our results. Because 

the resolution in our study is about 3 km, a fault-gouge zone 

several hundreds meters wide would be unresolvable.

Some workers have suggested that the high-velocity zone 

in the vicinity of the main shock and subsequent aftershocks 

is associated with brittle rocks more likely to sustain high 

stress (Lees, 1990; Lees and Malin, 1990; Nicholson and Lees, 

1992). This suggestion implies that the high-velocity prism 

southeast of the main shock is an asperity and, as such, may 

represent the locus for future large earthquakes. This physical interpretation is based on the apparent pattern at several 

points along the fault. For the Loma Prieta region, however, 

Figure 7.—Synthetic test of different velocity structures for each 

side of the San Andreas fault in study area (fig. 1). A, Contour plot 

of station corrections with same damping and smoothing parameters 

used in real data inversion. B, Contour plot of station corrections 

with strong damping of velocity, so that most of signal is mapped 

into station corrections. C, Cross section A–A' (fig. 2). Inversion 

with best damping reproduced velocity step at 18 km, same as in 

true (synthetic) data.

Figure 8.—Earthquake corrections. A, Total distance corrections. 

B, Time corrections.

0

0

0

0

0

0.025

0.025

0

20

40

60

DISTANCE, IN KILOMETERS

0 20 40

DISTANCE, IN KILOMETERS

-0.025

�

� ��

0

0

0 0.025

0.025 0.025

0.05

0.05

0.05

0.075

0.075

0.075

0.1 0.1

0.125

0.15

0 20 40

-0.15

-0.125

-0.1

-0.1

-0.1

-0.075

-0.075

-0.075 -0.05

-0.05

-0.05

-0.025 -0.025

-0.025

4.5 4.5

5 5

5.5

5.5 6 6

10

5

DEPTH, IN KILOMETERS

5 10 15 20 25 30 35

DISTANCE, IN KILOMETERS

� ��

�

�

0

100

200

#

-0.06 -0.04 -0.02 0.00 0.02 0.04 0.06

TIME CORRECTION, IN SECONDS

�

�

0

100

200

#

0.000 0.005 0.010 0.015 0.020

DISTANCE CORRECTION, IN KILOMETERS
E137

an alternative explanation is that the high-velocity anomaly 

represents a prism of serpentinite (Allen, 1968).

Serpentinite is a common rock type, occurring in large 

sheets, in the Franciscan assemblage northeast of the San 

Andreas fault (Hanna and others, 1972). Although serpentinite samples from the Loma Prieta region exhibit P-wave 

velocities of only 5.5 to 5.6 km/s (N.I. Christensen, written 

commun., 1993), partially serpentinized ultramafic rock 

containing more than 50 volume percent serpentine showed 

higher P-wave velocities of 6.2 to 6.6 km/s (Christensen, 1966, 

1982). Thoroughly serpentinized ultramafic rocks typically 

have a very low resistivity (Parkhomenko, 1982), and low 

resistivities have been observed between the San Andreas and 

Zayante faults in the vicinity of the Loma Prieta main shock 

(Eberhart-Phillips and others, 1990).

The most distinctive attribute of serpentinized rocks may 

be their high Poisson’s ratio (Christensen, 1982; Gildner and 

Caruso, 1992), which can be detected with VP /VS inversion; 

however, a study of VP /VS structure in the Loma Prieta region 

(Thurber and Atre, 1993) does not have the resolution needed 

for accepting or rejecting the serpentinite-filled-fault model.

Wang (1984) dismissed the possible presence of serpentinite 

because of observed low velocities in gouge of the fault zone. 

Our study, however, reveals high velocities in the fault zone 

southeast of the main shock below 6- to 8-km depth, and so the 

serpentinite model is reasonable. If the high-velocity prism is, 

indeed, serpentinite, we recognize that this identification may 

contradict the fault-zone-asperity model suggested above, and 

so this high-velocity prism would represent a relatively weak 

part of the fault more likely to exhibit small earthquakes or 

36°50�

37°00�

37°10�

 5.42 km/s 4.35 km/s

-10.0 % Slowness 12.0 

Depth: 2.5 km

N

A ��

B B�

C ��

D ��

 5.87 km/s 4.71 km/s

-10.0 % Slowness 12.0 

Depth: 3.5 km

N

36°50�

37°00�

37°10�

121°40�

121°30�

 6.07 km/s 4.88 km/s

-10.0 % Slowness 12.0 

Depth: 4.5 km

N

121°40�

12130�

0 5 10 KILOMETERS

 6.26 km/s 5.03 km/s

-10.0 % Slowness 12.0 

Depth: 5.5 km

N

Figure 9.—Three-dimensional tomographic results of P-wave inversion in study area (fig. 1), displayed as horizontal slices 1 km apart. Model 

is represented as percentage of perturbation from initial slowness model, where darker regions represent lower velocity and lighter regions 

represent higher velocity. Scale for each slice is from constant −10 to 12 percent. See figure 2 for locations of cross sections in first slice.

THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION

36°50�

37°00�

37°10�

 6.45 km/s 5.18 km/s

-10.0 % Slowness 12.0 

Depth: 6.5 km

N

 6.65 km/s 5.34 km/s

-10.0 % Slowness 12.0 

Depth: 7.5 km

N

36°50�

37°00�

37°10�

121°40�

121°30�

 6.80 km/s 5.47 km/s

-10.0 % Slowness 12.0 

Depth: 8.5 km

N

121°40�

121°30�

0 5 10 KILOMETERS

 6.90 km/s 5.54 km/s

-10.0 % Slowness 12.0 

Depth: 9.5 km

N
E138 EARTHQUAKE OCCURRENCE

36°50�

37°00�

37°10�

 6.95 km/s 5.58 km/s

-10.0 % Slowness 12.0 

Depth: 10.5 km

N

 6.98 km/s 5.61 km/s

-10.0 % Slowness 12.0 

Depth: 11.5 km

N

36°50�

37°00�

37°10�

121°40�

121°30�

 7.00 km/s 5.62 km/s

-10.0 % Slowness 12.0 

Depth: 12.5 km

N

121°40�

121°30�

 7.01 km/s 5.64 km/s

-10.0 % Slowness 12.0 

Depth: 13.5 km

N

Figure 9.—Continued

 7.03 km/s 5.65 km/s

-10.0 % Slowness 12.0 

Depth: 14.5 km

N

36°50'

37°00'

37°10'

121°40'

121°30'

0 5 10 KILOMETERS

 7.06 km/s 5.67 km/s

-10.0 % Slowness 12.0 

Depth: 15.5 km

N

121°40'

121°30'

Figure 10.—Four cross sections perpendicular to the San 

Andreas fault in study area (fig. 1), from southeast (A–A') 

to northwest (D–D'). Shading represents percentage of perturbation from initial slowness model, where darker areas 

represent lower velocity and lighter areas represent higher 

velocity. Contour lines denote absolute velocity (that is, reference model plus perturbation) in cross section. Unshaded 

areas are those with no raypaths. Small circles, horizontally 

projected earthquakes in a ±3-km swath. SAF, San Andreas 

fault; ZVF, Zayante-Vergeles fault.

3.5

44.5

5 5.5

6

6.5

10

0

15

5

0 5 10 15 20 25 30 35

� ��� ��� ��

DISTANCE, IN KILOMETERS

DEPTH, IN KILOMETERS

3.5

44.5 55.5

6.5

6.5

10

0

15

5

� ��� ��� ��

3.5

4

4.5

5

5.5

6

6.5

10

0

15

5

� ��� ��� ��

3.5

44.55

5.5

6

10

0

15

5

� ��� ��� ��
E139

southeast of the main shock. This high-velocity prism does 

not exist northwest of the main shock (figs. 9, 10). The characteristics of the San Andreas fault may change in the Loma 

Prieta region, and the change of slip during the earthquake 

may have been geologically controlled. More detailed studies 

of the area southeast of the main shock are needed to confirm 

this change in the San Andreas fault.

CONCLUSIONS

In this study, we have calculated and analyzed joint velocity, 

hypocenter, and station-correction inversions, using the traveltimes from aftershocks of the 1989 Loma Prieta earthquake. 

We found a regional gradient in station corrections across the 

fault, with particularly high negative station-delay perturbations occurring in the Salinian block near and southeast of the 

main shock. The shallow (3–4-km depth) region of our model 

exhibits a distinct velocity variation across the San Andreas 

fault, although in deeper sections the structures appear to be 

more heterogeneous.

The velocity model contains a low-velocity zone between 

the San Andreas and Zayante-Vergeles faults in the uppermost 10 km of the crust. This low-velocity zone is interpreted 

as a structure of sedimentary rocks exposed to the northwest 

and separated from the Salinian block by the ZayanteVergeles fault. Below 10-km depth, no consistent change 

is observed between the Salinian block and the Franciscan 

assemblage.

There appears to be a high degree of correlation between 

aftershock activity and localized high-velocity anomalies in 

the area southeast of the main shock. Although this zone of 

high-velocity anomalies may represent brittle rocks associated 

with a fault-zone asperity that failed after the main shock, we 

observed no evidence to suggest that the zone may be a prism 

of serpentinite. The serpentinite exhibits high velocities and is 

potentially less competent than the surrounding country rock, 

providing a sector along the fault more likely to be associated 

with numerous smaller earthquakes or creep behavior.

We hypothesize that the geology of the San Andreas fault 

zone changes in the area of the main shock, from a serpentinite-filled zone southeast of the main shock to a hard-rock 

zone northwest of the main shock.

ACKNOWLEDGMENTS

This research was supported by U.S. National Science Foundation grants EAR–9011441 and EAR–8657206. We thank 

Lynn Dietz for her help in obtaining the original data, and 

Mark T. Brandon for helpful discussions and suggestions.

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creeping behavior. By this interpretation, the aftershocks are 

concentrated in the weaker prism of serpentinite southeast 

of the main shock. A more detailed magnetic and resistivity 

investigation may be able to distinguish between those two 

alternative and apparently contradictory interpretations.

Rupture-model investigations (Beroza, 1991; Hartzell 

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1991) of the main shock showed different slip characteristics 

along the fault, from predominantly strike slip southeast of 

the hypocenter to predominantly dip slip northwest of the 

hypocenter. This change in slip direction, which occurs within 

a small distance of 12 to 20 km, has raised questions about 

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A high-velocity prism interpreted as serpentinite appears 

Figure 11.—Schematic representation of the low- and high-velocity regions in study area (fig. 1) along vertical cross sections A–A'

through C–C' (fig. 2). SAF, San Andreas fault; ZVF, Zayante-Vergeles 

fault. Low-velocity region, sedimentary block between the San 

Andreas and Zayente-Vergeles faults; high-velocity region, thin 

prism of serpentinite that correlates with seismicity.

THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION
E140 EARTHQUAKE OCCURRENCE

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THREE-DIMENSIONAL TOMOGRAPHIC ANALYSIS OF THE LOMA PRIETA REGION
E142 EARTHQUAKE OCCURRENCE
E143

By Shashank R. Atre and Clifford H. Thurber, University of Wisconsin, Madison

CONTENTS

—————

 Page 

Abstract----------------------------------------------------------------------- 143

Introduction------------------------------------------------------------------ 143

S−P data---------------------------------------------------------------------- 144

Inversion method------------------------------------------------------------ 146

Inversion tests---------------------------------------------------------------- 150

Results of three-dimensional inversion----------------------------------- 152

Discussion-------------------------------------------------------------------- 154

Conclusions------------------------------------------------------------------ 160

Acknowledgments----------------------------------------------------------- 161

References cited------------------------------------------------------------- 161

—————

ABSTRACT

We present a three-dimensional model for variations in the 

ratio of P- to S-wave velocity (VP /VS ) along the Loma Prieta 

rupture zone, derived from portable-instrument S−P arrival 

times and earthquake locations determined in a three-dimensional P-wave-velocity model. We devoted significant effort 

to identifying S-to-P-converted waves, which were prevalent 

at a few of the stations used. The inversion method uses threedimensional ray tracing in an existing laterally heterogeneous 

P-wave-velocity model and a constant VP /VS ratio to determine 

the initial paths for S waves. The model parametrization allows 

spatial variations in VP /VS ratio on a three-dimensional grid 

with node spacings of 3 to 15 km horizontally and 3 to 5 km 

vertically. The reduction in data variance is nearly 65 percent 

for the single-iteration inversion.

We interpret the VP /VS model in conjunction with other 

available geologic and geophysical information. Our major 

findings are (1) high VP /VS ratios along the rupture zone in the 

upper crust, indicative of upthrusted gabbroic rocks; (2) predominantly low VP /VS

 ratios in the lower crust, indicative of an 

underlying quartz-rich basement; and (3) two low-VP /VS-ratio 

areas in the near surface associated with known quartz-rich 

rock sequences, one of sandstone and the other of granite 

and (or) granodiorite. We also tentatively identify a boundary between areas of low and slightly high VP /VS

 ratio at depth 

in the immediate vicinity of the Loma Prieta main-shock 

hypocenter. Our results demonstrate the diagnostic power of 

a three-dimensional VP /VS model in combination with other 

geologic and geophysical information.

INTRODUCTION

The Loma Prieta earthquake sequence of October 17, 1989, 

provides an outstanding opportunity for detailed study of the 

structure and setting of a large California earthquake and its 

aftershocks. These events occurred within the U.S. Geological 

Survey (USGS)’s central California seismic network (Calnet), 

and so precise locations for the earthquake sequence are available. In addition, an array of three-component digital recorders densely covered much of the area above the aftershock 

sequence, mainly from the main shock toward the northwest 

(Simpson and field team, 1989). These data have been made 

available to the seismological community as part of the Program for Array Seismic Studies of the Continental Lithosphere 

(PASSCAL) of the Incorporated Research Institutions for 

Seismology (IRIS).

The data collected during the PASSCAL deployment, along 

with supplemental data from Calnet and other portable-instrument 

deployments, are well suited for the application of three-dimensional seismic-imaging techniques (Thurber, 1983; Thurber and 

Aki, 1987). The broadly distributed seismicity along the San 

Andreas and other nearby fault zones recorded at the PASSCAL 

and USGS stations (fig. 1), along with data from numerous portable-array, quarry-blast, and seismic-refraction studies (Wesson and 

others, 1973; Mooney and Luetgert, 1982; Bakun and McLaren, 

1984; Mooney and Colburn, 1985), provides excellent spatial 

coverage for imaging. The three-dimensional P-wave-velocity 

structure was estimated by Eberhart-Phillips and others (1990) 

and Lees (1990). Owing to problems in the maintaining absolute 

time on the PASSCAL portable instruments (A.L. Lerner-Lam, 

oral commun., 1990) our present analysis utilizes S-minus-P (S−P) 

differential times only.

We used observed arrival-time differences between Pand S-wave first arrivals at the PASSCAL portable-array 

sites, augmented by some portable-instrument data from the 

University of California, Berkeley, and the USGS, to carry 

E143

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

THREE-DIMENSIONAL VARIATIONS IN VP / VS RATIO

IN THE LOMA PRIETA REGION1

1

University of Wisconsin, Madison, Geophysical and Polar Research 

Center contribution 536.

EARTHQUAKE OCCURRENCE
E144 EARTHQUAKE OCCURRENCE

out a single-step inversion for three-dimensional variations 

in the ratio of P- to S-wave velocity (VP /VS ) along the Loma 

Prieta rupture zone. The analysis incorporates the preliminary 

three-dimensional P-wave-velocity model of Eberhart-Phillips and others (1990) to determine the initial raypaths. The 

assumptions built into the inversion for VP /VS ratio (rather 

than directly for P- and S-wave velocities) do not necessarily limit this analysis to a single-step inversion; however, a 

rigorous iterative inversion should also ultimately include 

hypocenter and P-wave-velocity perturbations. This endeavor 

will be undertaken with a greatly enlarged data set, including 

calibration-shot data obtained in 1991 by the USGS (D.M. 

Eberhart-Phillips, oral commun., 1991).

S−P DATA

The process of estimating the S−P time for an event 

simply involves picking the P-wave-arrival time, t

P, and the 

S-wave-arrival time, tS, and then subtracting tP from tS. The 

identification of the P phase is generally easy, given adequate 

signal-to-noise ratio. Some workers have expressed concern 

?

?

?

?

?

122°15' 121°45'

37°15'

122°00'

0 5 10 km

SAN ANDREAS FAULT

CALAVERAS FAULT

SARGENT FAULT

PACIFIC

 OCEAN

 37°00'

ANDY

RIDR UCB5

UCB2

UCB3

UCB6

GOOD

WAWA

JIMV

KROC

WARB

SAR

SAOS

WVRD ANNS

DBWN

UCB4

WSWD

HUTR ZAYA

SMTH

OLSP

LOMP

RADF

RDWD

CRG LYON

KKO

HOLY

Figure 1.—Loma Prieta region, Calif., showing locations of study area (box), major faults, and epicenters (circles) of events used for 

VP /VS inversion. Triangles, seismic stations.
E145

regarding the effect of the antialias filter of the PASSCAL 

instruments on the onset of P waves. A careful examination of 

representative data with varying signal-to-noise ratio indicates 

that the antialias filter does not affect our data (J.A. Fowler, 

oral commun., 1991). Typically, a P-wave arrival can be 

picked within ±10 ms, whereas picking the S-wave arrival is 

almost never easy. Because an S-wave arrives in the P-wave 

coda, the coda interferes with the S-wave. Thus, unlike for 

the P wave, no sharp and clear onset may be evident for the 

S wave.

Picking the S wave generally is subject to additional 

uncertainty because of the possible presence of anisotropy 

and (or) the generation of S-to-P-converted waves. The Sto-P-converted wave arrives before the actual S wave and, if 

not identified correctly, would result in too early an S-wave 

pick, causing an apparent increase in the S-wave velocity. 

Propagation of the S wave in an anisotropic medium depends 

on several factors, and if the effects of anisotropy are not 

considered, this dependence could also result in a mispick. 

In an anisotropic medium, the polarization vectors of the S

wave are expected to be mutually orthogonal, with a degree 

of splitting dependent on the direction of propagation and the 

path length (Keith and Crampin, 1977; Babuska and Cara, 

1991). If the horizontal components are rotated in this fixed 

direction of polarization, the S wave will split into a fast and 

a slow S wave, indicating the presence of anisotropy. We and 

other researchers from the University of California, Santa 

Cruz (Susan Schwarz, oral commun., 1990) have examined 

the data for anisotropy and found some evidence for weak 

splitting (max 50 ms), which is typically within our estimated 

S-wave-picking uncertainty.

Distinguishing the S wave from an S-to-P-converted wave 

requires that the horizontal-component seismograms (northsouth and east-west) be rotated into the radial and transverse 

components. Examples of vertical-component, horizontalradial-component, and horizontal-transverse-component 

seismograms for events recorded at PASSCAL stations, along 

with particle-motion plots for selected time windows, are 

plotted in figures 2 through 9.

An event almost directly south of station WARB (fig. 1) is 

shown in fig. 2A. The P-wave arrival is sharp enough to be 

picked with an estimated uncertainty of ±20 ms. To identify 

the S-wave arrival, we use the particle-motion plot, even 

though a clear arrival at approximately 22.8 s could be picked 

as the S-wave arrival. First, we plot the horizontal particle 

motion of the direct P-wave to confirm that the azimuth of 

the P-wave arrival is generally along the station-epicenter 

line. Here, the particle-motion plot shows the azimuth of the 

P-wave arrival to be orthogonal to the station-epicenter line. 

Similar plots of P-wave arrivals at this station from other 

events also showed the same phenomenon, leading us to 

believe that a channel flip may be present at this station. Thus, 

for station WARB, the horizontal components were rotated 

90°, resulting in a radial P-wave first arrival (fig. 2B). A plot 

of a time window of the horizontal components within the 

expected range of the S-wave arrival is shown in figure 2C, 

and on the basis of its orthogonality to the P wave, this wave 

is identified as a direct S wave, arriving at approximately 

22.82 s. Thus, what appears here to be the S-wave arrival on 

the raw seismograms is, indeed, the S-wave arrival. Another 

example for station WARB shows a strong phase at 21.9 s 

near the expected S-wave-arrival time (fig. 3A). A plot of the 

horizontal components (fig. 3B) shows that the azimuth of 

the direct P wave is close to the station-epicenter line once 

it is rotated 90°. A plot of the horizontal components within 

the time window expected for the S-wave arrival (fig. 3C) 

clearly shows that the initial particle motion of the phase on 

the radial component has nearly the same arrival azimuth 

as the direct P wave, and so this phase is identified as an 

S-to-P conversion. This phase is followed by a phase that is 

orthogonal to the S-to-P-converted phase and is identified 

as a direct S wave. The time difference between these two 

phases is about 100 ms.

Examples for station ZAYA (fig. 1) are plotted in figures 4 

and 5. The seismograms in figure 4A show a sharp P wave 

on the vertical component and what appears to be a sharp S

wave. In figure 4B, the azimuth of the direct P-wave arrival is 

a bit off the station-epicenter line, presumably owing to lateral 

heterogeneity. Once again, a plot of the horizontal components 

within the time window expected for the S-wave arrival (fig. 

4C) shows that the initial particle motion is due to a phase 

with nearly the same arrival azimuth as the direct P wave. 

This phase is followed by a phase that is clearly orthogonal 

to the azimuth of the direct P wave. We interpret this later 

phase to be the direct S wave, and identify the earlier phase 

as an S-to-P conversion. The time difference between these 

two phases is only 36 ms, within the uncertainty we associate with our best S-wave picks. Another example for station 

ZAYA, from an event east-southeast of the station, again 

shows the P wave arriving off azimuth (fig. 5B). A plot of the 

horizontal components within the time window expected for 

the S-wave arrival (fig. 5C) shows a small phase identifiable 

as an S-to-P conversion.

More noteworthy examples of S-to-P-converted waves at 

stations WVRD and WAWA (fig. 1) are highlighted by the 

pairs of seismograms plotted in figures 6 and 7 and in figures 8 

and 9, respectively. One example from station WVRD shows a 

strong phase on the radial component, arriving at 13.35 s (fig. 

6A). A plot of the two horizontal components within a time 

window between 13.3 and 14 s (fig. 6C) clearly shows that the 

initial motion has the same azimuth as the direct P wave (fig. 

6B). The S wave does not arrive until 13.7 s on the transverse 

component. The time difference between the onset of these 

two phases is about 350 ms, within the range of S−P time 

perturbations that we expect to quantify. Another example 

from station WVRD again shows a strong phase on the radial 

component (fig. 7A), arriving at 23.95 s, whereas the direct S

wave apparently does not arrive until 24.25 s. Two examples 

from station WAWA for events to the north-northeast of the 

station each show a strong phase on the radial component, 

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION
E146 EARTHQUAKE OCCURRENCE

arriving at 22.45 s (fig. 8A) and 12.72 s (fig. 9A). The direct 

P-wave arrival is off azimuth and has the same orientation 

(figs. 8B, 9B). The first-arriving phase on the horizontal 

components for both events is close to the station-epicenter 

line (figs. 8C, 9C ). The direct S wave for both events arrives 

approximately 120 ms later, at approximately 22.57 and 12.84 

s, respectively (figs. 8A, 9A). If ample care is not taken, such 

possible mispicks would result in an apparent increase in Swave velocity and a decrease in VP /VS

 ratio.

We are confident that our picks were of the true S wave. 

We took great care to identify S-to-P-converted waves, as 

indicated above. We routinely used particle-motion plots to 

ascertain the characteristics of the S waves and reduce the 

uncertainty of each pick. The accuracy of a pick, however, 

still depends on the sharpness of the arrival. We assigned 

quality factors to our picks on the basis of the uncertainty of 

the arrival time. If the onset of the S wave had an estimated 

uncertainty of <50 ms, we assigned the pick a quality factor 

of 0. Similarly, for uncertainties of ±75, ±100, ±125, and ±200 

ms in arrival time, we assigned a quality factor of 1, 2, 3, and 

4, respectively. S−P times were weighted by a factor of (1/2)n

in our inversions, where n is the quality factor, except that data 

with a quality factor of 4 were excluded from the inversion. 

The S−P data determined in this study were submitted to the 

Loma Prieta Data Archive Project of the National Information 

Service for Earthquake Engineering.

INVERSION METHOD

Our procedure of inverting for VP /VS ratios is similar in concept to that of Walck (1988), but with one major difference. 

Vertical

Transverse

Radial

5

5

5

5

TIME, IN SECONDS

A

B C

Figure 2.—Seismograph records from station WARB (fig. 1). A, P- and S-wave arrivals on vertical and two horizontal components from an 

event at lat 37°10.13' N., long 122°2.51' W., and 12.3-km depth. B, Particle-motion plot of direct P wave for a time window between 20.50 

and 20.85 s. C, Particle-motion plot of direct S wave for a time window between 22.70 and 23.05 s.
E147

The arrival time tij for a P or S wave at station j from event 

i is given by

 

 (1)

where t0 is the event origin time and u is the P- or S-wave 

slowness (1/V). If the VP /VS ratio is treated as a constant initially, 

then the raypaths are identical for both P and S waves, and 

we can express the S−P time difference dtij as

 

 

 (2)

Note the removal of dependence on event origin time, but the 

continued implicit dependence on event location as one end 

point of the path integral.

We can now develop a linearized inversion procedure for 

determining variations in VP /VS ratio. We begin by adopting 

a three-dimensional P-wave-velocity structure defined on 

a three-dimensional grid of nodes (Eberhart-Phillips and 

others, 1990), assuming an initially constant VP /VS ratio. We 

made a “reasonable” choice for an initial VP /VS ratio of 1.75 

on the basis of the results of preliminary inversions and an 

analysis of Wadati diagrams. Observed S−P time differences 

dtij were determined for a substantial set of events in the 

PASSCAL data set that have also been accurately located 

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION

Vertical

Transverse

Radial

A

B C

5

5

5

5

TIME, IN SECONDS

Figure 3.—Seismograph records from station WARB (fig. 1). A, P- and S-wave arrivals on vertical and two horizontal components from an 

event at lat 37°6.92' N., long 122°0.01' W., and 14.8-km depth. B, Particle-motion plot of direct P wave for a time window between 20.25 

and 20.50 s. Note that azimuth of P wave is close to station-epicenter line. C, Radial-transverse particle-motion plot of direct S wave for 

a time window between 22.90 and 23.15 s. Note that initial particle motion is along nearly the same azimuth as direct P wave. Near 23.02 

s, direct S wave arrives with a particle motion orthogonal to azimuth of initial phase, which is identified as S-to-P conversion.

tij = t0 + uds

path ∫ ,

dt

V

V

V

ds ij

P

S

P

= �

�

�

�

�

�

�

�

�

�

�

�

�

� 1

path

.
E148 EARTHQUAKE OCCURRENCE

by the USGS, using both “two dimensional” (Dietz and 

Ellsworth, 1990) and “three dimensional” velocity models 

(Eberhart-Phillips and others, 1990). We used both of these 

sets of velocity models and hypocenter locations to assess the 

sensitivity of the VP /VS inversion to assumed event locations 

and an initial velocity model. We calculated raypaths and 

expected S−P time differences dtij* by using the three (or 

two)-dimensional P-wave-velocity model and a constant VP /

VS ratio, according to the pseudo-bending three-dimensional 

ray-tracing method of Um and Thurber (1987). This step is 

distinct from the approach of Walck (1988), who did not use 

the available three-dimensional P-wave-velocity model of 

Walck and Clayton (1987) to trace rays for the backprojection of S−P residuals. We then related the S−P time misfit 

Dtij=dtij−dtij* to perturbations in VP /VS ratio at the nodes of 

the three-dimensional grid, and inverted for VP /VS ratio in a 

single step, keeping the P-wave velocities and hypocenter 

locations fixed. Note that we can calculate revised S-wave 

traveltimes by first applying (VP /VS) corrections to determine 

a grid of S-wave velocities for ray tracing and then redetermining the expected S−P times, thus directly assessing the 

improvement to the data fit, along with the usual measures 

of solution uncertainty and resolution.

This strategy for determining three-dimensional variations 

in VP /VS ratio is analogous to single-step inversion of earthquake P-wave-arrival-time residuals keeping the event locaA

B C

5

5

5

5

TIME, IN SECONDS

Vertical

Transverse

Radial

Figure 4.—Seismograph records from station ZAYA (fig. 1). A, P- and S-wave arrivals on vertical, radial, and transverse components from an 

event at lat 37°12.57' N., long 122°5.15' W., and 9.3-km depth. B, Radial-transverse particle-motion plot of direct P wave for a time window 

between 11.25 and 11.50 s. Note that azimuth of P wave is not along station-epicenter line. C, Radial-transverse particle-motion plot of direct 

S wave for a time window between 13.60 and 13.85 s. Note that initial particle motion is along same azimuth as direct P wave. Around 13.68 

s, direct S wave arrives with a particle motion orthogonal to azimuth of initial phase, which is identified as S-to-P conversion.
E149

tions fixed. We invoke Fermat’s principle to justify the use 

of the (three dimensional) P-wave path to assess the spatial 

variation in VP /VS ratio, so that we neglect the variation of the 

“final” (S wave) path from the “initial” (P wave) path (which 

would be nil for a truly constant VP /VS ratio). These raypath 

differences should have only a second-order effect on our 

estimation of the variation in VP /VS ratio. Indeed, if spatial 

variations in VP /VS ratio are modest (say, max 5 percent), this 

assumption is quite valid. However, the complex geology of 

the Loma Prieta region (fig. 1), as well as the level of lateral

P-wave-velocity variations determined by Eberhart-Phillips 

and others (1990) (>20 percent in places), leads us to suspect 

that the variation in VP /VS ratio may be substantial and an 

iterative solution will eventually be required. Thus, we view 

this single-step inversion as the first stage in our efforts, to 

be followed in the future by an iterative determination of Pwave-velocity structure, VP /VS

 structure, and event locations 

from P-wave-arrival and S−P times.

If we are interested in using variations in VP /VS ratio (or, 

equivalently, Poisson’s ratio) to infer geologic structure and 

mechanical properties (including, for example, pore pressure), then the approach of inverting directly for variations 

in VP /VS ratio has certain significant advantages. It is well 

known that a direct comparison of three-dimensional P- and 

S-wave-velocity models to infer VP /VS ratios is hampered by 

the tendency for S-wave-velocity models to be more poorly 

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION

A

B C

5

5

5 5

TIME, IN SECONDS

Vertical

Transverse

Radial

Figure 5.—Seismograph records from station ZAYA (fig. 1). A, P- and S-wave arrivals on vertical, radial, and transverse components from an 

event at lat 37°9.19' N., long 122°4.06' W., and 11.1-km depth. B, Radial-transverse particle-motion plot of direct P wave for a time window 

between 10.25 and 10.50 s. Note that azimuth of P wave is not along station-epicenter line. C, Radial-transverse particle-motion plot of direct 

S wave for a time window between 12.35 and 12.6 s. Note that initial particle motion is along same azimuth as direct P wave. Around 12.5 

s, direct S wave arrives with a particle motion orthogonal to azimuth of initial phase, which is identified as S-to-P conversion.
E150 EARTHQUAKE OCCURRENCE

resolved than P-wave-velocity models (Eberhart-Phillips, 

1989). Thus, variations in the P-wave-velocity model tend to 

be more extreme, whereas variations in the S-wave-velocity 

model tend to be damped, resembling more closely the initial model. As a result, inferred VP /VS

 ratios might not reflect 

constraints that the data do (or do not) place on the variations 

in VP /VS ratio, but rather would tend to be controlled simply 

by the perturbation in P-wave velocity. If, instead, we invert 

directly for three-dimensional perturbations in VP /VS ratio, the 

solution will produce a relatively faithful representation of 

significant variations in VP /VS ratio where the data require it 

(and it is resolved), and elsewhere VP /VS ratios will be unperturbed, owing either to absence of heterogeneity or absence 

of resolution.

INVERSION TESTS

We made some tests to assess the need for a three-dimensional model to satisfy the data, and to examine the sensitivity of our results to the initial model (P-wave-velocity 

structure and corresponding hypocenters). Inversions for a 

one-dimensional (that is, vertically varying) model resulted 

in a structure with high near-surface and low midcrustal 

A

B C

5

5

5

5

TIME, IN SECONDS

Vertical

Transverse

Radial

Figure 6.—Seismograph records from station WVRD (fig. 1). A, P- and S-wave arrivals on vertical, radial, and transverse components from an 

event at lat 37°2.32' N., long 121°49.02' W., and 4.0-km depth. B, Radial-transverse particle-motion plot of direct P wave for a time window 

between 10.40 and 10.80 s. C, Radial-transverse particle-motion plot of direct S wave for a time window between 13.35 and 14.00 s. Note 

that initial particle motion is along nearly the same azimuth as direct P wave. Near 13.70 s, direct S wave arrives with a particle-motion 

orthogonal to azimuth of initial phase, which is identified as S-to-P conversion.
E151

VP /VS ratios (table 1). Including station corrections in the 

one-dimensional inversion resulted in a sharp reduction in 

the perturbation of near-surface VP /VS ratios and insignificant 

changes elsewhere. In both one-dimensional cases, variance reduction was less than 30 percent. Variance reduction 

should be significantly greater for a three-dimensional model 

to be justified.

We also examined the sensitivity of the inversion to the 

initial hypocenters and P-wave-velocity model. Preliminary 

three-dimensional VP /VS models were determined initially 

with a modest-size data set (31 earthquakes, 269 S−P times), 

using both the “three-dimensional initial” (Eberhart-Phillips 

and others, 1990) and “two-dimensional initial” (Dietz and 

Ellsworth, 1990) velocity models and corresponding hypocenters. The pattern of features in the two solutions were 

quite similar, but the “three-dimensional initial” velocity 

model yielded a slightly superior data fit (root-mean-square 

[rms] S−P time residual of 0.13 versus 0.14 s, from initial 

values of 0.20 and 0.21 s, respectively), with a lower rms 

amplitude of model perturbations (rms (VP /VS) of 0.09 

versus 0.11) relative to the “two-dimensional initial” velocity model. Thus, we adopt the “three-dimensional initial” 

velocity model as our preferred initial model and confine our 

discussion below to results derived from that model.

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION

A

B C

5

5 5

5

TIME, IN SECONDS

Vertical

Transverse

Radial

Figure 7.—Seismograph records from station WVRD (fig. 1). A, P- and S-wave arrivals on vertical, radial, and transverse components from 

an event at lat 36°57.35' N., long 121°42.93' W., and 13.4-km depth. B, Radial-transverse particle-motion plot of direct P wave for a time 

window between 19.45 and 19.65 s. C, Radial-transverse particle-motion plot of direct S wave for a time window between 23.95 and 24.35 

s. Note that initial particle-motion is along nearly the same azimuth as direct P wave. Around 24.25 s, direct S wave arrives with a particle 

motion orthogonal to azimuth of initial phase, which is identified as S-to-P conversion.
E152 EARTHQUAKE OCCURRENCE

Systematic errors in the initial P-wave-velocity model 

may still be propagated into biased perturbations in VP /VS

ratio. The comparison above of two different sets of initial 

velocity models and their corresponding hypocenters suggests that our inversion is reasonably robust. The study by 

Schwartz and Nelson (1991) also indicates good stability 

of hypocenters, whether they are determined with P-wavearrival times only (Calnet data) or S−P arrival times only 

(PASSCAL data), and whether a two-dimensional (Dietz and 

Ellsworth, 1990) or three-dimensional (Eberhart-Phillips 

and others, 1990) velocity model is used (both assuming a 

constant VP /VS ratio).

RESULTS OF

THREE-DIMENSIONAL INVERSION

The results of the three-dimensional inversion from our 

complete S−P data set (128 earthquakes, 1,054 S−P times) 

are plotted in figure 10. Estimated VP /VS ratios range from 

1.57 to 2.00, corresponding to a range of Poisson’s ratios 

from 0.16 to 0.35. Note that these variations are significant 

relative to the uncertainty in VP /VS ratio of about 3 percent 

estimated from the inversion, corresponding to an uncertainty in VP /VS

 ratio about 0.05. This inversion used an initial 

VP /VS ratio of 1.75; using an initial VP /VS ratio of 1.80 resulted 

A

B C

5

5 5

5

TIME, IN SECONDS

Vertical

Transverse

Radial

Figure 8.—Seismograph records from station WAWA (fig. 1). A, P- and S-wave arrivals on vertical, radial, and transverse components from 

an event at lat 37°10.95' N., 121°58.56' W., and 2.2-km depth. B, Radial-transverse particle-motion plot of direct P wave for a time window 

between 20.20 and 20.40 s. C, Radial-transverse particle-motion plot of direct S wave for a time window between 22.45 and 22.65 s. Note 

that initial particle motion is along nearly the same azimuth as direct P wave. Around 22.57 s, direct S wave arrives with a particle motion 

orthogonal to azimuth of initial phase, which is identified as S-to-P conversion.
E153

in a solution that differed by no more than 0.02 from the 

model shown. We also performed a three-dimensional inversion with station corrections. As expected, the resulting VP /VS

model differed significantly only in the near-surface grid 

layer, where most perturbations were cut approximately in 

half. Surprisingly, despite the greater number of parameters, 

the variance reduction of less than 60 percent was somewhat 

less than that for the three-dimensional inversion without 

station corrections (64 percent). Thus, we have adopted 

the three-dimensional model without station corrections 

as our best model.

In the shallow subsurface (layer 2, fig. 10B), VP /VS ratios 

are generally high (min 1.80), especially along the San 

Andreas fault zone. This result is not surprising because 

near-surface rocks typically have high VP /VS ratios (Nicholson 

and Simpson, 1985), owing in part to extensive microcracks 

(O’Connell and Budiansky, 1974). Two areas are exceptions: 

just west of the San Andreas fault, near stations OLSP (x=−5 

km, y=0 km; VP /VS =1.57) and WAWA (x=−15 km, y=−10 

km; VP /VS =1.67); and just east of the Sargent fault where it 

merges with the San Andreas fault, near stations WVRD and 

SAOS (x=2 km, y=−10 km; VP/VS=1.68) (see fig. 1 for locations). At shallow depths (layer 3, fig. 10C), VP/VS

 ratios are 

persistently high, though somewhat varying, along the San 

Andreas fault. A downward continuation of anomalously low 

VP /VS ratios also occurs at stations OLSP and WAWA.

A

B

C

5

5

5

5

TIME, IN SECONDS

Vertical

Transverse

Radial

Figure 9.—Seismograph records from station WAWA (fig. 1). A, P- and S-wave arrivals on vertical, radial, and transverse components from 

an event at lat 37°9.46' N., long 122°0.20' W., and 10.10-km depth. B, Radial-transverse particle-motion plot of direct P wave for a time 

window between 10.70 and 10.90 s. C, Radial-transverse particle-motion plot of direct S wave for a time window between 12.7 and 12.90 

s. Note that initial particle motion is along nearly the same azimuth as direct P wave. Around 12.84 s, direct S wave arrives with a particle 

motion orthogonal to azimuth of initial phase, which is identified as S-to-P conversion.

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION
E154 EARTHQUAKE OCCURRENCE

Layer 4 (fig. 10D) displays only one feature that we view as 

significant. An area of moderately high VP /VS ratio (1.80–1.87) 

is present along and just southwest of most of the San Andreas 

fault (x=−5 to 2 km, y=−20 to 0 km). Good model resolution 

becomes fairly restricted to the immediate vicinity of the 

seismically active fault-zone area in layer 5 (fig. 10E), and so 

we have little information on the VP /VS structure northeast of 

the San Andreas fault. However, we do find a pattern of low 

VP /VS ratios along and southwest of most of the San Andreas 

fault, in sharp contrast to the high VP /VS ratios along the fault 

at shallower depths. In layer 6 (fig. 10F), good resolution is 

quite limited; however, a strong VP /VS -ratio low occurs in the 

area just northwest of the main-shock epicenter (x=−5 km, 

y=−10 to 0 km; VP /VS =1.55−1.69).

Rows of the resolution matrix for selected nodes of our 

model are listed in tables 3 through 6. The patterns are quite 

typical: resolution is highest for the diagonal element, with 

two or three neighboring nodes having resolutions a factor 

of 2 to 5 smaller and the other nodes having resolutions a 

factor of 8 to 10 smaller or even less. We consider the model 

resolution to be quite satisfactory.

DISCUSSION

Our results indicate significant variations in VP /VS ratio 

laterally and with depth throughout the study area (fig. 1). 

We note that the VP /VS ratios in layer 2 (0-km depth, fig. 10B), 

layer 3 (3-km depth, fig. 10C), layer 4 (7-km depth, fig. 10D), 

layer 5 (11-km depth, fig. 10E), and layer 6 (16-km depth, 

fig. 10F) vary by 27, 15, 12, 7, and 17 percent, respectively. 

Previous studies have shown that VP /VS ratios are sensitive 

to pore-fluid pressure and crack or pore geometry, and are 

influenced by the distribution of crack aspect ratios (Tatham, 

1982). VP /VS ratios in crystalline rocks, which are known to 

have significant porosity, are strongly affected by the degree 

of flatness of the cracks because S-wave velocity is more sensitive to aspect ratio than is P-wave velocity (Tatham, 1982). 

In sedimentary rocks, composition more than crack geometry 

may be the dominant factor in controlling the VP /VS ratio 

 0 1.95 0.88 5

 3 1.77 .96 3

 7 1.71 .97 3

 11 1.71 .97 3

 16 1.62 .91 4

 24 1.79 .14 5

VP /VS Diagonal Uncertainty

 ratio resolution (percent)

Node

depth

(km)

Table 1.— One-dimensional Vp / Vs model without station corrections

0–3 4

3–7 22

7–11 44

11–16 59

16–24 11

Number

 of 

events

Depth

range

(km)

Table 2.—Depth distribution of 

aftershocks used for inversion.

(Wilkens and others, 1984). VP /VS ratios are also influenced 

by confining pressure, pore pressure, pore-fluid saturation, 

and to a lesser degree, temperature. VP /VS ratios decrease with 

depth as microcracks close under increasing confining pressures (Nur and Simmons, 1969; O’Connell and Budiansky, 

1974) because S-wave velocity increases faster than P-wave 

velocity, owing to closure of microcracks at depth. Studies 

suggest that most microcracks in crustal rocks close when 

confining pressures reach 1 to 2 kbars (at 3.5–7-km depth) 

and that further increases in pressure with depth have little 

effect on the VP /VS ratio (Wilkens and others, 1984). Thus, 

below 6- or 7-km depth, the composition and mineralogy of 

the rock strongly influences the VP /VS ratio.

For shallow crustal rocks, composition can still significantly 

affect inplace VP /VS ratios. In particular, laboratory studies show 

that free quartz has a low Poisson’s ratio (0.077; Simmons 

and Wang, 1971)—that is, a VP /VS ratio of 1.48—and that even 

the bulk Poisson’s ratio of a rock containing 20 to 40 volume 

percent free quartz is relatively low (Ukawa and Fukao, 1981). 

Kern and Richter (1981) and Christensen and Wepfer (1989) 

also concluded that quartz-rich rocks have low Poisson’s ratios 

but that feldspar-rich rocks have high Poisson’s ratios.

Clearly, VP /VS ratios depend on several factors, and the 

variation in VP /VS ratio in rocks is due to one or more these 

factors. The variations in VP /VS ratio for different rock types 

overlap, and so it is difficult to assign a unique VP /VS ratio 

to a given rock type. Thus, additional information is also 

necessary, such as independently estimated P-wave-velocity structure, as well as surface and subsurface geologic 

and geophysical information. The independently estimated 

P-wave-velocity structure in the study area (fig. 1; EberhartPhillips and others, 1990) also shows considerable variations 

in P-wave velocity laterally and with depth. We note that the 

P-wave velocities in layers 2, 3, 4, 5, and 6 of the model of 

Eberhart-Phillips and others (1990) vary by 57, 54, 12, 6, 

and 4 percent, respectively. This observation clearly shows 

that P-wave velocity varies significantly above 7-km depth. 

P-wave velocity, like VP /VS ratio, is sensitive to several factors, such as porosity, aspect ratio, pore-fluid composition, 

and fluid saturation. Thus, any interpretation of the geologic 

structure based only on P-wave velocity can be misleading 
E155

above 7-km depth. Below 7-km depth, however, lateral 

variations in P-wave velocity are significantly smaller (<12 

percent) in the study area, probably because of closure of 

microcracks due to high confining pressures, resulting in a 

reduced influence of porosity and pore fluids. Thus, for 7-km 

depth and greater, P-wave velocity and VP /VS ratio together 

could provide some information about the mineralogy of the 

rock, whereas P-wave velocity alone might not.

The significance of VP /VS ratio as a diagnostic tool over P-wave 

velocity for geologic interpretation can be clearly noted below 

10-km depth. In layer 5 (fig. 10E), the estimated P-wave-velocity structure of Eberhart-Phillips and others (1990) shows a 

mere 6-percent lateral variation in P-wave velocity, thus not 

providing much evidence about rock composition. In contrast, 

the VP /VS ratios clearly suggest that rocks with different mineralogies could be present at this depth. Studies by Molotova 

and Vassil’ev (1960), Press (1966), and Christensen (1989), 

however, indicate enough nonuniqueness in VP /VS ratios even 

for high P-wave velocities (below 7-km depth) that geologic 

interpretations based only on seismic-velocity information can 

be misleading. Thus, when making geologic interpretations in a 

region, it is crucial to incorporate information on (1) the density 

of specific rock types and their lateral variations, as inferred 

from gravity data; (2) the depth and position of rocks that 

produce anomalous magnetic fields, as inferred from magnetic 

data; and (3) surface geology. In combination with information 

from other geophysical and geologic investigations, variations 

in VP /VS ratio provide powerful diagnostic evidence regarding 

the geologic composition and structure of a region.

Magnetic and gravity anomalies in the study area (fig. 1), based 

on the maps by Brabb and Hanna (1981), Chuchel and Jachens 

(1990), and Jachens and Griscom (this chapter), are mapped 

in figures 11 and 12, respectively. These maps of geophysical 

anomalies are invaluable in our geologic interpretation.

Two conspicuous magnetic anomalies occur southwest of 

the San Andreas fault in the study area (fig. 1): near Boulder 

Creek and Corralitos (fig. 11). The Boulder Creek anomaly 

is a northwest-southeast-trending, elongate magnetic high 

with an amplitude of 120 γ between the Zayante and Butano 

faults, where Cenozoic sedimentary rocks are present at the 

surface. Magnetic-field gradients indicate that the source of 

the Boulder Creek anomaly is below about 1.5-km depth and 

may extend significantly deeper (Hanna and others, 1972). 

This magnetic high coincides with a region of average to 

high VP /VS ratios in layers 2, 3, and 4 (figs. 10B, 10C, and 

10D, respectively) and is sandwiched between two northsouth-striking gravity highs to the east and west (fig. 12). 

The Corralitos anomaly is somewhat more complex: two 

magnetic highs flank a magnetic low between the Zayante 

and San Andreas faults and a paired magnetic high and low 

southwest of the Zayante fault.

The Boulder Creek anomaly and the part of the Corralitos 

anomaly between the Zayante and the San Andreas faults 

resemble the Logan anomaly situated southeast of the study 

area (fig. 1). The Logan anomaly is associated with exposed 

quartz gabbro and anorthosite gabbro (Ross, 1970), which 

have relatively high densities (Clark and Rietman, 1973) and 

a strong normal magnetization (Brabb and Hanna, 1981). The 

map by Hanna and others (1981) also shows an outcrop of 

gabbro about 10 km southwest of the center of the Boulder 

Creek anomaly. A gabbroic source rock for the Boulder Creek 

and Corralitos anomalies that lies between the Zayante and 

San Andreas faults would also be consistent with the high 

VP /VS ratio of 1.81 to 2.0 in our model (VP /VS =1.84–1.96 for 

gabbro; Molotova and Vassil’ev, 1960; Press, 1966). A gravity low occurs over the magnetic high between the Zayante 

and San Andreas faults. The absence of a gravity high at the 

Corralitos anomaly may be due to the presence of a thick 

section of clastic marine sedimentary rocks at the surface 

(McLaughlin and others, 1988).

The magnetic low within the Corralitos anomaly southwest of the Zayante fault coincides with the regions of low 

VP /VS ratio in layers 2, 3, and 4 (figs. 10B, 10C, and 10D, 

respectively) and is close (2 km southeast) to an outcrop of 

granodiorite. Granodiorite also crops out in the Bald Mountains west-northwest of the magnetic low (Hanna and others, 

1981). Thus, these granitoids are also present southwest of 

the Zayante fault. We note that the trace of the Zayante fault 

separates the low-VP /VS-ratio area to the southwest from the 

high-VP /VS-ratio area to the northeast in layer 2 (0-km depth, 

fig. 10B). The projection of the fault trace downward to layers 

3 and 4 (figs. 10C and 10D, respectively) also shows a clear 

separation between the areas of low and high VP /VS ratio, suggesting that the Zayante fault extends downward to at least 

7-km depth. A correlation of stratigraphic units in this area 

supports a downward extension of the Zayante fault to at least 

6-km depth (R.J. McLaughlin, oral commun., 1992).

Granite and granodiorite commonly contain 20 to 40 volume 

percent quartz, a diamagnetic mineral, and so give rise to negative magnetic anomalies. The presence of quartz also helps 

to lower the VP /VS ratio of a rock. Laboratory studies on the 

granodiorite of Lakeside, Calif., yielded a VP /VS ratio of 1.54 

(Press, 1966). Similarly, laboratory studies on the granodiorite 

of Weston, Mass., also yielded a VP /VS ratio of 1.54 (Birch and 

Bancroft, 1940); a study of the mineral composition of this 

granodiorite showed that it contained 33.5 volume percent 

quartz. The P-wave velocities in these two granodiorites were 

4.88 and 4.78 km/s, respectively. The three-dimensional Pwave-velocity models of Eberhart-Phillips and others (1990) 

and Foxall and others (1993) at a depth (below mean sea level) 

of 3 km show P-wave velocities of 4.5 to 5.0 km/s southwest 

of the Zayante fault, similar to the P-wave velocities in granodiorite at very shallow depths, as cited above and elsewhere 

(Press, 1966). Thus, we interpret the anomalously low VP /VS

ratios in layers 2, 3, and 4, (figs. 10B, 10C, and 10D, respectively), west of the Zayante fault, to be due to the presence of 

granodiorite and rocks of similar composition.

A second area of anomalously low VP /VS ratio is present in 

layer 2 (0-km depth, fig. 10B), just northeast of the Sargent 

fault. The area is underlain by a zone of average to high VP /VS

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION
E156 EARTHQUAKE OCCURRENCE

-30 -20 -10 0 10 20

-30

-20

-10

0

 10

SAN ANDREAS

 FAULT

SARGENT

 FAULT

ZAYANTE

 FAULT

A

x

y

-- 00 12 02

-05 14 08 06

-03 25 -16

-- -23 07 -07

-- 10 07 00

H

L

H

15

-09

15 -09

07

10

-30 -20 -10 0 10 20

-30

-20

-10

0

 10 B

L

x

y

-30 -20 -10 0 10 20

-30

-20

-10

0

 10

-- 00 08 08

03 08 10 01

-01 00 08

-01 -13 04 09

11 04

H

H

-04

00

01

-01

07

-- 11 00 C

x

y

-03

-- 00 01 --

03 09 01 00

-02 02 01-01 00

 04 07 04

-- 08

H

H

-01

00

-30 -20 -10 0 10 20

-30

-20

-10

0

 10 03 01

 03

00

D

x

y

-- --

01 -06-14 --

-02 -13-06 00

-01 -10 -01

-- 00 -0102

L

-01 -03 --

00

-02-01

00

-30 -20 -10 0 10 20

-30

-20

-10

0

 10

--

E

x

y

-- -- -- -- --

-- -02 -01 -- --

-- -14 -06 -- --

-- -08 02 -- --

-- 05 -- --

L

 05

-30 -20 -10 0 10 20

-30

-20

-10

0

 10

H

F

x

y
E157

ratio in layer 3 (fig. 10C). This low-VP /VS-ratio region coincides with a magnetic low and a moderate gravity high. These 

observations are consistent with the geologic interpretation 

by R.J. McLaughlin (oral commun., 1990) that this area has 

a 2- to 3-km-thick near-surface layer of quartz-rich sandstone 

underlain by underthrust ophiolitic basement. Studies by 

Christensen (1989) show that the Poisson’s ratio of the Mount 

Simon and St. Simon sandstones of northern Illinois can be 

as low as 0.185 and 0.142, respectively, under moderate (1–2 

kbars) confining pressures, corresponding to VP /VS ratios of 

1.61 and 1.55. Sample mineralogic descriptions of these 

two rocks show that the Mount Simon sandstone contains 

84 volume percent quartz and the St. Simon sandstone 91 

volume percent quartz. Both of these rocks have a porosity of 8 to 9 percent and were from a depth of only 670 m. 

Here, the high quartz content of the rocks seems to control 

0.0000 -0.0078 -0.0124 0.0123 -0.0037

-.0098 -.0500 .0414 -.0468 -.0266

-.0128 .0373 .5108 -.0018 -.0309

.0000 -.0648 .1139 .0068 -.0265

.0000 -.0054 -.0211 -.0251 .0000

.0000 .0000 -.0113 -.0002 .0000

-.0040 -.0081 -.0065 -.0052 .0453

-.0145 -.0378 .1685 .0248 .0038

.0000 -.0144 .0296 .0107 .0008

.0000 -.0054 -.0164 -.0276 .0000

.0000 .0000 -.0052 -.0011 .0000

.0027 .0186 .0192 -.0029 .0000

-.0047 .0268 .0204 -.0208 -.0033

.0000 .0251 .0085 .0237 .0039

.0000 .0070 .0015 .0016 .0000

.0000 .0000 .0006 .0000 .0000

.0000 .0049 .0101 -.0002 .0000

-.0017 .0131 -.0016 .0275 .0000

.0000 -.0019 .0013 -.0058 .0000

.0000 .0017 -.0001 .0021 .0000

.0000 .0000 .0000 .0000 .0000

.0000 .0002 .0007 .0000 .0000

.0000 .0005 -.0050 .0000 .0000

.0000 -.0005 .0048 .0000 .0000

.0000 .0023 .0060 .0000 .0000

Table 3.—Resolution matrix for layer 2 node of three-dimensional inversion model.

[Diagonal element is listed in bold]

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION

Figure 10.—Results of three-dimensional inversion. A, Schematic map of study area in Loma Prieta region, Calif. (fig. 1), showing locations 

of PASSCAL, University of California, Berkeley, and GEOS stations (triangles) and mapped faults for reference. y-axis of grid is rotated 

41° west of north, following three-dimensional model of Eberhart-Phillips and others (1990). B–F, Final VP /VS ratios for model nodes in grid 

layers 2 through 6 at depths of 0, 3, 7, 11, and 16 km, respectively. Grid layer at 24-km depth was totally unresolved, and so no results are 

plotted for that layer. H, high; L, low. x- and y-axes in kilometers. Resolution of VP /VS ratio: large numbers, >0.45; medium-size numbers, 

0.15–0.45; small numbers, <0.15. Contour interval, 0.10; dashed where resolution is poor. Lateral node spacing ranges from 3 to 15 km 

across strike (northeast-southwest) and from 10 to 15 km along strike (northwest-southeast) of the San Andreas fault. Variance reduction for 

single-step inversion, 64 percent, with a root-mean-square perturbation of 0.09 and an average estimated error of 3 percent for moderately 

well resolved nodes (diagonal element >0.15). Depth distribution of events is listed in table 2.
E158 EARTHQUAKE OCCURRENCE

the Poisson’s ratio. Similarly, the high quartz content of the 

sandstone in the study area (fig. 1) appears to dominate the 

VP /VS ratios.

The VP /VS ratios in layer 2 (0-km depth, fig. 10B) range from 

1.80 to 2.0. A correlation with the surface-geology map by 

Brabb and Hanna (1981) shows that the VP /VS ratio of 2.00 at 

x=−5 km, y=−10 km, occurs about 4 km southwest of outcrops 

of diabase and gabbro in contact with outcrops of serpentinite 

and Franciscan rocks. Similarly, the VP /VS ratios of 1.95 at x=−5 

km, y=−20 km, of 1.91 at x=0 km, y=−10 km, and of 1.82 at 

x=0 km, y=−20 km, occur very close to (within 5 km of) outcrops of serpentinite and diabase. The outcrops of serpentinite 

and diabase near the nodes at x=−5 km, y=−10 km, and at x=0 

km, y=−10 km, are surrounded by a thick section of Tertiary 

marine sedimentary rocks. The gravity low visible in figure 

12, which extends from y=10 km to y=−10 km near x=−5 km, 

marks approximately the boundary of this sedimentary section. The outcrops of serpentinite and diabase near the nodes 

at x=0 km, y=−20 km, and at x=5 km, y=−20 km, are also surrounded by sedimentary rocks with P-wave velocities of 3.5 to 

4.5 km/s (Eberhart-Phillips and others, 1991). The VP /VS ratios 

of diabase from eastern Siberia and from Centerville, Va., for 

example, range from 1.8 to 2.3. Similarly, the VP /VS ratios of 

serpentinite range from 2.0 to 2.1 (Press, 1966). Samples of 

serpentinite collected near Loma Prieta by N.I. Christensen 

(written commun., 1991) have a VP /VS ratio of 1.93 to 1.95 over 

a depth range of 0 to 3 km. Both diabase and serpentinite are 

rich in ferromagnetic minerals, and so a magnetic contribution 

to the Boulder Creek anomaly due to these outcrops and their 

subsurface extensions (if they exist) is possible. The range in 

VP /VS ratio of 1.8 to 2.0 for this area is also consistent with the 

VP /VS ratios of sedimentary rocks.

Comparing our model with those of Eberhart-Phillips and 

others (1990) and Foxall and others (1993) for layer 2 (0-

km depth, fig. 10B), we note that a region of lower P-wave 

velocities (∼4.0–4.5 km/s) in their models coincides with the 

high VP /VS ratios between y=10 km and y=−10 km near x=0 

km in our model. This lower-P-wave-velocity region corresponds to the thick section of Tertiary marine sedimentary 

rocks. The P-wave velocities in a serpentinite and a diabase at 

shallow (<1 km) depths, however, were determined to range 

from 5.5 to 5.9 km/s and from 6.2 to 6.8 km/s, respectively 

0.0000 0.0000 -0.0037 -0.0010 -0.0011

-.0070 -.0239 .0053 -.0041 -.0023

-.0162 .0392 -.0144 -.0039 -.0004

.0000 .2741 -.0215 -.0061 .0038

.0000 -.0029 -.0068 -.0009 .0000

.0000 .0000 -.0048 -.0012 .0000

.0108 -.0049 .0103 .0008 .0009

-.0548 .0130 - .0182 .0038 .0020

.0000 .5955 .0428 -.0178 .0038

.0000 -.0053 -.0125 -.0071 .0000

.0000 .0000 -.0012 .0002 .0000

.0060 .0056 .0005 -.0026 .0000

-.0144 -.0218 -.0038 .0069 .0003

.0000 .1529 .0713 .0004 -.0008

.0000 .0704 .0239 -.0014 .0000

.0000 . 0000 .0003 .0000 .0000

.0000 .0064 -.0025 -.0003 .0000

.0053 .0301 .0059 .0060 .0000

.0000 -.0330 -.0399 -.0156 .0000

.0000 .0143 .0177 -.0049 .0000

.0000 .0000 .0000 .0000 .0000

.0000 -.0054 .0005 .0000 .0000

.0000 -.0146 -.0054 .0000 .0000

.0000 .0268 .0067 .0000 .0000

.0000 .0035 .0170 .0000 .0000

Table 4.—Resolution matrix for layer 3 node of three-dimensional inversion model.

[Diagonal element is listed in bold]
E159

(Birch, 1960). The P-wave-velocity models of Eberhart-Phillips and others (1990) and Foxall and others (1993) do not 

show these variations in P-wave velocity in the near-surface 

rocks. Thus, the sedimentary section is apparently contributing significantly more to the high VP /VS

 ratios in layer 2, and 

so the contribution from the outcrops of serpentinite and 

diabase may be mostly confined to affecting the magnetic 

anomalies only. Notwithstanding, the low-VP /VS-ratio areas at 

x=2 km, y=−10 km, and at x=7 km, y=−10 km, in our model 

coincide with the higher-P-wave-velocity areas (4.5–5.2 

km/s) in their models. The higher-P-wave-velocity area 

southwest of the San Andreas fault corresponds to outcrops 

of granodiorite and quartz diorite, which, because of their 

high quartz contents, can be expected to have lower VP /VS

ratios, consistent with our results. The low-VP /VS-ratio area 

southwest of the Zayante fault in layer 3 (3-km depth, fig. 

10C) probably indicates an extension of the granodiorite to 

depth. The corresponding P-wave velocity in this area ranges 

from 5.2 to 5.5 km/s (Eberhart-Phillips and others, 1990), 

within the expected range of P-wave velocities in granodiorites of 4.4 to 6.27 km/s at 0- to 3-km depth estimated from 

laboratory studies (Press, 1966). A broad high-VP /VS-ratio 

area also exists at 3-km depth along the San Andreas fault, 

consistent with the results of Eberhart-Phillips and others, 

who indicated the presence of high-P-wave-velocity rock 

at 3-km depth between the Sargent and San Andreas faults. 

They interpreted this rock to be a part of upthrust ophiolitic 

basement (presumably gabbro).

In layer 4 (7-km depth, fig. 10D), the VP /VS highs corresponding to the Boulder Creek and Corralitos anomalies are visible 

along the San Andreas fault, along with a low-VP /VS-ratio zone 

close to the coast. The presence of this low-VP /VS-ratio zone 

is consistent with the continued presence of such quartz-rich 

basement rocks as the quartz diorite, granodiorite, and granite 

that crop out on Ben Lomond and in the Bald Mountains to the 

northwest (Hanna and others, 1981). Below 7-km depth, VP /VS

ratios are mostly average (1.72–1.77) to low (1.55–1.69). In 

particular, a distinct discontinuity is evident between 7- and 

11-km depth that separates the high-VP /VS-ratio area along the 

fault zone that we have associated with shallower mafic rocks 

from the major VP /VS low visible in layer 5 (11-km depth, fig. 

10E). This observation agrees quite well with the seismic0.0000 0.0000 0.0031 0.0048 -0.0002

.0045 .0134 .0146 .0077 -.0031

-.0092 .0184 .0131 .0036 .0004

.0000 .0049 -.0020 .0054 .0060

.0000 .0001 .0005 .0000 .0000

.0000 .0000 .0037 .0030 .0000

.0090 .0182 -.0110 -.0043 -.0049

-.0068 -.0405 .0131 - .0026 .0042

.0000 .0301 .0000 .0013 -.0014

.0000 .0005 .0013 .0006 .0000

.0000 .0000 -.0009 -.0015 .0000

-.0062 -.0203 .0006 .0011 .0000

.0189 .1401 -.0316 .0021 .0005

.0000 -.0357 -.0127 .0006 -.0024

.0000 .0066 .0053 -.0009 .0000

.0000 .0000 -.0024 .0000 .0000

.0000 .0195 -.0103 -.0006 .0000

.0146 .7302 .0498 -.0156 .0000

.0000 .0037 -.0190 -.0057 .0000

.0000 -.0107 -.0079 .0025 .0000

.0000 .0000 .0000 .0000 .0000

.0000 .0290 .0015 .0000 .0000

.0000 .1935 .0161 .0000 .0000

.0000 .0258 .0121 .0000 .0000

.0000 -.0062 .0090 .0000 .0000

Table 5.—Resolution matrix for layer 5 node of three-dimensional inversion model.

[Diagonal element is listed in bold]

THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION
E160 EARTHQUAKE OCCURRENCE

refraction studies by Fuis and Mooney (1990), who reported a 

midcrustal boundary at about 10-km depth, and by Foxall and 

others (1993), who also claim that rocks of different composition are present at 7- versus 11-km depth. However, the interpretation of a gneissic midcrust by Fuis and Mooney (1990) 

(VP /VS =1.75–1.94; Press, 1966) is inconsistent with the low 

VP /VS ratios (1.65–1.68) near the San Andreas fault between 

y=0 km and y=−20 km obtained in our study, but possibly 

consistent with the average VP /VS ratio (1.72–1.77) between 

y=0 km and y=10 km (fig. 10E). In any case, the VP /VS ratios 

between y=0 km and y=−20 km at 11- and 16-km depth are 

inconsistent with those typically observed for gabbro, which 

has been suggested by Foxall and others (1993) to be present 

at these depths. Our interpretation, based on VP /VS ratios, is that 

the midcrust in the study area (fig. 1) is composed mainly of 

granitic rocks; that is, it is a part of the Salinian block. Wideangle seismic-reflection/refraction data obtained in the Loma 

Prieta region by Brocher and others (this chapter) showing 

that seismic velocities at 11- to 16-km depth range from 6.0 

to 6.4 km/s also suggest that the rocks at these depths may be 

a part of the Salinian block. 

In layer 6 (16-km depth, fig. 10F), we again observe a conspicuous VP /VS low along and southwest of the San Andreas 

fault near its intersection with the Sargent fault (fig. 10F), and 

possibly average to high VP /VS ratios slightly to the southeast. 

Resolution elsewhere at this depth is too poor to permit interpretation. Interestingly, these two anomalies bracket the 1989 

Loma Prieta main-shock hypocenter, possibly indicating a 

major variation in rock type in the immediate vicinity of the 

main shock that may play a role in the mechanical behavior of 

the fault. One major goal of our future work will be to improve 

the model resolution within this depth range to determine the 

significance of these particular anomalies.

CONCLUSIONS

A model of the three-dimensional velocity structure along 

the Loma Prieta rupture zone indicates the presence of 

significant variations in VP /VS ratio both laterally and with 

depth. Our model can account for 64 percent of the original variation in S−P time (assuming a constant VP /VS

 ratio). 

0.0000 0.0000 0.0000 -0.0023 0.0009

-.0124 .0020 .0015 -.0013 .0024

.0105 .0011 -.0005 .0078 -.0013

.0000 .0008 .0191 .0090 .0016

.0000 .0036 .0069 .0035 .0000

.0000 .0000 -.0009 -.0015 .0000

-.0150 -.0036 .0008 .0001 -.0001

-.0017 .0019 -.0056 -.0002 -.0028

.0000 .0268 -.0030 -.0008 .0044

.0000 .0056 .0092 .0084 .0000

.0000 .0000 .0000 -.0007 .0000

-.0020 .0086 -.0018 .0020 .0000

.0114 -.0226 .0060 -.0023 .0005

.0000 -.0138 -.0018 -.0007 .0081

.0000 .0098 .0169 .0039 .0000

.0000 .0000 .0017 .0000 .0000

.0000 -.0150 -.0008 .0001 .0000

.0270 .0258 .0015 .0008 .0000

.0000 .1563 .0083 .0132 .0000

.0000 -.0794 -.0365 .0113 .0000

.0000 .0000 .0000 .0000 .0000

.0000 -.0263 -.0010 .0000 .0000

.0000 .1417 .0120 .0000 .0000

.0000 .4937 .0940 .0000 .0000

.0000 .0542 .0087 .0000 .0000

Table 6.—Resolution matrix for layer 6 node of three-dimensional inversion model.

[Diagonal element is listed in bold]
E161

Using other geologic and geophysical information, we can 

interpret most of the anomalous features in our model. Two 

major findings are supporting evidence for the presence of 

mafic rocks at shallow depth along the fault zone (high VP /VS

ratios) and for a crustal discontinuity at about 10-km depth 

(a transition from mostly high to mostly low VP /VS ratio with 

depth). Unanticipated VP /VS lows near the surface are accountable for by the presence of quartz-rich rocks. Our model 

supports the suggestion that the Zayante fault represents a 

major boundary to 7-km depth, with low VP /VS ratios (mostly 

granodiorite) to the southwest. Northeast of the San Andreas 

fault, a near-surface VP /VS anomaly supports the presence of 

a 2- to 3-km-thick sandstone sequence. At 16-km depth, a 

possible boundary between high- and low-VP /VS-ratio rocks 

coincides with the location of the Loma Prieta main-shock 

hypocenter. With additional data from recent USGS passive and active experiments in the Loma Prieta region, we 

expect to be able to improve the resolution of our image of 

the Loma Prieta rupture zone.

ACKNOWLEDGMENTS

This research was supported by U.S. National Science 

Foundation grants EAR–9011784 and EAR–9117464; 

computations at the University of Wisconsin-Madison’s Geophysical and Polar Research Center computing facility were 

partly supported by U.S. National Science Foundation grant 

EAR–9019555. We thank Donna Eberhart-Phillips and Lynn 

Dietz for providing their P-wave-velocity models and event 

locations for our use; Bill Foxall and Susan Schwartz for 

sharing their data and ideas; Bob Jachens, Bob McLaughlin, 

and David Malone for helpful discussions; Bill Ellsworth 

and John Vidale for their reviews of the manuscript; Ray 

Wells for seeing the chapter to completion; and the IRIS 

Data Management Center staff for assistance in accessing 

the PASSCAL data.

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THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION
E162 EARTHQUAKE OCCURRENCE

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THREE-DIMENSIONAL VARIATIONS IN Vp/Vs RATIO IN THE LOMA PRIETA REGION
E164 EARTHQUAKE OCCURRENCE
E165

CONTENTS

Page

Abstract---------------------------------------------------------------------- E165

Introduction------------------------------------------------------------------ 165

Geologic setting------------------------------------------------------------- 167

Traveltime data-------------------------------------------------------------- 167

Inversion procedure--------------------------------------------------------- 169

Results------------------------------------------------------------------------ 171

 Velocity and resolution---------------------------------------------- 171

 Relocation of events------------------------------------------------- 178

 Determinations of fault-plane solutions--------------------------- 178

 Stress field inferred from focal mechanisms---------------------- 179

Discussion-------------------------------------------------------------------- 182

Conclusions------------------------------------------------------------------ 185

Acknowledgments---------------------------------------------------------- 185

References cited------------------------------------------------------------- 186

ABSTRACT

Three-dimensional models from P-wave velocities and 

the ratio of P- to S-wave velocity (VP /VS ) for the Loma Prieta 

region were developed from the inversion of local traveltime data (21,925 P-wave arrivals and 1,116 S-wave arrivals) from earthquakes, seismic-refraction shots, and blasts 

recorded at 1,700 stations in the U.S. Geological Survey’s 

Northern California seismic network (Calnet) and numerous 

portable-seismograph deployments. The velocity and density 

models and microearthquake hypocenters reveal a complex 

structure that includes a San Andreas fault extending to the 

base of the seismogenic layer. A high-P-wave-velocity body 

extends the length of the Loma Prieta rupture zone and fills 

the 5-km-wide volume between the main-shock rupture and 

the San Andreas and Sargent faults. We suggest that this 

body controls both the pattern of background seismicity on 

the San Andreas and Sargent faults and the extent of rupture 

during the main shock, thus explaining how the background 

seismicity outlined the along-strike and depth extent of the 

main-shock rupture on a different fault plane 5 km away. New 

aftershock focal mechanisms, based on three-dimensional ray 

tracing through the velocity model, support a heterogeneous 

postseismic-stress field and cannot resolve a uniform faultnormal compression. Both the subvertical (or steeply dipping) 

San Andreas fault and the fault surfaces that ruptured in the 

1989 Loma Prieta earthquake are parts of the San Andreas 

fault zone, this section of which does not have a single type 

of characteristic earthquake.

INTRODUCTION

At first glance, the 1989 Loma Prieta earthquake appeared 

to be an expected and easily understood event. By filling a 

gap in the background seismicity along the San Andreas fault 

with right-lateral slip, it fit the forecasts by Moths and others 

(1981), and the initial epicentral location to the west of the 

San Andreas fault was attributable to a mislocation of the 

main shock caused by across-fault velocity contrasts. With 

further investigation, however, it soon became clear that the 

earthquake occurred on a dipping structure, west of the surface 

trace of the San Andreas fault, and that it was an oblique-slip 

event with a thrust component and a spatially varying amount 

of slip and rake angle. Considering the varying strike of 

the San Andreas fault trace, Dietz and Ellsworth (1990 and 

this chapter) inferred that the Loma Prieta slip-and-rupture 

geometry is ideal for accommodating plate motion and that 

an additional subvertical San Andreas fault segment at the 

depth of the Loma Prieta rupture is not needed. Schwartz and 

others (1990) also presented a model without a subvertical San 

Andreas fault segment, but they instead required a 65° NE.-

dipping fault in addition to the 65° SW.-dipping Loma Prieta 

rupture. Olson (1990) examined background seismicity and 

inferred an active subvertical San Andreas fault in addition 

to the southwest-dipping Loma Prieta rupture.

These observations lead to several questions that we 

investigate in this paper by inverting local-earthquake and 

controlled-source arrival-time data to determine a detailed 

three-dimensional velocity model for the Loma Prieta region. 

How does the fault surface for this earthquake relate to the 

San Andreas fault and the major geologic units? Can spatially 

varying material properties account for the spatial variations 

in the amount of slip? Why did the main shock fill a microTHE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

SEISMOTECTONICS OF THE LOMA PRIETA REGION DETERMINED FROM

THREE-DIMENSIONAL P-WAVE VELOCITIES, VP / VS RATIOS, AND SEISMICITY 

By Donna M. Eberhart-Phillips1 and Andrew J. Michael, U.S. Geological Survey

1

Current affiliation: Institute of Geological and Nuclear Sciences, 

Dunedin, New Zealand.

E165
E166 EARTHQUAKE OCCURRENCE

seismicity gap despite occurring on a subparallel surface 

several kilometers from the background seismicity? Was 

the post-main-shock state of stress heterogeneous or simple 

fault-normal compression?

Several previous studies have focused on the problems of 

the three-dimensional velocity structure in the Loma Prieta 

region (fig. 1). Eberhart-Phillips and others (1990a, b) discussed an inversion of primarily aftershock data recorded by 

the U.S. Geological Survey (USGS)’s permanent northern 

California seismic network (Calnet). Lees (1990) and Shalev 

and Lees (this chapter) used Calnet aftershock data to produce 

a tomographic image of velocity contrasts in the aftershock 

zone. Their approach, however, did not include new ray 

tracing through the three-dimensional structure, a step that 

Kissling and others (1994) showed to be critical when the 

sources are inside the structure, as is the case here. Foxall 

and others (1993) included three-dimensional ray tracing 

in their inversions of Calnet aftershock data to determine 

the P-wave-velocity structure of the Loma Prieta region. In 

this paper, we seek to improve the spatial resolution and, 

most importantly, to obtain a model based on VP /VS ratios. 

These goals require combining the largely vertical component Calnet data with data from varied portable-instrument 

deployments. To improve the spatial distribution of sources 

and to remove some of the tradeoffs between velocity and 

hypocentral location, we also use various active sources, 

including explosions, throughout the Loma Prieta region 

and airguns deployed to the southwest. This expanded data 

set yields not only VP /VS ratios in addition to P-wave velocities, but also a more complete image of the crust, instead of 

focusing only on the region within and above the zone of 

seismicity. A velocity model with a larger volume than just the 

aftershock zone is important for such purposes as including 

realistic crustal rheology in numerical models (for example, 

Eberhart-Phillips and Stuart, 1992).

Using the new three-dimensional velocity model, we relocated 8,577 events that occurred in the aftershock zone from 

1968 through September 1993. We determined new focal 

mechanisms for 3,964 of these events, about 2,600 of which 

we judged to be sufficiently reliable to use in seismotectonic 

and stress interpretations. With this new data set, we now 

determine whether the three-dimensional ray tracing requires 

SAN ANDREAS FAULT

CALAVERAS FAULT

Santa Clara

Basin ZAYANTE-VERGELES

FAULT

.

.

.

.

.

Loma Prieta

Onshore-Offshore

Profiles

SARGENT FAULT

North American

plate

Pacific

plate

PA

OCEAN

CIFIC

EXPLANATION

0 10 20 KILOMETERS

37°50'

37°40'

37°30'

37°20'

37°10'

37°00'

36°50'

36°40'

36°30'

122°40' 122°20' 122°00' 121°40' 121°20' 121°00' 120°40'

SAN 

FRANCISCO

BAY

SAN 

FRANCISCO

MARIN

PENINSULA

LOS ALTOS

DAVENPORT GILROY

HOLLISTER

Salinian

Franciscan

Great Valley

assemblage

sequence

block

SANTA CRUZ

MOUNTAINS

Figure 1.—Loma Prieta region, Calif., showing generalized geology (adapted from Irwin, 1990), locations of major faults 

(dotted where inferred), approximate surface projection of rupture zone (shaded area), and velocity grid for three-dimensional 

model (within rectangle). Main-shock focal mechanism is plotted at hypocenter.
E167

any changes in previous tectonic interpretations. We use the 

new seismicity catalog to reevaluate the stress state before 

and after the main shock by considering the stress inferred 

from focal mechanisms. Finally, we combine the velocity 

models with the seismicity results to interpret faults and 

consider the relations between seismic velocity and seismogenic behavior.

GEOLOGIC SETTING

The Loma Prieta region is dominated by the Pacific-North 

American plate boundary (fig. 1), of which the major element 

is the San Andreas transform fault. Over the past 23 m.y., 

as much as 315 km of strike-slip movement has occurred on 

this fault, so that the Salinian block of plutonic rocks on the 

southwest is juxtaposed against the Franciscan subduction 

complex on the northeast (Page, 1990). During the fault’s 

history, however, various factors have caused the geology 

to be more complex than a simple strike-slip fault separating uniform blocks of material. The fault has changed its 

position, at times stranding slivers of material (Sims, 1993). 

The plate motion changed slightly about 3.5 Ma, so that a 

compressional component exists (Harbert and Cox, 1989), 

generally taken up off the San Andreas fault (Zoback and 

others, 1987).

The Loma Prieta region is transitional between the San 

Francisco Bay region and the narrower creeping section 

of the San Andreas fault. In the San Francisco Bay region, 

geodetic analysis of velocity data shows that the plate boundary extends approximately 50 km northeastward of the San 

Andreas fault (Lisowski and others, 1991). In contrast, the 

geometry of the plate boundary in the creeping section is 

considered to be relatively simple with the deformation 

through the lithosphere centered below the surface trace of the 

San Andreas fault (Furlong and others, 1989). The velocity 

field in the transitional region is not well modeled by block 

motion (Lisowski and others, 1991), and a bend in the fault 

has formed in the Loma Prieta region.

Fuis and Mooney (1990) summarized the crustal structure 

and geology of this transitional region. Northeast of the San 

Andreas fault, mafic crust is inferred at 8-km depth below 

the Franciscan assemblage. Directly northeast of the fault, 

several thrust faults complicate the picture by bringing in 

slivers of ophiolite and sedimentary rock, as described by 

McLaughlin and Clark (this chapter). Southwest of the San 

Gregorio fault lies an inactive accretionary prism consisting of the San Simeon terrane (Franciscan assemblage) and 

Tertiary sedimentary rocks. Page (1990) inferred that Salinian basement lies on both sides of the San Gregorio fault, so 

that the San Gregorio fault does not consistently bound the 

Salinia terrane. Page and Brocher (1993) and Brocher and 

others (this chapter) further delineated the crustal structure 

and tectonic history of the transitional region by showing that 

oceanic crust underlies the accretionary prism and Salinian 

block. They inferred that oceanic crust extends from the 

deep-sea floor to the San Andreas fault, dipping landward 

8°–10°. Holbrook and others (1996) interpreted wide-angle 

profiles across the fault near San Francisco and parallel to 

San Francisco Bay. They also inferred that active fault zones 

extend throughout the crust and that Pacific oceanic crust 

may not extend under the Salinian block.

The structure of the San Andreas fault at depth in the 

Loma Prieta region is unclear from surface geology. Just as 

evidence of near-surface heterogeneity exists from slivers of 

rock, other slivers or blocks of material may exist at greater 

depth along the fault that influence fault behavior. To study 

the plate boundary and understand seismic deformation, the 

fault zone needs to be better defined within the larger-scale 

setting of crustal blocks.

TRAVELTIME DATA

For this study, we tried to assemble the most complete 

traveltime data set possible (fig. 2). In previous studies, we 

used primarily Calnet data and solved for P-wave velocity only (Eberhart-Phillips and others, 1990a; Michael and 

Eberhart-Phillips, 1991). To increase the variety of raypaths 

and, therefore, the resolution of the three-dimensional inversion, we now include data from the temporary stations that 

were deployed after the Loma Prieta main shock and from 

three-component stations, so that we can solve for VP /VS ratio 

in addition to P-wave velocity.

Several arrays of temporary instruments were deployed 

after the main shock. A total of 60 seismic cassette recorders 

were deployed by the USGS. These instruments are normally 

used for recording seismic-refraction shots, can record only 

for a maximum of 13-minute intervals, and can record a 

maximum of 30 minutes on a single tape (Murphy, 1988). 

With a field crew changing tapes, the instruments recorded 

for six 13-minute periods on October 19, 1989, and during 

those 78 minutes of operation they recorded 59 aftershocks. 

The University of California, Berkeley (UCB), deployed 

six DR100 recorders from October 19 to November 16. The 

USGS deployed General Earthquake Observing System 

(GEOS) recorders at 94 different sites from October 17, 1989, 

to March 14, 1990 (Mueller and Glassmoyer, 1990). Reftek 

recorders were deployed by the Program for Array Seismic 

Studies of the Continental Lithosphere (PASSCAL) of the 

Incorporated Research Instruments for Seismology (IRIS) at 

21 sites from October 27 to November 6, and at six sites from 

November 7 to November 23. We obtained P-wave-arrival 

and S−P times for PASSCAL, UCB, and selected GEOS stations from C.H. Thurber (written commun., 1993).

The best spatially distributed three-component array was 

deployed in fall 1990 by Takauchi and Evans (1995) for the 

purpose of recording teleseisms; however, numerous aftershocks were also recorded during the deployment. A total of 

29 sites were occupied for 5 months, using 25 5-day analog 

SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E168 EARTHQUAKE OCCURRENCE

recorders, 3 GEOS recorders, and 2 Reftek recorders. For this 

study, we used 69 events that form a spatially well distributed 

data set from the 1990 array. To improve the spatial distribution of sources, we also used 85 events that were recorded on 

Calnet before the Loma Prieta main shock.

To reduce the tradeoff between hypocentral location and the 

velocity model, we also used arrival-time data from known 

sources, including 19 quarry blasts recorded on Calnet and 

seven shots recorded on both Calnet and seismic-refraction 

arrays before 1989. Those seismic-refraction data cross the 

San Andreas fault near Gilroy (Mooney and Colburn, 1985), 

traverse the Santa Clara Valley 15 km south of Gilroy (Mooney 

and Luetgert, 1982), and run southwest of the San Andreas 

fault from Los Altos to Davenport (Boken and Mooney, 

1982).

In May 1991, the USGS conducted a seismic-refraction profile across the Zayante-Vergeles, San Andreas, and 

Sargent faults (Murphy and others, 1992). Six shots were 

recorded on a 75-station distributed array reoccupying sites 

of temporary earthquake stations; at most of these stations, 

horizontal seismometers were deployed in addition to vertical 

seismometers. Three shots were recorded on a 25-km-long 

across-fault profile with a 150-m station spacing. Three shots 

were recorded on a long profile from the Marin peninsula, 

along the San Francisco peninsula, through the Loma Prieta 

region. We include P-wave-arrival times (and S-wave-arrival 

times, where available) from all of the shots and stations in 

the study area (fig. 1).

The onshore/offshore data collected by Brocher and others 

(1992) provide another valuable data set, particularly for the 

Salinian/southwestern block (fig. 1), which is poorly sampled 

by the aftershock data and onshore stations. Brocher and others 

(1992) deployed 5-day recorders onshore to record offshore 

airgun profiles. In their analysis, the recording site is considered the shotpoint, and the source locations are considered 

the receivers. They interpreted both first and later arrivals to 

infer that the crust southwest of the San Andreas fault has a 

northeast-dipping Moho with a layer of subducted oceanic 

37°40'

37°30'

37°20'

37°10'

37°00'

36°50'

36°40'

122°40' 122°20' 122°00' 121°40' 121°20' 121°00'

-45

-25

-10

-3410

18

26

35

50

69

Y=

-60

-32

-15

-7

15

25

42

X=-4,-1,2,5,8

Stations:

Calnet station

Other station

Sources:

Event

Blast

Shot

Grid point

0 10 20 KILOMETERS

Figure 2.—Loma Prieta region, Calif., showing locations of sources and receiver stations superimposed on fault base map. Dots, velocity 

grid. Origin of grid is at the San Andreas fault near main-shock hypocenter (fig. 1). x is parallel to fault, positive to northeast; y is normal 

to fault, positive to southeast.
E169

crust (see Brocher and others, this chapter). From the seismicreflection lines, they were able to carefully map the bathymetry 

and the bottom of the Neogene sedimentary deposits along 

the airgun profiles. To use their data in our inversions, we 

stripped off the water and sedimentary layers. We traced rays 

in their models and then adjusted the traveltimes by the time 

traveled in the water and sedimentary layers, and changed 

the “receiver” locations to be at the raypath location at the 

base of the Neogene sedimentary deposits. To allow for some 

uncertainty in this correction process, we treated these shots 

as blasts by allowing the origin time to be free.

Our data set thus includes 1,700 stations. We use 273 events 

with 20,284 observations, and 54 shots and blasts with 2,757 

observations, including 21,925 P-wave arrivals and 1,116 

S-wave arrivals.

INVERSION PROCEDURE

The local aftershock arrival-time data are simultaneously 

inverted for hypocenters and three-dimensional velocity structure, as developed by Thurber (1983) and Eberhart-Phillips 

(1990) and described by Thurber (1993) and Eberhart-Phillips 

(1993). The algorithm has been revised to solve simultaneously for P-wave velocity and VP /VS ratio, using the same 

formulation as that of Thurber and Atre (1993). We prefer 

to obtain P-wave-velocity and VP /VS -ratio models rather 

than P- and S-wave-velocity models because VP /VS ratio is 

similar to Poisson’s ratio and is an important parameter in 

characterizing rock properties and rheology. Particularly, 

in such cases as ours where S-wave-velocity data are less 

numerous and of poorer quality than P-wave-velocity data, 

S-wave velocities would be poorly resolved relative to P-wave 

velocities, making the interpretation of VP /VS -ratio variations 

difficult (Eberhart-Phillips, 1990). We assume that, given a 

three-dimensional heterogeneous P-wave-velocity model and 

unknown S-wave velocities, estimating an S-wave-velocity 

model from P-wave velocities using a constant VP /VS ratio 

would be more useful than assuming a homogeneous or onedimensional S-wave-velocity model. Alternatively, we could 

assume a crustal lithology with only one-dimensional variations in rigidity, despite known three-dimensional variations 

in compressibility, implying large variations in Poisson’s 

ratio wherever the three-dimensional pattern differed from 

the one-dimensional pattern.

The velocity of the medium is parametrized by assigning 

velocities at the intersections (gridpoints) of a three-dimensional grid. The raypaths are calculated with an approximate 

three-dimensional ray-tracing algorithm that produces curved, 

nonplanar raypaths which are defined by points more closely 

spaced (1 km) than the velocity gridpoints (3–20 km). The 

solution is obtained by iterative damped least-squares inversion. For each iteration, new raypaths are determined, the 

hypocenter solution is included, and parameter separation is 

carried out. Throughout the inversion, weighting is applied to 

each observation on the basis of both the size of the traveltime 

residual and the source-receiver distance. The full-resolution 

matrix is then calculated.

The damping parameter for each inversion is chosen 

empirically by evaluating a tradeoff curve of data variance 

and solution variance, because the damping will vary with 

the model grid and the data set (Eberhart-Phillips, 1986). 

This selective damping greatly reduces the data variance, with 

only a moderate increase in solution variance. Note that for a 

damped-least-squares inversion, any significantly large damping parameter will reduce the resolution, so that the diagonal 

resolution will be far from 1. Thus, the resolution matrix must 

be evaluated by other means than simply by inspecting the 

diagonal elements.

We performed a series of simultaneous hypocenter-velocity 

inversions, progressively increasing the complexity of the 

model (Eberhart-Phillips and Michael, 1993). This procedure 

obtains a reasonable estimate of velocities throughout the 

entire volume. We aimed at obtaining a velocity model that 

would be useful for other purposes, such as inclusion of 

three-dimensional crustal rheology in numerical modeling 

of deformation. Thus, we preferred an inversion procedure 

that provides simple but reasonable velocities outside the 

high-resolution area, rather than giving either undefined 

velocities or uninterpretable velocity artifacts outside the 

immediate aftershock zone. The pattern of subsurface velocity variations probably will resemble the overall distribution 

of lithologic units and faults that are mapped at the surface. 

Thus, we assume greater variation in material properties in 

the direction normal than parallel to the major faults, and 

we treat the traveltime data so as to optimize structures subparallel to the San Andreas fault. Using tectonic knowledge 

to parametrize a velocity model is a widely used strategy. 

For example, Abers and Roecker (1991), Foxall and others, 

(1993), Thurber and others (1997), and Shalev and Lees 

(this chapter) all used grids designed to maximize the resolution of across-fault velocity heterogeneity. Alternatively, 

an even grid spacing or smoothing can be used. Parsons 

and Zoback (1997) used 10-km horizontal smoothing for 

a San Francisco Bay model that included the northwestern 

part of our model, and Dorbath and others (1996) used an 

even 5-km grid spacing for a Hollister velocity model that 

included the southeastern part of our model (fig. 1). Both of 

these models show a similar along-strike extent of velocity 

features; however, they do not resolve across-fault velocity 

gradients so sharply.

We do not incorporate theoretical velocity information 

into the model, but obtain initial models from the data with 

progressive inversions (Kissling and others, 1994). For a 

synthetic example with a sharp across-fault gradient, the true 

model was recovered more accurately and with fewer artifacts 

when progressive inversions were done (Eberhart-Phillips and 

others, 1995). An alternative single-step inversion with evenly 

spaced grid points across the San Andreas fault is discussed 

below in the section entitled “Results.” The grid area is shown 

SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E170 EARTHQUAKE OCCURRENCE

in figure 2. For all inversions, stations outside the grid are 

allowed station corrections to avoid smearing external velocity 

anomalies into the grid.

We begin with a one-dimensional inversion of the entire 

data set for P-wave velocities and hypocenters. We position 

a sharp velocity gradient to approximate the Moho at 25-km 

depth, with P-wave velocity fixed at 8 km/s at 26-km depth. 

The position of the Moho is poorly constrained, but the 25-

km depth fits the data slightly better than do other depths 

(data variance, 2 percent lower than for 22- or 28-km depth), 

consistent with the modeling of wide-angle data by Holbrook 

and others (1996). Brocher and others (this chapter) determine a dipping Moho from later arrivals. Thus, we assume a 

Moho that dips from 15-km depth at 85 km southwest of the 

San Andreas fault to 25-km depth at the fault and then is flat 

farther northeast (fig. 3A), and we evaluate this location with 

two-dimensional inversions, using only onshore-offshore 

data. A dipping Moho fits the data slightly better but is poorly 

constrained by first-arrival data.

We next obtain an approximately two-dimensional Pwave-velocity model of regional variations across the plate 

boundary. We use the onshore-offshore data, which provide 

almost all of our raypath coverage southwest of the Loma 

Prieta region (fig. 1), together with a subset of 71 spatially 

distributed earthquakes, so that the two-dimensional inversion 

does not overly focus on the aftershock zone. The inversion is 

not purely two dimensional but allows for some variation in 

velocity along the fault, using two y-grids, one at y=−45 km 

(45 km northwest) and the other at y=69 km (69 km southeast). 

The z-grids are at 0-, 2-, 4-, 6-, 8-, 10-, 12-, 15-, and 18-km 

depth. We use a constant VP /VS ratio with the two-dimensionalinversion results as our initial three-dimensional model. The 

S-wave data come almost entirely from events and stations in 

the aftershock zone and so do not sample regional variations. 

We selected a uniform VP /VS ratio of 1.74 from Wadati plots. 

The two-dimensional model has a 24-percent reduction in 

variance from the one-dimensional model.

The coarse three-dimensional grid has 20- to 34-km spacing in the y-direction, and the initial P-wave-velocity model 

is interpolated from the two-dimensional P-wave-velocity 

model. The three-dimensional inversions use the full set of 

earthquakes and shots. The coarse three-dimensional inversion 

4

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DEPTH, IN KILOMETERS

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x-DISTANCE, IN KILOMETERS

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F

Figure 3.—Initial one-dimensional P-wave-velocity model (A), with a dipping Moho estimated from Brocher and others (1992), and 

southwest-northeast cross sections of coarse three-dimensional P-wave-velocity model for grid layers at y=–45 km (B), y=–10 km (C) 

y=10 km (D), y=35 km (E), and y=69 km (F). Contour interval, 0.25 km/s. Crosses in figures 3B through 3F, hypocenters of events used 

in inversion.
E171

reduces the variance by a further 22 percent relative to the onedimensional model; the fine three-dimensional inversion has 

a y-grid spacing of 6 to 20 km, and the initial P-wave-velocity and VP /VS-ratio models are interpolated from the coarse 

three-dimensional inversion. Only the well-sampled central 

part of our model, from x=−7 to x=15 km, is free in the fine 

three-dimensional inversion; outside this area, velocities are 

fixed to the coarse three-dimensional values. The fine threedimensional inversion reduces the variance by a further 27 

percent relative to the one-dimensional model.

Some near-surface variations in velocity may not be resolvable with our relatively smooth three-dimensional gridded 

model, and so they should be included to improve the locations 

of the hypocenters. Using the fine three-dimensional results, 

we use additional iterations with lower damping where we 

solve only for velocities in the upper 2 km within the finely 

gridded central area and allow station corrections for stations 

located outside of this area. This procedure reduces the variance by a further 5 percent relative to the one-dimensional 

model, so that the progressive three-dimensional inversion 

reduces the variance by a total of 78 percent. The final rootmean-square (rms) traveltime residual is 0.118 s.

Hypocenter locations will be most accurate in the central 

part of the model, where most of the inversion hypocenters 

occur. To evaluate the spatial variation in hypocenter-location 

accuracy, we relocated the shots and blasts, using the threedimensional model and only the Calnet data. The four shots 

in the profile across the San Andreas fault near the main shock 

locate within 0.27 km of their true epicenters. Shots near the 

fault at y=−36 km and across the fault at y=15 km locate within 

0.36 km of their true epicenters. At the southeast end of the 

model, blast relocations are not so accurate, mislocating 1 

km southwest at y=35 km and 1.7 km southwest at y=45 km. 

Thus, some of the apparent offset between the southeastern 

epicenters and the San Andreas fault trace may be due to lower 

hypocentral accuracy in the southeastern part of the model. A 

set of station corrections determined solely for this area would 

likely move the events back onto the San Andreas fault, but we 

chose to keep the location process consistent for the entire data 

set, rather than to use multiple sets of station corrections.

RESULTS

VELOCITY AND RESOLUTION

The resolution matrix describes the distribution of information for each node, such that each row is the averaging vector 

for a given parameter. The relative size and pattern of the 

offdiagonal elements show how the information is smeared. 

For a node to be adequately resolved, its resolution should be 

peaked and have no significant contribution from nonadjacent 

nodes. For a succinct way of comparing the resolution, we 

calculate the spread function (Michelini, 1991; Michelini and 

McEvilly, 1991), which compresses each row of the resolution 

matrix into a single number which describes how strong and 

peaked the resolution is for that node. As described in detail 

by Toomey and Foulger (1989), the spread function is a better 

way to describe the resolution than by examining only the 

diagonal element, because the diagonal resolution depends 

strongly on grid spacing and damping. As defined by Michelini 

and McEvilly (1991), the spread function Sj for a given node 

is computed from all the elements skj of the corresponding row 

of the resolution matrix, weighted by their distance Djk from 

the node, and Sj is the diagonal element Sjj:

This relation combines two issues in evaluating resolution: 

the amount of information and the amount of smearing. The 

first factor makes the spread function small for nodes with 

large diagonal resolutions. The summed terms make the 

spread function large for nodes that have significant averaging from other nodes, and particularly large if the contributing nodes are more distant. Because the spread function 

combines these factors, it does not have such physical units 

as kilometers, similar to the resolution matrix, which also 

does not have physical units. The range of values obtained 

for the same data set will vary, depending on the chosen grid 

spacing and damping.

We also illustrate the pattern of image blurring in lowresolution areas by showing contours of the averaging vectors 

for nodes with significant smearing. For local earthquake 

tomography, this procedure is more useful than showing 

ellipses because the smearing commonly is in only one direction and ellipses would wrongly imply that the averaging is 

centered on the node. Each node has a row of the resolution matrix, with offdiagonal elements for all other nodes, 

showing the spatial averaging for that node. For each row 

of the resolution matrix, we plot where the resolution is 70 

percent of the diagonal element. The resolution for the coarse 

three-dimensional model is plotted in figure 4. Resolution 

contours are shown for only those nodes with such smearing 

that the resolution contour extends beyond an adjacent node; 

at other nodes, we would expect the spatial averaging to be 

represented by the grid spacing.

For small values (<2) of the spread function, the model velocity is representative of the volume surrounding a given node, 

that is, the volume between that node and adjacent nodes. For 

moderate values (2–4) of the spread function, the velocity may 

be averaging a larger volume, and so small features may not be 

imaged or may appear as broader features. The largest values 

(>4) of the spread function indicate little or no information, and 

the associated velocity in the damped-least-squares inversion 

remains close to that in the initial model.

A comparison of the coarse three-dimensional model and 

inversion hypocenters (fig. 3) with the resolution plots (fig. 

4) indicates that the resolution is best where the density of 

SEISMOTECTONICS OF THE LOMA PRIETA REGION 

S

s

s

s j D

j

jk

j

jk

k

N

=



































 =

log ∑ . 1

2

1
E172 EARTHQUAKE OCCURRENCE

events is highest. The resolution is generally poor below 12-

km depth, except where deeper events occurred. The surface 

nodes generally are not constrained by raypaths independent 

of the 2-km-deep nodes. The resolution of VP /VS ratio is generally low, and virtually no information is obtainable outside 

the central area with three-component aftershock stations. 

Because the damping allows only small velocity perturbations in areas of low resolution, velocity features that are 

artifacts are avoided. However, the spatial shape of velocity 

features may be inexact, and spatially small velocity features 

will not be imaged in poorly resolved areas. For example, the 

southwestern part (x=−32 to −15 km) of the y=69 km grid 

(fig. 3F) shows a higher velocity than that in the initial model 

at 0- to 12-km depth; however, the resolution contours are 

0

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1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5

SPREAD FUNCTION

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x-DISTANCE, IN KILOMETERS

DEPTH, IN KILOMETERS

A

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I J

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B

y = –45 km y = –45 km

y = –10 km

y = 10 km

y = 35 km

y = 69 km y = 69 km

y = 35 km

y = 10 km

y = –10 km

Figure 4.—Resolution plots of P-wave velocity (A, C, E, G, I) and VP /VS ratio (B, D, F, H, J) for coarse three-dimensional P-wave-velocity model (figs. 3B–3F) of velocity grid in the Loma Prieta region (figs. 1, 2) for grid layers at y=– 45 km (A, B), y=–10 km (C, D) y=10 

km (E, F), y=35 km (G, H), and y=69 km (I, J). Spread function is shown with contours of averaging vector (row of resolution matrix) 

for nodes with smearing greater than grid spacing. Contours, which are 70 percent of diagonal resolution, are plotted where resolution 

contour extends farther than adjacent nodes, but not for nodes with near-zero resolution.
E173

vertically smeared. Thus, velocities may be high below 4-km 

depth and relatively low in the uppermost few kilometers of 

the crust, but any shallow, moderately low velocity cannot 

be imaged distinct from a deeper high velocity.

P-wave velocities in the southwestern part of our model 

range from 6.1 to 6.35 km/s at 5- to 15-km depth and vary 

only slightly from northwest to southeast. Overall, the range 

of P-wave velocities in this Salinian material is similar to 

the range of 6.0 to 6.4 km/s at 5- to 15-km depth modeled 

by Holbrook and others (1996) from wide-angle data near 

San Francisco. The central part of our model (x=−7 to 8 km) 

shows more complexity, with distinct gradients and high 

velocity features.

In the northeastern part our model, the area between the 

Calaveras and San Andreas faults (x=15 km) contains a large 

volume of material at 2- to 10-km depth with a P-wave velocity ranging from 5.1 to 5.9 km/s (fig. 5). Along the Santa 

Clara Basin (fig. 1; x=15 km, y=−45 to −25 km), however, 

the shallow velocities are lower, and the deeper velocities are 

higher, with 6.25-km/s material at 10-km depth. Some velocity contrast is evident across the Calaveras fault, with higher 

P-wave velocities on the northeast side below 6-km depth. 

P-wave velocities in the northeastern part of our model are 

also similar to those in Holbrook and others’ (1996) block 

model of San Francisco Bay, except that Holbrook and others 

inferred a discontinuity that does not appear in our model; 

they also showed material with P-wave velocities ranging 

from 5.5 to 6.1 km/s at 3- to 17-km depth overlying material with P-wave velocities ranging from 6.9 to 7.3 km/s at 

17- to 24-km depth.

Cross sections of the fine three-dimensional model are 

shown in figure 6, and the resolution is plotted in figure 7. 

Figure 5.—Map views of final three-dimensional model of velocity grid in the Loma Prieta region (figs. 1, 2), showing surface 

traces of faults (thin white lines) and locations of hypocenters (crosses) within 1 km of each layer for grid layers at 4-km (A), 

6-km (B), 8-km (C), and 10-km (D) depth. Faults: CAL, Calaveras; SAF, San Andreas; SAR, Sarmiento; ZV, Zayante-Vergeles. 

P-wave-velocity contour interval, 0.5 km/s, with 5.0- and 6.0-km/s contours in bold. Yellow heavy straight line in figure 5C

denotes location of cross section in figure 12E.

SEISMOTECTONICS OF THE LOMA PRIETA REGION 

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+

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+

+ +

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+

+

+

+

+

x-DISTANCE, IN KILOMETERS

y-DISTANCE, IN KILOMETERS

P-WAVE VELOCITY, IN

KILOMETERS PER SECOND

-32

-15

-7

-1

5

15

25

42

-45 -25 -10 4 18 35 50 69

-40 -30 -20 -10 0 20 40 6 10 30 50 0 70

-30

-20

-10

0

10

20

30

40

6

6 6

6

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+ +

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CAL

SAF SAR

ZV

A

B

C

D

3 4 5 6 7

+ Background Seismicity

+ Aftershocks

Depth: 4 km Depth: 8 km

Depth: 6 km Depth: 10 km
E174 EARTHQUAKE OCCURRENCE

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 4

5

6

7

SAF

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 4

5

6 6

7

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ZV

SAF

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4

5

6

7

+

+

+

+

+

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+

+

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ZV

SAF SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 4

5

6

6

7

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 5

6

6

7

+

+

+

+

+

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+

+

+

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 4

5

6

7

+

+

+

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+

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 4

5

6

7

+ +

+

+

+

+

+

+

+

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+

+

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4

5

6

7

+++++ ++

+ + + +

+

+ +

+

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+

+

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+

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ZV

SAF

SAR

CAL

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4

5

6

6

7

++

+

+

+

+

+

+

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+

+

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+

+

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+

+

+

+

+

+

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+

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SAFCAL

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 4

5

6 6

7

SAFCAL

-15 -7 -4 -1 2 5 8 15

18

12

6

0 4 4

5

6

7

+

+

+

+

+

+

++

+

+

+

+

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+

+

+

+

+

+

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+

+

+

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+

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+

+

+

+

ZV

SAF

SAR

LP

SV

SW NE

DEPTH, IN KILOMETERS

x-DISTANCE, IN KILOMETERS

P-WAVE VELOCITY,

IN KILOMETERS PER

SECOND

A

y=–45 km

y=–25 km

y=–10 km

y=–3 km

y=4 km

y=10 km

y=18 km

y=26 km

y=35 km

y=50 km

y=69 km

3 4 5 6 7

+ Background

 Seismicity

+ Aftershocks

Figure 6.—P-wave-velocity (A) and VP /VS-ratio (B) cross sections of final three-dimensional model of velocity grid in 

the Loma Prieta region (figs. 1, 2), showing locations of mapped fault traces and of hypocenters within 3 km of each 

layer. Faults: CAL, Calaveras; SAF, San Andreas; SAR, Sarmiento; ZV, Zayante-Vergeles. Contour intervals: P-wave 

velocity, 0.5 km/s, with 4.0-, 5.0-, 6.0-, and 7.0-km/s contours in bold; VP /VS ratio, 0.05, with 1.7, 1.8, and 1.9 contours in 

bold. In cross sections for y=4 km, estimated locations of southwest-dipping fault surface (LP) and the subvertical San 

Andreas fault (SV) are shown; both LP and SV are considered parts of the San Andreas fault system at depth.
E175

The velocity images (fig. 6) extend farther southwestward 

than the nodes in the fine three-dimensional inversion (x=−7 

to 15 km) so as to plot all of the relocated aftershocks. The 

resolution is generally good at 4- to 12-km depth. Key features of the P-wave-velocity model are a large low-velocity 

body in the southwest, a deep high-velocity body in the 

southwest, and several small, shallow high-velocity bodies. 

The low-velocity body parallels the San Andreas fault from 

the northwest to y=20 km and is most conspicuous at y=4

km (fig. 6), where it is clearly bounded on the northeast by 

the San Andreas fault and its P-wave velocity ranges from 

2.8 km/s at the surface to 5.7 km/s at 6-km depth. This lowvelocity body appears to be bounded by a vertical fault on 

the southwest, as shown in the y=−3 km and y=4 km cross 

-15 -7 -4 -1 2 5 8 15

18

12

6

0 1.75 1.751.751.75 1.751.75 1.75 SAF

-15 -7 -4 -1 2 5 8 15

18

12

6

0 1.8

1.8

1.8

1.75

1.75

1.75

+

+

+

+ +

+

+

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+

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ZV

SAF

-15 -7 -4 -1 2 5 8 15

18

12

6

0 1.75 1.8 1.75

1.75

1.75

1.75

1.85

1.85

1.85

+

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ZV

SAF

SAR

VP / VS RATIO

-15 -7 -4 -1 2 5 8 15

18

12

6

0 1.8

1.75

1.75

1.75

1.75

1.75

+

+

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0

1.8

1.8

1.8 1.8

1.8 1.8

1.75 1.75

1.75

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0

1.8

1.75

1.75 1.75

1.75

+

+

+

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0

1.8

1.75 1.75

1.75

1.75

1.75

+

+

+

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0

1.75

1.75

1.75 + +

+

+

+

+

+

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+

+

+

++

+++

+

+

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+ + +

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++

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+

++

+ +

+

+

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+

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+

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+

+

+

++

+

++

+

+++

+

+

++++

+

+

+

+

+

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+

+

++

+++

+

+

+

+

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++

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+ +

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+

+

++

+

+

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+

+

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+

+

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+

+

+

+

+++++

+

++ +

+

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+

++

+ +

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+ +

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+ +++

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++

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+ +

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+

+ + + +

+ +

+ +

+

+

+ +

+

++

+ +

+ ++

++

+ +

+

++ + ++ ++

++

+

+

+

+

+

+

+

+

++

+

++

+

+

+

+

+

+

+ +

ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

18

12

6

0

1.75

1.75

+++++ ++

+ + + +

+

+ +

+

+ + ++ +

+ +

+ +

+

+

+

+++

+

+

+++

++

++ ++

+

+

++

+

+

+

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+

+

++

++++

++

+

+

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+

++

+

+++

+

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+

+

+

+

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+ + + ++

+++

+

+

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++

+

++

++ +

+

+

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+

+++

+

+

+++

+

+

+

+

+

+

+

+

+++

+

+

++

+

+

+

++++++

+

+

+ +

+

++

+

+++++

+

+ +

+

++

++

+

+

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+

++

+

+

+

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+

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++

++

+

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+ ++

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++

+ +

+

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+ +

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++

+

++ +

+

+

+ +

+ + ++

+ +

+

++ +

+

++++

+ + ++

+

+ ++

+

++

+

+

+

+

+++

+++ + +

+

+

+

+

+

+

+

+

++++ + +

+

+

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+

++

++

+ +

+

+ + + +

+ +

++ + +

++ +

+

+ ++

+ +

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+ ++

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++

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+ +

++ +

+

+ +

+

+ ++ +

+ +

+

++

+

+

+ +

+

++++ +

+

+

++

+ + +

++

+

+ +

+

+ +++

+ +

+ + +

+

+

+

++ +

+ +

+

+

+

+

++

++++

+

+

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+

+ +

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++

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++ +

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+ ++

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+

+ + + + + +

+

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+

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+ +

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+ ++

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++

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+ + + +

+

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+

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+

+

+

+

+

+

+++++ + + +

+

+ +

+ +

+ ++ ++ +

+ + ++++

+

+

+

+

+

+

+

+

+ +

+

+

+

+ + + +

+ + ++ ++ ++

+

+

+

+

++

+

+ + +

+

+ + + +++

+ +

+ + + +

+

+

+ + ++ +

+

+

+ ++ +

+ + +

+ + + + + + +

+

+

+

+

+

+ + ++

+

+ +

+

+

+ + + + +

+ +

+

+ +

++++

++++ + + +

+ + + +

++

++ + +

+

++

+

+

+ ++ + +

+

+ +

+

+

++

+

+

+

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+ + +

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+

+

+

+ + + +++ + +++

+

+

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+ ++

+

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+ +++ + +

+

+ +

+

++ +

+

+ +

++

+ +

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++

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++ +

+

++

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+

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+ +

+ ++

+

+

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+ +

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+ +

+ +

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+

+ ++

+

+ + + +++

+ +

+

+ +

+ ++

+ ++ ++

+ +

++

+

+ ++

+ + +

+

+

+ + + +

+ + +

+ + +

+

+++

+++

+

+++

+

+

++

+

+++ + ++ + +

+

ZV

SAF

SAR

CAL

-15 -7 -4 -1 2 5 8 15

18

12

6

0

1.75

++

+

+

+

+

+

+

+

+ +

+

+

+

+ +

+

+ ++

+

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+

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++++++

+ + ++

+

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+ ++

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++

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+

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+ ++ ++

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+

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+

+ +

+

+ + +++

SAFCAL

-15 -7 -4 -1 2 5 8 15

18

12

6

0

SAFCAL

LP

SV

SW NE

B

1.70 1.75 1.80 1.85 1.90

+ Background 

 Seismicity

+ Aftershocks

DEPTH, IN KILOMETERS

x-DISTANCE, IN KILOMETERS

y=–45 km

y=–25 km

y=–10 km

y=–3 km

y=4 km

y=10 km

y=18 km

y=26 km

y=35 km

y=50 km

y=69 km

Figure 6.—Continued

SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E176 EARTHQUAKE OCCURRENCE

sections (fig. 6). Given the resolution and grid spacing, however, this boundary could as well dip southwest. The resolution 

is fairly even over the length of the low-velocity body, and so 

we can be confident of its location centered near y=4 km.

Below 6-km depth, the three-dimensional model shows an 

overall higher P-wave velocity on the southwest side of the 

vertical projection of the San Andreas fault trace, as well as 

several distinct high-velocity bodies. One high-velocity body 

parallels the fault on the southwest, is centered at y=10 km, 

and ranges in P-wave velocity from 6.5 to 6.6 km/s. This 

high-velocity body is about 20 km long and dips southwest 

at 8- to 12-km depth. The resolution is fairly even in this 

vicinity, and so the high-velocity body is most likely a distinct 

feature, rather than reflecting the uniformly high velocity on 

0

5

10

15

–5 0 5 10 15

y=69 km 

–5 0 5 10 15

0

5

10

15

y=–45 km 

0

5

10

15

–5 0 5 10 15

y=–25 km 

0

5

10

15

–5 0 5 10 15

y=–10 km 

10

15

y=–3 km 

0

5

–5 0 5 10 15

0

5

10

15

–5 0 5 10 15

y=4 km 

0

5

10

15

–5 0 5 10 15

y=10 km 

0

5

10

15

–5 0 5 10 15

y=18 km 

0

5

10

15

–5 0 5 10 15

y=26 km 

1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5

Spread Function

0

5

10

15

–5 0 5 10 15

y=35 km 

0

5

10

15

–5 0 5 10 15

y=50 km 

SW NE

x-DISTANCE, IN KILOMETERS

DEPTH, IN KILOMETERS

Figure 7.— Resolution plots of VP /VS ratio for final three-dimensional model (fig. 6) of velocity grid in the 

Loma Prieta region (figs. 1, 2). Spread function is shown with contours of averaging vector (row of resolution matrix) for nodes with smearing greater than grid spacing. Contours, which are 70 percent of diagonal 

resolution, are plotted where resolution contour extends farther than adjacent nodes but not for nodes with 

near-zero resolution.
E177

the southwest side of the fault. The high-velocity body appears 

to be about 3 km thick, although the resolution decreases 

below 8-km depth. A synthetic test of a subset of these data 

was performed by Kissling and others (1994). Below 10-km 

depth, the high-velocity side of their inversion model had a 

10-km-thick high-velocity feature in addition to the expected 

overall high velocity. Thus, below 10-km depth, the southwestern high-velocity body could be part of a broader feature 

extending farther southwestward. For example, Brocher and 

others (this chapter) describe a high-velocity feature below 

10-km depth that extends 35 km southwestward to the San 

Gregorio fault.

On the northeast side of the San Andreas fault is a moderately high velocity zone (6.1–6.3 km/s) extending from 

Figure 8.— Comparison of three-dimensional inversion (upper plots), with even across-fault grid spacing and one-dimensional initial model, with progressive inversion of three-dimensional model (lower plots), showing P-wave velocity in grid layers at y=–25 km (A), y=–3 km 

(B), y=10 km (C), and y=35 km (D) and VP /VS ratio in grid layers at y=–3 km (E) and y=10 km (F). Dots, inversion grids. Contour intervals: 

P-wave velocity, 0.25 km/s; VP /VS ratio, 0.02. Faults: CAL, Calaveras; SAF, San Andreas; SAR, Sarmiento; ZV, Zayante-Vergeles.

SEISMOTECTONICS OF THE LOMA PRIETA REGION 

0

10

20

SW NE ZVSAF Cal

4

5

5

6

6

0

10

20

−20 −10 0 10 20 30

l l l

4

5 5

6 6

0

10

20

ZVSAFSarCal

5

5

6

6

0

10

20

−20 −10 0 10 20 30

l l l l

5

5

6

6

0

10

20

ZVSAFSarCal

4 5

5

6

6

0

10

20

−20 −10 0 10 20 30

l l l l 5

5

6

6

x-DISTANCE, IN KILOMETERS

DEPTH, IN KILOMETERS

A

B

C

y=–25 km

y=–3 km

y=10 km

0

10

20

y= 35 km

ZVSAFCal

5

5

6

6

0

10

20

-20 -10 0 10 20 30

l l l 4

5

5

6

6

0

10

20

y= –3 km

ZVSAFSarCal

0

10

20

-20 -10 0 10 20 30

l l l l

1.76

0

10

20

y= 10 km

ZVSAFSarCal

1.76

0

10

20

-20 -10 0 10 20 30

l l l l

1.76

1.72

1.72

1.72

1.76

1.80

1.72

1.82

1.76

1.72

1.76

1.78

x-DISTANCE, IN KILOMETERS

DEPTH, IN KILOMETERS

D

E

F
E178 EARTHQUAKE OCCURRENCE

y=−18 to y=20 km at 6- to 12-km depth. The resolution in 

this area is adequate, and some hypocenters occur within it, 

and so this feature is not caused simply by smearing from 

the southwestern high-velocity body. In the uppermost 4 

km of crust, shallow high-velocity features, with a P-wave 

velocity of 5.9 km/s at 2-km depth, follow the Sargent fault 

from y=10 to y=18 km; these features are well imaged at 

y=−3 km, where the shallow resolution is best.

In the southeastern part of our model (y=26–69 km), the 

velocity structure is simpler, with uniformly higher velocity 

on the southwest. Within the general southwestern highvelocity body is a distinct narrow high-velocity zone, with 

a P-wave velocity of 6.8 to 7.0 km/s at 6-km depth, that is 

15 km long and extends from y=35 to y=50 km. We note 

that the resolution is poor below 12-km depth, and so we are 

uncertain of the depth extent of this high-velocity zone.

The VP /VS model has a moderate resolution from y=−25 

to y=−3 km, and the most conspicuous feature is a high 

VP /VS ratio (max 1.89) within the large low-P-wave-velocity 

area southwest of the San Andreas fault. At y=−3 km, the 

high VP /VS ratio is centered on the fault, rather than in the 

low-velocity material. This feature would be intriguing if 

it were imaging some property of the fault zone itself, but 

it is just as likely to be a smearing of high VP /VS ratios on 

both sides of the fault. Apparently, the various high-velocity 

features have different VP /VS ratios: the high-velocity zone 

northeast of the fault has a high VP /VS ratio, and the highvelocity body on the southwest has a low VP /VS ratio.

As described above in the section entitled “Inversion 

Procedure,” we take into account the general pattern of 

geology and the uneven data distribution by using a series 

of inversions, with more closely spaced nodes near the 

San Andreas fault, where most events occur. To evaluate 

the effect of these assumptions on the resulting velocity 

features, we also compute an inversion with evenly spaced 

nodes in the across-fault direction and the one-dimensional 

initial crustal model (fig. 3A), without two-dimensional or 

coarse three-dimensional inversion steps. These results are 

compared with our progressive-inversion results in figure 8. 

No velocity features are evident in our progressive model 

that are not evident in the evenly spaced model; however, 

the features appear broader in the evenly spaced model 

because their spatial position cannot be defined so well by 

the coarser spacing near the San Andreas fault. The evenly 

spaced model also differs in having many single-point 

velocity anomalies and low-velocity zones in its poorly 

resolved parts. Only a couple of these differences appear 

to be reasonable, such as the more distinct low-velocity zone associated with the Calaveras fault at y=10 km 

(fig. 8C). In general, the differences do not appear to be 

believable features that could be interpreted to represent 

crustal structure. For example, the low-velocity zone on 

the southwest edge at y=−25 km (fig. 8A) and the velocity 

complexity below 10-km depth at y=10 km (fig. 8C) are 

more likely to be inversion artifacts. Therefore, the evenly 

spaced model would be less useful for other applications, 

such as providing three-dimensional rheology in numerical modeling.

RELOCATION OF EVENTS

To improve the image of the seismicity with respect to 

the new velocity model, we relocated 8,577 events with 

the final three-dimensional model and station corrections. 

These events, which occurred from 1968 through September 

8, 1993, were recorded by Calnet, have hand-timed arrival 

times, are all of M≥1.5, and were located in the central part 

of our model. The standard Calnet catalog uses four different one- and two-dimensional velocity models with station 

corrections to locate events in this area and achieves an 

average rms traveltime-residual misfit of 0.13 s. With the 

single three-dimensional velocity model, the average rms 

misfit drops by 25 percent to 0.096 s.

In the figures that follow, we plot only the best located 

of the new hypocenters. Each hypocentral location uses at 

least 15 stations and has an rms misfit of less than 0.12 s. 

In general, the hypocenter locations do not differ substantially from those in other studies (Dietz and Ellsworth, this 

chapter), but they do provide a self-consistent comparison 

between the hypocenter locations and the three-dimensional 

velocity structure. These hypocenters were also used in 

figures 5 and 6.

DETERMINATIONS OF

FAULT-PLANE SOLUTIONS

For the events located with the three-dimensional model, 

first-motion fault-plane solutions are determined with the 

program FPFIT (Reasenberg and Oppenheimer, 1985), 

using the three-dimensional raypaths. Requiring 25 or more 

hand-timed readings, and rejecting solutions that did not 

converge well or had multiple possible minimums, reduces 

the number of solutions to 3,964 events, of which 2,624 

are considered to be well determined because they have 

30 or more observations. The 90-percent-confidence limits 

in strike, dip, and rake were limited to 40°, 43°, and 50°, 

respectively; the observations were well distributed over the 

focal sphere, and the rms misfit was less than 0.24. These 

values were picked by making histograms of the various 

functions and then selecting values to keep most of the solutions while rejecting the tails of the distributions. A review 

of the solutions shows that these criteria are adequate. The 

average 90-percent-confidence limit on the three angles for 

the selected solutions is 23°.

The distribution of focal mechanisms is illustrated in 

figure 9 by maps of the different types of mechanisms in the 

background seismicity and in the aftershocks. The strike-slip 

mechanisms were divided into right and left lateral, signify-
E179

ing the type of slip on the nodal plane most nearly parallel 

to the San Andreas fault. Thus, the events plotted as right 

lateral are consistent with the stress field expected in the 

Loma Prieta region; however, an event plotted as right lateral 

could be left lateral on a plane subperpendicular to the San 

Andreas fault. The P-axes are shown for thrusting events, 

and the T-axes for normal events. Right-lateral and thrust 

solutions predominate, with a greater proportion of thrusting events in the north half of the aftershock zone, as might 

be expected from the main-shock slip distribution, which 

indicates a higher proportion of reverse slip for the northern 

part of the rupture (Beroza, 1991; Wald and others, 1991). 

Numerous left-lateral and normal aftershocks, however, are 

not inconsistent with the main-shock slip. The left-lateral 

and normal aftershocks are not on isolated structures but 

are distributed throughout areas of right-lateral and thrusting events, similar to the one-dimensional mechanisms of 

Oppenheimer (1990).

STRESS FIELD INFERRED

FROM FOCAL MECHANISMS

Focal-mechanism data were inverted by the method of 

Michael (1984) to obtain the uniform component of the stress 

field under the assumption that the constitutive relation is linear 

and isotropic. The uniform-stress tensor that best explains the 

focal mechanisms in a data set is expressed by the three axes 

A B

0 10 20 KILOMETERS

Right-lateral-event hypocenters Left-lateral-event hypocenters

Before main shock After main shock

Figure 9.—Distribution of focal mechanisms in the Loma Prieta region (fig. 1) before (upper plots) and after (lower plots) October 17, 1989, 

main shock, calculated using three-dimensional raypaths. A, Events with focal mechanisms consistent with right-lateral slip on a nodal plane 

subparallel to the San Andreas fault. B, Events with focal mechanisms consistent with left-lateral slip on a nodal plane subparallel to the San 

Andreas fault. C, P-axes of events with thrust mechanisms. D, T-axes of events with normal mechanisms. Note that because we categorize 

strike-slip events on the basis of a nodal plane subparallel to the San Andreas fault, events on cross-structures may be misclassified, for 

example, in figure 9A, where an intermediate fault links the Sargent and Calaveras faults.

SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E180 EARTHQUAKE OCCURRENCE

of principal stress and the stress ratio φ=(S2−S3)/(S1−S3), where 

S1, S2, and S3 are the principal stresses numbered from most to 

least compressive. The 95-percent-confidence regions were 

determined by the bootstrap method (Michael, 1987a), and an 

average misfit angle β (angle between observed and predicted 

slip directions) was computed that measures the difference 

between the observed slip direction and the predicted direction 

of maximum tangential traction. The angle β can be used as 

a measure of the success of inversion; if it exceeds 47°, then 

the inversion has failed to find a uniform-stress tensor, most 

likely because the true stress field is highly heterogeneous 

(Michael, 1991).

The data set is divided into two time periods, before and 

after the main shock, and four hypocentral volumes for a total 

of eight subsets, as was done for the preliminary Loma Prieta 

data set by Michael and others (1990). The four hypocentral 

volumes were obtained by first dividing the epicenters into 

three areas: the central area, which contains the main-shock 

rupture; and the areas to the north and south of this central 

area. The main-shock rupture did not extend to the surface, 

and so the central area was further subdivided at 5-km depth 

into shallow and deep volumes, with the deep volume containing the main-shock rupture. For each event, the nodal plane 

oriented closer to a strike of northwest-southeast was selected 

to represent the slip plane. Owing to the complexity of these 

mechanisms, however, when computing the confidence limits, 

the odds of choosing the wrong fault plane from the two nodal 

planes were estimated at 50 percent (Michael, 1987b).

The major coseismic change is found in the main-shock 

rupture area (deep central area, fig 10). Before the main 

shock, the motion on fault planes subparallel to the San 

Andreas fault was primarily right-lateral strike slip due to a 

uniform stress field with the axis of principal stress S1 oriented 

slightly east of north. The aftershocks in the deep central 

area, however, are a completely heterogeneous set of focal 

mechanisms that cannot be explained by a uniform-stress 

0 10 20 KILOMETERS

C D

Thrusting-event P-axes Normal-event T-axes

After main shock Before main shock

Figure 9.—Continued
E181

tensor, as evidenced by the extremely high misfit, (β=61°). 

These results from the three-dimensional focal mechanisms 

are similar to those of the initial study by Michael and others 

(1990), and thus continue to support their conclusion that the 

stress field in the rupture zone became highly heterogeneous 

after the main shock.

Above the main-shock rupture zone (shallow central area, 

fig. 10), the number of data is insufficient to determine the 

previous stress field, but aftershocks show a stress field with 

the axis of principal stress S1 oriented east-northeast and the 

axis of principal stress S3 vertical, consistent with thrusting. 

The northern area has a similar absence of background seismicity, but the aftershocks consisted primarily of fault-parallel 

thrusts due to fault-normal compression. In the southern area, 

right-lateral strike-slip events dominate both time periods, 

and the stress regime shows approximately north-southward 

compression, with the axis of principal stress S2 vertical both 

before and after the main shock. Michael and others (1990) 

suggested that the main shock increased the stress in the southern area, leading to a lower β value during the aftershocks, 

but this conclusion is not supported by the three-dimensional 

mechanisms.

The high β value of 61° in the main-shock rupture zone 

suggests that the uniform component of the stress field, 

which is what the stress inversions attempt to recover, is 

smaller than the spatially varying component, which the 

inversions treat as noise. Although some nonzero uniform 

component to the stress field is likely, we cannot recover 

it within the 95-percent-confidence limits of the inversion 

results (Michael, 1991). Zoback and Beroza’s (1993) conclusion that the total stress drop on the main-shock fault plane 

(Michael and others, 1990) will lead to a uniform component 

After main shock

3

Very

heterogeneous

stress

Before main shock

North

Central

South

Shallow Deep Shallow Deep

Insufficent

data N=2

Insufficent

data N=4

N=68, β=31o

N=333, β=18o

N=49, β=17o

N=75, β=18o N=402, β=61o

N=641, β=19o

1

2

2

3

1

2

3

1

2

3

1

2 3

20 km

North

Central

South

Selection

Areas

Deep &

Shallow

1

Figure 10.—Results of stress inversion before and after October 17, 1989, main shock for four areas originally defined by Michael and others (1990): north of main-shock rupture (north), south of main-shock rupture (south), above main-shock rupture (shallow central), and near 

main-shock rupture (deep central). 1–3, three axes of principal stress from most compressional to most extensional, with their 95-percentconfidence areas; N, number of inverted events; b, average misfit angle between observed and predicted slip directions.

SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E182 EARTHQUAKE OCCURRENCE

of fault-normal compression during the aftershocks is reasonable, but we consider this conclusion to be unresolvable 

with the data. 

DISCUSSION

The aftershocks show a fault dipping southwest to near 

the base of the crust, but could there also be a subvertical 

fault that extends below the surface trace of the San Andreas 

fault? The aftershock zone would intersect a subvertical San 

Andreas fault at 5- to 6-km depth. Above that hypothetical 

intersection, the San Andreas fault is indicated by a subvertical lateral velocity contrast in the 0- to 4-km grids. In the 

northwestern and central parts of our model, from y=−45 to 

y=18 km, the P-wave velocity in the higher shallow material is on the northeast side; and in the southeastern part of 

our model, from y=26 to y=69 km, the P-wave velocity is 

higher on the southwest side (fig. 5A). With the station array 

and velocity grid we used, we are unable to discern whether 

or not an additional narrow low-velocity zone is present at 

the fault.

Does the San Andreas fault extend below 6-km depth? In 

the southeastern part of our model, the fault is defined by a 

lateral-velocity contrast down to 15-km depth, with a high 

P-wave velocity on the southwest (y=35–50 km, fig. 6). The 

resolution plots (y=35–50 km, fig. 7) show that whereas the 

spread function increases below 10-km depth, the nodes 

near the San Andreas fault are not dominated by smearing 

from other areas. Therefore, they should show a weaker, 

broader fault image than the actual fault, but not be artifacts. 

A synthetic example of a simple fault model with a subset 

of the event data set showed that a velocity gradient across 

the fault could be determined down to 15-km depth but that 

it was much weaker than in the true model (Kissling and 

others, 1994). In the density model of Jachens and Griscom 

(this chapter), a narrow, high-density (2.91 g/cm3) fragment 

adjacent to the San Andreas fault broadens with depth into 

a large, 5-km-thick high-density body (fig. 11C) that corresponds to the large high-velocity zone we image on the 

southwest side (figs. 5, 6). They suggested that the shallow 

high-density fragment, a mapped gabbro, is a small piece that 

has been splintered off the larger deeper body.

Northwest of y=18 km, below 6-km depth, the San Andreas 

fault is not simply defined in the P-wave-velocity image. 

A relatively high velocity is measured on both sides of a 

subvertical San Andreas fault, and several distinct highvelocity bodies are present. The VP /VS ratio varies across 

the fault, implying the high-velocity body is not a single 

feature: the high P-wave velocity on the northeast side has a 

high VP /VS ratio, but the deeper high P-wave velocity on the 

southwest side has a low VP /VS ratio. Additionally, Jachens 

and Griscom’s (this chapter) density model near y=−3 km 

shows that the high-velocity features have different densities 

3 4 5 6 7

+ Background Seismicity

+ Aftershocks

1.70 1.75 1.80 1.85 1.90

+ Background Seismicity

+ Aftershocks

NW SE NW SE

DEPTH, IN KILOMETERS

y-DISTANCE, IN KILOMETERS

A B

P-WAVE VELOCITY (VP) IN VP/VS RATIO

KILOMETERS PER SECOND

-40 -20 0 20

-20 -15 -10 -5 0

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-40 -20 0 20

-20 -15 -10 -5 0

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-40 -20 0 20

-20 -15 -10 -5 0

1.8

1.8

1.75

1.75 1.75 1.85

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+++++++++

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-40 -20 0 20

-20 -15 -10 -5 0

1.8

1.8 1.8

1.75

1.75

1.75

1.75 1.85

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VP at x= –1 VP at x= 2

VP /VS at x= −1 VP/ VS at x= 2

Figure 11.—P-wave-velocity (upper plots) and VP /VS ratio (lower plots) cross sections for grid layers at x=–1 km (A) and x=2 km (B), 

showing hypocenters within 3 km of each cross section. Contour intervals: P-wave velocity, 0.5 km/s, with 4.0-, 5.0-, 6.0-, and 7.0-km/s 

contours in bold; VP /VS ratio, 0.05, with 1.7, 1.8, and 1.9 contours in bold. Irregular white outlines, 1- and 2-m-slip contours of Wald and 

others (1991).
E183

4 4

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-15 -7 -4 -1 2 5 8 15

18

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ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

12

6

0

ZV

SAF

SAR

-15 -7 -4 -1 2 5 8 15

12

6

0

ZV

SAF

SAR

CAL

2.37

2.48

2.52

2.56

2.59

2.67 2.91

2.5952.77

2.729 Edge of Model

2.87

2.17

2.67

2.67

gr Ts

gb

Jou

Jou Kjf

Ts, Ks, Kjf

2.17

2.87

2.91

2.642.612.55

2.48

2.37 2.17 2.87

gr

gb

Ts Ts & Qal

Ts & Qal

2.67

2.67

2.67

4

5

6

7

-15 -7 -4 -1 2 5 8 15

18

12

6

0

+++++ ++

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ZV

SAF

SAR

CAL

NNW SSE

0 10 20 30 40 50 60

0 -15 -5

4 5

6

6

DISTANCE SUBPARALLEL TO y, IN KILOMETERS

10

6.4

6.4

A

B D

E

C

x-DISTANCE, IN KILOMETERS

DEPTH, IN KILOMETERS

SW Density at y=−3 km Density at y=35 km NE

VP at y=−3 km VP at y=35 km

VP along HVB

Figure 12.— Comparison of density models of Jachens and Griscom (this chapter) with 

three-dimensional P-wave-velocity (VP) model, showing density (A, C; all values in grams 

per cubic centimeter) and P-wave velocity (B, D; all values in kilometers per second) for 

grid layers at y=–3 and 35 km. P-wave-velocity contour interval, 0.5 km/s, with 4.0-, 

5.0-, 6.0-, and 7.0-km/s contours in bold. Faults: CAL, Calaveras; SAF, San Andreas; 

SAR, Sarmiento; ZV, Zayante-Vergeles. gb, gabbro; gr, granitic and older metamorphic 

rocks; Jou, Coast Range ophiolite; Kjf, Jurassic and Cretaceous Franciscan complex; Ks, 

Cretaceous sedimentary rocks; Qal, Quaternary alluvial deposits; Ts, Tertiary sedimentary rocks. E, Cross section along length of main high-velocity body (HVB; see fig. 5C), 

which trends slightly more northerly than the San Andreas fault. Cross section is subparallel to y-axis, starting at origin and ending at y=–9 and y=59 km. P-wave-velocity contour 

interval, 0.5 km/s, with 4.0-, 5.0-, 6.0-, and 7.0-km/s contours in bold; 6.4-km/s P-wavevelocity contour is added to better delineate high-velocity body. Dashed contours emphasize 

lack of resolution at depth and therefore our inability to delineate bottom of high-velocity 

body.

SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E184 EARTHQUAKE OCCURRENCE

across the fault and that the deeper southwestern feature has 

a higher density (fig. 11A).

Thus, the high-velocity slivers adjacent to the San Andreas 

fault on the northeast at 6- to 8-km depth and on the southwest 

at 8- to 15-km depth have different material properties and 

so are different rock types separated by the fault and are not 

a single feature that crosses it. Other studies that developed 

only a P-wave-velocity model interpreted the high-velocity 

body as a single body that would cross a subvertical San 

Andreas fault (Foxall and others, 1993; see Shalev and Lees, 

this chapter).

We infer that the most likely boundary of the southwestern high-velocity body is the San Andreas fault extending 

from the surface throughout the seismogenic zone. A few 

aftershocks and previous microearthquakes are consistent 

with the San Andreas fault (as also noted by Olson, 1990). 

We cannot determine the fault dip accurately with our grid 

spacing and resolution, and so the fault may be either subvertical or steeply southwest dipping in the Loma Prieta region. 

We are confident, however, the San Andreas fault is distinct 

from the Loma Prieta rupture surface and that it does not dip 

northeast. The fault surfaces can be visualized from aftershock 

locations. We cannot define them exactly, and so we prefer 

not to draw the faults on every velocity cross section, but for 

illustration purposes the two faults are sketched in the y=4 

km section (fig. 6).

In addition to a subvertical San Andreas fault, other faults 

are also imaged in the velocity model. A southwest-dipping 

thrust fault is correlated with the southwestern Santa Clara 

Valley thrust belt that has emplaced the Sierra Azul block (see 

McLaughlin and Clark, this chapter). This thrust fault is not 

defined by a lineation of seismicity but seems to bound the 

seismicity such that aftershocks in this area occur primarily in 

the hanging-wall material (y=−3 km, fig. 6). Most of the aftershocks within this volume have thrust mechanisms (fig. 9B).

No single clear feature in the velocity model reveals the fault 

segment that ruptured in the 1989 Loma Prieta earthquake 

(fig. 6), although small velocity variations might be missed 

below 10-km depth as the spread function increases (fig. 7). 

Along the segment from y=−10 to y=−3 km, the rupture zone 

has the appearance of a thrust, with higher velocities in the 

upthrown block. Farther southeast from y=4 to y=10 km, the 

plane defined by aftershocks is centered in the high-velocity block, not below it. The Loma Prieta fault segment thus 

appears to vary in its characteristics in the velocity image 

along the length of rupture. The slip also varies: the earthquake 

was an oblique-slip event, with at least two patches of slip 

with different proportions of reverse- and strike-slip motion 

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

Michael and Eberhart-Phillips (1991) showed that geologic 

structure as imaged in three-dimensional velocity models 

is related to seismogenic behavior, on the basis of models 

of four earthquakes, including preliminary results from the 

1989 Loma Prieta earthquake. In the asperities with greatest 

moment release, higher-velocity material is in contact with 

the fault, and the velocity image of the fault is more complex than in creeping sections, where the fault is imaged as 

a simple strong across-fault velocity contrast. Although our 

model is coarse at the southeast end, other studies focusing 

on the creeping section also image the fault with a relatively 

uniform velocity contrast. For example, Dorbath and others 

(1996), using a uniform 5-km grid spacing in their study of 

the intersection of the Calaveras and San Andreas faults, also 

characterized the San Andreas fault by an across-fault velocity 

contrast at y=26 km and farther southeast.

In our final model (fig. 11), we can consider details of the 

high-velocity body and the Loma Prieta rupture zone. The 

southeastern rupture patch is correlated with a large, distinct 

high-velocity body within the general southwestern high-velocity feature. The deepest part of the high-slip patch is correlated 

with the deepest part of the high-velocity body, although the 

resolution is weak below 10-km depth and the high-velocity 

body may extend to 18-km depth all along its length to the 

southeast. The high-velocity body is a more nearly linear feature 

than the Loma Prieta rupture (figs. 5, 12E). It trends slightly 

more northerly than the San Andreas fault and is closest to the 

fault near the main shock. Southeast of the main shock, aftershock locations show that the Loma Prieta rupture bends toward 

the subvertical San Andreas fault (fig. 6). Thus, the rupture is 

nonplanar and consists of at least two different fracture surfaces. The high-velocity body cannot be an undeformed block 

with earthquake displacements only on the block boundaries. 

Instead, we propose that the high-velocity body adjoining the 

San Andreas fault at depth has many internal fracture surfaces 

that could sustain rupture in different earthquakes. The highvelocity body may have been built up of many fragments over 

time coming off the mafic oceanic crust below, or it may have 

been emplaced as a large coherent fragment that developed 

internal structure as it was transported into the San Andreas 

fault bend in the Santa Cruz Mountains.

The primary feature in our final model is a volume of 

higher-than-normal VP /VS ratio in the hanging wall of the 

Loma Prieta rupture zone at the northwest end of the mainshock rupture. This feature, also seen by Thurber and others 

(1995), is associated with low P-wave velocity and low 

density. The low-velocity material on the southwest corresponds to a 9-km-wide, 8-km-deep low-density (2.37–2.67 

g/cm3) body (figs. 12A, 12B). The horizontal base of the 

high-VP /VS -ratio zone at about 8-km depth may be a lithologic 

boundary between marine sedimentary rocks above and 

gabbroic basement below (see McLaughlin and Clark, this 

chapter). Because high VP /VS ratios correlate with high pore 

pressures (Nur and Simmons, 1969), this correlation may 

provide an explanation for the northwest termination of the 

Loma Prieta asperity. High pore pressure in that volume of 

material could reduce the effective pressure, possibly causing 

velocity-strengthening behavior. Thus, as the rupture began to 

propagate upward into the sedimentary rocks, it could have 

undergone a change in frictional behavior and had difficulty 

propagating farther.
E185

In light of the aftershock locations and velocity features, 

we suggest that the fault surface which ruptured in the 1989 

Loma Prieta earthquake should not be considered as the only 

active fault segment, or as a tectonic fault independent of the 

subvertical San Andreas fault; instead, the approximately 5-

km-wide zone could be considered as the active San Andreas 

fault zone. Some events may rupture the subvertical San 

Andreas fault, which bounds the southwestern high-velocity 

mafic fragment. This high-velocity mafic fragment probably contains various fracture surfaces, any of which might 

rupture in an earthquake that would release the strain built 

up on the San Andreas fault zone. For any particular event, 

the exact rupture surface may depend on such factors as the 

residual-stress pattern after previous events. This model of 

the San Andreas fault zone explains how the Loma Prieta 

rupture could fill in a microseismicity gap that had been 

observed on the subvertical San Andreas fault (Moths and 

others, 1981; U.S. Geological Survey staff, 1990). Given 

the various fault structures, however, this section of the San 

Andreas fault zone does not consistently rupture in a single 

type of characteristic earthquake but rather in varied events, 

such as the 1906 and 1989 earthquakes (for example, Segall 

and Lisowski, 1990).

CONCLUSIONS

Defining the spatial extent of lithologic units and materialproperty changes in the subsurface is a challenge best met 

by combining several geophysical parameters. In this study, 

we have developed new three-dimensional models for the 

Loma Prieta region. Inverting for VP /VS ratios circumvents 

the problems associated with determining VP /VS ratios from 

separate P- and S-wave-velocity models that have different 

resolution patterns.

To reasonably estimate velocities throughout the model, 

taking into account the overall pattern of mapped geology, 

we performed a series of inversions of progressively increasing complexity. An alternative model with even grid spacing 

across the fault, and a one-dimensional initial model, show 

similar but broader velocity features in the aftershock zone 

but also many single-point velocity anomalies and low-velocity zones in poorly resolved areas.

We interpret subvertical San Andreas fault through the 

Loma Prieta region from an examination of P-wave velocities, VP /VS ratios, microearthquake hypocenters, and densities. In the section southeast of the Loma Prieta rupture, the 

fault is simply defined by an across-fault P-wave-velocity 

gradient with higher velocity on the southwest at all depths. 

The velocity image is more complex near the Loma Prieta 

rupture, with low velocities on the southwest side above 7-

km depth and relatively high velocities on both sides of the 

fault below that depth.

Relocated aftershocks help define fault surfaces and show 

that the Loma Prieta rupture is associated with a high-velocity body, likely to be a mafic fragment, although the rupture 

surface does not bound this high-velocity body.

Fault-plane solutions computed with three-dimensional 

raypaths show a range of focal mechanisms similar to that in 

previous one-dimensional fault-plane solutions. Inversions 

of focal mechanisms for uniform-stress tensors in subregions 

are also similar to those in previous one-dimensional results. 

The aftershocks in the main-shock rupture zone cannot be 

fitted by a uniform-stress tensor and provide evidence for 

a coseismic-stress change resulting from a near-total stress 

drop during the main shock that left a small heterogeneous 

residual-stress field.

The high-velocity fragment appears to be a key feature in 

controlling the pattern of slip and the length of the fault segment that ruptured, as well as the distribution of background 

seismicity on the subvertical San Andreas fault. Thus, the 

fault surface that ruptured in the 1989 Loma Prieta earthquake, and the subvertical San Andreas fault, do not behave 

independently and could both be considered parts of the active 

San Andreas fault zone.

ACKNOWLEDGMENTS

Use of all the data collected after the earthquake turned 

out to be a large task, and we understand why so many 

data are commonly overlooked when excellent Calnet data 

are also available and already analyzed. Cliff Thurber and 

Shashank Atre combined forces and helped with this task 

by timing some of the data from portable instruments. We 

are also indebted to Bill Foxall for the UCB data, Chuck 

Mueller for the GEOS data, Gary Fuis for data from seismic 

cassette recorders, John Evans for the 5-day data from the 

teleseismic network, Cliff Thurber and Shashank Atre for 

the PASSCAL data, Tom Brocher for the onshore-offshore 

profiles, and Walter Mooney for his seismic-refraction 

profiles. Still, most of the data came from the USGS Calnet 

team, who completely timed and analyzed the records from 

the prolific aftershock sequence. Without the permanent 

network, our view of the Loma Prieta region would be hazy 

and grossly limited in time. Jim Luetgert, Bruce Julian, and 

John Evans all provided software to help analyze these data 

sets. The 1991 field experiment required too many personnel 

from both the USGS and Stanford University to name. We 

enjoyed many useful discussions with Greg Beroza, Tom 

Brocher, Rufus Catchings, Bill Ellsworth, Lynn Dietz, John 

Gephart, Bob Jachens, Will Kohler, Al Lindh, Bob McLaughlin, Dave Oppenheimer, Dal Stanley, Cliff Thurber, Mark 

Zoback, and Mary Lou Zoback. We thank Tom Brocher, Dave 

Oppenheimer, and Stephanie Ross for their reviews of the 

manuscript. We dedicate this paper to Ben Page, who, with 

his keen insight and unflagging enthusiasm for investigating the tectonics of the San Andreas fault system, pushed 

us to make the most geologically relevant interpretations of 

seismic-velocity inversions.

SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E186 EARTHQUAKE OCCURRENCE

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SEISMOTECTONICS OF THE LOMA PRIETA REGION 
E188 EARTHQUAKE OCCURRENCE
E189

THE LOMA PRIETA, CALIFORNIA, EARTHQUAKE OF OCTOBER 17, 1989—

GEOLOGIC SETTING AND CRUSTAL STRUCTURE

EARTHQUAKE OCCURRENCE

TELESEISMIC TOMOGRAPHY OF THE LOMA PRIETA REGION:

IMPLICATIONS FOR STRAIN PARTITIONING

By Yoko Takauchi and John R. Evans, U.S. Geological Survey

CONTENTS

————— Page

Abstract------------------------------------------------------------------------- 189

Introduction-------------------------------------------------------------------- 189

Method-------------------------------------------------------------------------- 191

 Data----------------------------------------------------------------------- 191

 Inversion technique------------------------------------------------------ 191

 Crustal correction------------------------------------------------------- 194

Results-------------------------------------------------------------------------- 195

 Unstripped models------------------------------------------------------ 195

 Stripped models: Upper-crustal structure---------------------------- 196

 Resolution---------------------------------------------------------------- 197

Discussion---------------------------------------------------------------------- 199

Conclusions-------------------------------------------------------------------- 202

Acknowledgments------------------------------------------------------------- 203

References cited--------------------------------------------------------------- 203

—————

ABSTRACT

From teleseismic traveltime residuals, we derive threedimensional velocity models of the Loma Prieta region to 

71-km depth. We use an extended ACH inversion process that 

minimizes parametrization artifacts while retaining an explicit 

resolution matrix. We also “strip” upper-crustal structure 

from our data, inverting the remaining teleseismic traveltime residuals to test the robustness of the deeper structures 

we image. For stripping, we use three methods: wave-front 

tracing through each of two local-earthquake-tomography 

models, and an iterative technique that maximally removes 

the upper-crustal signature from our data. Absent stripping, 

the shallow crustal structure we determine generally agrees 

with local-earthquake-tomography results, with discrepancies 

consistent with resolution effects and wave effects at strong 

velocity contrasts across some faults. Deeper structures are 

not significantly affected by crustal stripping, indicating that 

they are not much affected by smearing of shallow, largemagnitude velocity anomalies. We observe many horizontal 

velocity gradients beneath the San Andreas fault at all depths, 

including in the upper mantle. These gradients suggest that the 

fault was a deep, narrow shear locus at least at one time. Low 

velocities are observed beneath the Loma Prieta highlands 

near the Moho; this feature probably is caused by a crustal 

root, approximately 2 to 5 km thick and 10 to 30 km across. 

We also resolve two conspicuous upper-mantle low-velocity 

features between 45- and 70-km depth, separated right laterally 

along the San Andreas fault by about 45 km. The San Andreas 

fault (or its adjacent predecessor, the Pilarcitos fault) formed 

in this region by at least 6 Ma. Estimated slip rates along 

the San Andreas fault system suggest that this 45-km offset 

is comparable to the amount of shear occurring within just 

a few million years. If this pair of upper-mantle low-velocity features preserves the total upper-mantle slip at a locus 

beneath the San Andreas fault, this result is consistent with a 

model that posits rapid cooling of this volume after passage of 

the Mendocino Triple Junction and a current, broad locus of 

upper-mantle shear centered beneath the Hayward/Calaveras 

fault system to the northeast.

INTRODUCTION

The lithospheric structure of the Loma Prieta region of 

central California (fig. 1) is complicated by major and minor 

reverse- and strike-slip faults that together form the San 

Andreas fault system. Although the San Andreas fault is a 

manifestation of the northwest-striking boundary between the 

Pacific and North American plates at the surface, the deeper 

structure and position of the plate boundary are matters of 

debate.

The 1989 Loma Prieta earthquake (MW=6.9) was unusual 

for its relatively great depth and oblique reverse rupture. 

Furlong and Langston (1990) concluded that relatively low 

observed heat flow and thin upper crust in the region lead to 

strain accumulation at the base of the brittle crust (18–20-km 

depth), where the Loma Prieta rupture commenced. Below 

this depth, the lowermost Pacific crust can deform ductilely. 

They attributed this strain to a discrepancy in the displacement 

directions between the crustal and mantle parts of the Pacific 

plate where the crust moves relatively westward around a bend 

in the San Andreas fault, creating a basal traction toward the 

fault. If, as they proposed, the crust between sections of the 

fault is detached from the mantle somewhere in the lower 

crust, then crustal compressive strain will accumulate near 

the fault.

Zandt (1981) did the first comprehensive teleseismic-tomography study of central California, achieving 10- to 25-km 

E189
E190 EARTHQUAKE OCCURRENCE

ELEVATION, IN METERS

0 160 400 640 880 1040 1200

122°30' 122°00' 121°30'

37°

30'

37°

00'

KILOMETERS

0 10 20 30 40 50 60 70

A

B

B'

C

C'

D

D'

E

E'

F

G

G'

H

H'

HOSGRI FAULT

PACIFIC OCEAN

Monterey Bay

CALAVERAS FAULT

SAN

HAYWARD FAULT

San Francisco Bay

Santa Clara Valley

ZAYANTE /

FAULT

VERGELES

ANDREAS

FAULT

SARGENT FAULT

SAN GREGORIO -

A'

F'

Figure 1.—Loma Prieta region, Calif., showing locations of major faults and cross sections A–A' through H–H'. Dashed-line cross 

sections were shown only by Takauchi (1994). Both types of upright triangles, seismic stations used for principal-model series; both 

types of solid triangles, seismic stations used for detailed-model series.
E191

horizontal resolution from teleseismic P-wave-traveltime 

data recorded on the Northern California Seismic Network 

(NCSN) operated by the U.S. Geological Survey (USGS). 

His main findings include (1) correlation of a Tertiary marine 

sedimentary block between the San Andreas and Zayante 

faults with a modeled upper-crustal low-velocity feature, 

and (2) a low-velocity body in the lower crust (10–30-km 

depth) beneath Loma Prieta that he attributed to variation in 

the depth to the Moho.

Local-earthquake-tomography models by Eberhart-Phillips 

and others (1990) and Lees (1990) similarly reveal a lowvelocity wedge between the San Andreas and Zayante faults, 

extending to 8-km depth, that they interpreted as Tertiary 

marine sedimentary rocks. It also has a low Q value (Lees 

and Lindley, 1994). Both Lees (1990) and Eberhart-Phillips 

and others (1990) imaged a high-velocity body northeast of 

the San Andreas fault. In the near-surface, Eberhart-Phillips 

and others (1990) modeled it between the San Andreas and 

Sargent faults, whereas Lees placed it mostly northeast of the 

Sargent fault. This high-velocity body dips southwest to at 

least 18-km depth, where it underlies the previous low-velocity 

body in both models and includes the hypocenter of the 1989 

Loma Prieta earthquake. Lees inferred a relation between this 

high-velocity body and aftershock locations. Foxall and others 

(1993), who also derived a local-earthquake-tomography 

model of the southern Santa Cruz Mountains, inferred that 

the high-velocity body is the site of a stress concentration 

which caused the earthquake. Seismic-refraction studies (for 

example, Fuis and Mooney, 1990) also reveal the structure of 

the crust in this area, showing it thinning toward the Pacific 

Ocean and about 25 km thick beneath the Loma Prieta region 

(fig. 1).

Although previous local-earthquake-tomography models 

are useful for understanding the structure of the seismogenic 

crust, they do not reveal the lower-crustal or upper-mantle 

structure of the region because rays observed from local 

earthquakes do not penetrate that deeply. Previous teleseismictomography studies were of relatively low resolution, limited 

by the spacing of NCSN permanent stations. In this paper, 

we describe teleseismic-tomography models that resolve this 

deeper structure and allow inferences regarding the tectonics 

of the San Andreas fault system.

METHOD

DATA

Teleseisms used in our study were recorded from August 

to December 1990, using 27 temporary seismic stations to 

augment NCSN stations. The combination is a much denser 

array than that available to Zandt (1981). In this paper, we 

treat this array as two subarrays, with many stations shared 

between subarrays (fig. 1). The first subarray was for a study 

of the whole area to 60- to 70-km depth at ~6-km horizontal 

resolution. This less dense subarray comprised 20 temporary 

stations and 17 NCSN stations; the models derived from it 

are called the principal-model series. The second subarray 

was for a study in the vicinity of the intersection of the San 

Andreas and Sargent faults, where structure is particularly 

complex. This subarray comprised 14 temporary stations 

and 10 NCSN stations; the models derived from it, which are 

called the detailed-model series, reach 30- to 40-km depth at 

~3-km horizontal resolution. (A larger, regional-model series 

spanning the principal-model series, presented by Takauchi, 

1994, does not add much to the principal models and those by 

Zandt, 1981, and so is not described further here.)

The epicenters of the teleseisms we used are mapped in 

figure 2. The resulting distribution of rays within our target 

volumes is reasonably homogeneous, although events with 

northeastward backazimuths were far fewer than those with 

other backazimuths, as is typical in the Western United States 

(for P phases: 4 events from the northeast, 25 from the southeast, 11 from the southwest, and 23 from the northwest; for 

PKPdf phases: 5 events).

Relative arrival times of compressional P and PKPdf

phases were picked from filtered traces by a visual correlative 

method. Each pick was assigned a subjective quality based 

on noise and signal coherence, using the guideline that 3r

for “a”-quality picks should be about ±0.05 s, for “b”-quality 

picks about ±0.10 s, and for “c”-quality picks about ±0.25 s. 

We also reviewed the data in slowness-backazimuth space to 

identify errors, basing any corrective actions on the seismograms themselves. The final data set has 2,705 picks from 68 

teleseisms, including 544 “a”-quality, 1,193 “b”-quality, and 

968 “c”-quality picks.

INVERSION TECHNIQUE

Traveltime residuals Rij at the ith station for the jth event 

are calculated with respect to a known one-dimensional 

Earth model (Herrin, 1968). This “raw” residual includes the 

effects of origin-time uncertainty, hypocenter mislocation, 

and velocity heterogeneity outside the modeled volume. To 

mitigate these effects, relative residuals RRij are calculated by 

subtracting some reference residual Rj:

For Rj, we use a weighted mean residual

where wij is a weight derived from the pick qualities.

To invert these data for velocity structure, we used the 

ACH inversion method of Aki and others (1977) as extended 

by Evans and Achauer (1993) to include (1) an offset-andRRij =Rij −Rj.

TELESEISMIC TOMOGRAPHY OF THE LOMA PRIETA REGION: IMPLICATIONS FOR STRAIN PARTITIONING

Rj =

wij Rij ∑i

∑wij

i
E192 EARTHQUAKE OCCURRENCE

average horizontal-smoothing technique; (2) a layer-thinning 

vertical-smoothing technique; and (3) representation of the 

resolution matrix R in terms of “volume metrics,” irregular 

volumes of approximately the size, shape, and orientation 

of the smallest resolvable features. We use the more mature 

ACH method because its behavior is well understood and it 

produces an explicit resolution matrix R.

The target volume beneath the receiver array is divided into 

layers of approximately equal teleseismic traveltime, where 

each layer is subdivided into rectilinear blocks whose horizontal width is about equal to the mean station spacing. This 

dimension is the smallest horizontal size validly modeled by 

ACH tomography. Each block has an unknown fractional slowness deviation from the initial model’s (layer) slowness. The 

total thickness of the modeled volume approximately equals 

the aperture of the receiver array because below this depth, 

P-wave rays diverge significantly, no longer crisscrossing, so 

that deeper objects are poorly resolvable.

Succinctly, following Evans and Achauer (1993), we linearize 

the problem through block parametrization, disregarding refraction by the slowness perturbations (a second-order term because 

our rays do not have turning points within the model):

b = Gm,

where b is a vector derived from the traveltime residuals 

RRij, m is a vector containing the unknown fractional slowness deviations of each block (that is, the velocity model we 

seek), and G is a matrix derived from the unperturbed traveltimes of ray ij in block k. (Here, we leave out details related 

to data-quality weighting and parameter separation for the 

source term.) To estimate m, we use damped least squares to 

regularize the problem:

mˆ = (G + θI)−1b.

The solution, mˆ, is not iterated, since refraction is a secondorder effect. More modern methods use higher-order regularization, which is arguably more physical, but the smoothing 

procedures that we next apply obviate this advantage.

We chose an appropriate damping value by plotting tradeoff 

curves between damping, data fit, and model complexity (the 

Euclidean norm of mˆ). The optimal damping value is near 

a residual data variance that is comparable to the square of 

expected reading errors, and where neither data fit nor model 

complexity dominates the solution (fig. 3). Finally, the Moho 

is near 25-km depth in this region (for example, Fuis and 

Mooney, 1990). To improve ray-tracing accuracy, we ensure 

that a layer boundary falls at this depth. Horizontal block 

widths are ~6 and ~3 km for principal- and detailed-model 

series, respectively.

The processes of horizontal and vertical smoothing proceed 

in several steps, beginning with a “base” model and proceeding through layer-thinning and offset-and-average steps (as 

explained below). For the base model in each series, we chose 

layer thicknesses about 1.5 times the block width, a choice 

designed to yield fairly “well resolved” individual blocks. For 

initial crustal velocities, we used the average of the velocity 

models of Dietz and Ellsworth (this chapter) for the southwest 

and northeast sides of the San Andreas fault. Mantle velocities are those of Herrin (1968), joined to the crustal model 

by a linear gradient between 25- and 45-km depth. To this 

velocity structure we applied an Earth-flattening transform to 

trace rays more accurately. (The ACH inversion that we use 

is Cartesian.) Our choice of damping value was for this base 

model. We also used the base models to compute diagonal 

Figure 2.—Azimuthal equidistant maps of earthquake epicenters for 

this study (dots). Diamond in upper map is study area; diamond in 

lower map is its antipode. Dark circles in upper map are 025° and 

100° distant from center of study area; dark circle in lower map is 

110° distant from study area (thus, 070° distant from antipode).
E193

elements of the resolution matrix R and associated standard 

errors, because subsequent layer thinning invalidates these 

items (Evans and Achauer, 1993).

We incorporated vertical smoothing into the model by dividing each layer of the base model, except for the shallowest 

layer, into three sublayers of equal traveltime (tables 1, 2). 

Evans and Achauer (1993) showed that this layer-thinning 

method is a legitimate means of obtaining vertical smoothing in ACH teleseismic tomography, as long as the damping 

value is reduced in proportion to the increase in the number 

of unknowns (that is, in the number of modeled blocks).

One effect of layer thinning is to drastically reduce the 

values of the diagonal elements of the resolution matrix R. 

Although these diagonal elements are in no case an adequate 

exposition of resolution characteristics, with layer thinning 

that inadequacy is manifold, and so we clearly need a more 

complete description. For this purpose, Evans and Achauer 

(1993) introduced the “volume metric,” which is a figure of 

merit derived from any single column of the resolution matrix

R by summing its largest positive values until they reach some 

acceptable “equivalent diagonal-element value,” for example, 

0.75. The volume metric is the irregular portion of the model 

spanned by these summed blocks. Thus, each volume metric 

is a simple representation of the size, shape, and orientation 

of the model’s “impulse response” at the site of the diagonal 

element’s block—that is, it shows how far and in what directions any very small feature would be smeared out. Evans 

and Achauer showed that such volume metrics are effectively 

invariant under adjusted-damping layer thinning.

For horizontal smoothing, we applied an offset-and-average scheme. Each layer of the thinned-layer model is shifted 

horizontally in increments of 1/nth the block width (in our 

model, B the block width), and the shifted model is inverted. 

Every nonduplicating permutation of block shift is used, and 

so there are n2 (nine) shifted models. Models with each of 

these block-shift permutations are produced and averaged 

within each B-by-B-width block overlap (Evans and Achauer, 

1993, fig. 13.2). Between layer thinning and offset-and-average smoothing, then, the block volume is reduced to A7th that 

of the base model, although the resolvable volume remains 

nearly constant. We gain presentation quality and reduce the 

side effects of block parametrization, which can be significant 

and unintuitive.

50

2 1

5 10

20

100

200

500

0.0 0.05 0.10 0.15 0.20 0.25 0.30

0.002 0.004 0.006 0.008 0.010 0.012

A

SQUARE OF MODEL LENGTH

0.0 0.1 0.2 0.3 0.4 0.5

DATA VARIANCE AFTER INVERSION, IN SQUARE SECONDS

0.01 0.02 0.03

1 2 5

10

20

50

100

200

500 B

residual error = 0.050 s

residual error = 0.050 s

residual error = 0.075 s

residual error = 0.100 s

residual error = 0.075 s

residual error = 0.100 s

residual error = 0.125 s

residual error = 0.150 s

residual error = 0.175 s

Figure 3.—Damping tradeoff curves for principal model (A) and 

detailed model (B). Circled damping value was selected for use, 

on the basis of shape of curve (trading data fit against model 

complexity) and residual error after inversion, which should approximate reading error.

Base model Thinned-layer model

Depth to top Velocity Depth to top Velocity

(km) (km/s) (km) (km/s)

Station elevation 4.86 Station elevation 4.86

6.90 6.23 6.90 6.12

9.81 6.22

12.75 6.33

15.75 6.52 15.75 6.40

18.78 6.57

21.89 6.57

25.00 8.05 25.00 8.03

28.80 8.05

32.60 8.07

36.40 8.11 36.40 8.09

40.22 8.11

44.03 8.12

47.85 8.14 47.85 8.13

51.67 8.14

55.50 8.15

59.32 8.16 59.32 8.15

63.15 8.16

66.97 8.17

70.80 – – – 70.80 – – –

Table 1.—Initial parameters for principal inversion model

[Datum is sea level]

TELESEISMIC TOMOGRAPHY OF THE LOMA PRIETA REGION: IMPLICATIONS FOR STRAIN PARTITIONING
E194 EARTHQUAKE OCCURRENCE

CRUSTAL CORRECTION

We also addressed the issue of possible smearing of shallow 

velocity anomalies to greater depth, using the two methods 

of Dawson and others (1990). Because the shallowest layer 

generally has the largest velocity perturbations (unconsolidated sedimentary deposits, high porosity, weathering) and 

resolution is never perfect, the perturbations caused by these 

shallow upper-crustal structures might distort deeper images 

by smearing downward. To examine these effects, we (1) used 

iterative stripping of shallow structure and (2) separately 

corrected our data for two upper-crustal models derived by 

local-earthquake tomography.

Iterative stripping refers to an iterative inversion for 

velocities in the shallowest layer by itself. At each step, the 

velocity model resulting from inversion for this top layer is 

added back into the starting model of the next iteration, until 

velocity perturbations converge. The intent is to explain as 

many of the data as possible by this one-layer shallow model 

alone. The final, iterated velocity model is then used as the 

first layer of the starting model for a fully layered inversion 

of the entire target volume. Below the shallowest layer, 

this new starting model has the one-dimensional starting 

velocities of the original, whereas the shallowest layer is the 

heterogeneous model taken from the iteration scheme. This 

starting model effectively “corrects” the data, removing all 

residuals that can be (but not necessarily are) explained by 

shallow structure, thus allowing inversions of deeper layers 

to proceed with a minimum of smearing. We call such models 

iteratively stripped.

Although the iterative-stripping method is useful, it may 

overcorrect for shallow structure (Dawson and others, 1990) 

and so should be considered an end member in a continuum 

of models. (The unstripped model is the other end member.) 

Not all data that can be explained by shallow structure 

necessarily should be. To explore this issue, we used two 

recently derived local-earthquake-tomography models, those 

of Eberhart-Phillips and Michael (this chapter) and Harley 

Benz (digital commun., 1995).

Eberhart-Phillips and Michael (this chapter) derive detailed 

three-dimensional P-wave-velocity and VP /VS -ratio models 

of the Loma Prieta region from data recorded on NCSN stations and our temporary array, and from seismic refraction 

or other explosion data. They use the damped-least-squares 

method originated by C.H. Thurber (see Evans and others, 

1994). Eberhart-Phillips and Michael’s model (“crustal model 

A”) is specific to the Loma Prieta region and includes station-correction terms for many of our stations. Harley Benz 

determined P-wave velocities in a broad stretch of northern 

California (“crustal model B”); he used NCSN data for local 

earthquakes, applying the wave-front-tracing method of 

Podvin and Lecomte (1991) and a least-squares inversion 

with smoothness constraints.

To incorporate crustal models A and B into ours, we computed traveltimes through both models for our teleseismic 

wave fronts, using the method of Podvin and Lecomte (1991). 

First, we extracted the appropriate part of each model, creating a slowness mesh of 0.5-km block dimension. (Teleseismic 

wavelengths are ~6 km in the crust.) This remeshing was 

unambiguous for model B because it was derived with the 

same wave-front tracer; for model A, we simply made a trilinear interpolation to the center of each mesh cell from the 

nearest eight nodes of the crustal model. The mesh extends 

from the middle crust (10 km below sea level) to above the 

highest station. Except locally, the local-earthquake models 

are unreliable below the middle crust. In any case, most of the 

traveltime perturbation associated with these models comes 

from shallower structure.

Rather than initializing our plane-wave front on three sides 

of this mesh (the bottom and both sides facing the teleseism), 

we extracted a volume broad enough to initialize only the 

bottom without affecting the seismic array with diffractions 

along the uninitialized sides. We calculated the teleseismic 

traveltime through the slowness mesh (crustal model) and 

interpolated it to each station from the nearest time nodes 

(corners of 0.5-km mesh blocks). We converted these 

Base model Thinned-layer model

Depth to top Velocity Depth to top Velocity

(km) (km/s) (km) (km/s)

Station elevation 4.06 Station elevation 4.06

2.90 5.67 2.90 5.42

4.15 5.61

5.48 5.98

6.90 6.16 6.90 6.03

8.35 6.22

9.82 6.22

11.31 6.30 11.31 6.22

12.75 6.32

14.27 6.34

15.77 6.46 15.77 6.35

17.29 6.46

18.84 6.57

20.41 6.57 20.41 6.57

21.97 6.57

23.48 6.57

25.06 8.04 25.05 8.03

26.96 8.04

28.86 8.05

30.78 8.07 30.77 8.06

32.66 8.07

34.59 8.08

36.51 – – – 36.50 – – –

Table 2.—Initial parameters for detailed inversion model

[Datum is sea level]
E195

traveltimes to traveltime residuals by removing the initial 

wave-front slowness. The residuals were reduced to relative 

residuals by removing an appropriate mean residual.

Finally, we subtracted these “crustal relative residuals” from 

our observed teleseismic relative residuals and inverted the 

difference; we call these deterministically stripped inversions, 

or A- or B-stripped models. The result, if the crustal models 

were perfect and directly applicable, would be our estimate 

of deeper structure based on the teleseisms. However, differing resolution, different wavelengths (differing diffraction 

or evanescence effects near faults), and, possibly, anisotropy 

result in some fraction of the upper-crustal contribution to 

our teleseismic relative residuals being undercorrected or 

overcorrected. Our experience matches that of Dawson and 

others (1990)—that iteratively stripped models probably 

overestimate the crustal correction, whereas deterministically 

stripped crustal models derived from nonteleseismic sources 

commonly underestimate the crustal correction.

RESULTS

UNSTRIPPED MODELS

Maps and cross sections of our velocity models (fully 

smoothed by layer-thinning and offset-and-average steps) for 

the principal- and detailed-model series are shown in figures 

4 and 5, respectively. For each series, unstripped models are 

shown in the first position, iteratively stripped models in the 

last position, and models deterministically stripped by crustal 

models A and B in between. In this subsection, we discuss 

the unstripped models.

In general, our unstripped models of upper-crustal velocity 

structure agree with the results from earlier local-earthquake 

tomography (models A, B; Lees, 1990). We confirm thick 

low-velocity material southwest of the San Andreas fault (for 

example, figs. 4D, 5C), previously identified by teleseismic 

and local-earthquake-tomography studies and interpreted as 

Tertiary marine sedimentary deposits. Although the bottom 

of this low-velocity body reaches as deep as 20 km in our 

models (figs. 4D, 5C), its thickness is not well constrained by 

teleseismic tomography. A high-velocity body northeast of the 

Sargent fault also appears in our models (figs. 4A, 5A) and in 

the local-earthquake results of Lees. The equivalent feature 

reaches the surface between the San Andreas and Sargent 

faults in the model of Eberhart-Phillips and others (1990).

Many upper-mantle velocity gradients fall beneath the 

surface trace of the San Andreas fault (for example, figs. 4B, 

4C, 5B). We compared the subjective coincidence of such 

gradients with the coincidences along arbitrary lines. For 

this comparison, we used a set of lines passing through the 

center of the model at various azimuths. The linear sum of 

such gradient coincidences is in most cases larger for the real 

San Andreas fault than for these arbitrary lines. Therefore, we 

infer that the coincidence is real (see Takauchi, 1994).

In figures 4B, 4D, 5B, and 5C, a large, low-velocity body 

(feature L), centered at 25- to 35-km depth, is the most 

conspicuous feature. Although the depth to the center of 

this low-velocity body is greater than the nominal average 

Moho depth of 25 km and its thickness outwardly appears to 

be about 25 km, its true thickness may be much smaller, and 

its depth uncertain by as much as 10 km. (Exaggeration of 

the vertical dimensions of subhorizontal features is a wellknown artifact in teleseismic tomography, according to Evans 

and Achauer, 1993.) For example, the low-velocity body 

could well be caused by a horizontal, lenticular feature near 

the depth where the velocity anomaly appears to be widest 

(25–35-km depth). Indeed, many geometries of a low-velocity 

body with the width of the imaged feature and fitting within 

its vertical range could cause this anomaly. Thus, it could 

be caused by topography on the Moho, a crustal “root,” as 

suggested by Zandt (1981).

Near 58-km depth, we note two conspicuous low-velocity 

features adjacent to the San Andreas fault (M1, M2, figs. 4C, 

4E). Oceanic-plate geometry was complex enough during 

the final stages of subduction here (for example, Skaer, 

1989) to create such a velocity variation in many ways. For 

example, these two low-velocity features could be caused by 

one or more subducted ridge segments (still hot or relatively 

enriched) or by asthenospheric upwelling through weak or 

open parts of the subducted slab. Whatever their origin, these 

two low-velocity features are separated from one another by 

about 45 km right laterally along the San Andreas fault. If 

they were once contiguous, which we cannot determine, then 

they create a piercing point, a strain marker.

The outward “dips” of low-velocity features M1 and M2 

suggested by figure 4E are almost certainly artifacts of the 

inversion caused by smearing along the dominant ray direction (Evans and Achauer, 1993); M1 and M2 probably are 

caused by the presence of more compact bodies near the 

depths of their maximum velocity anomalies. Similarly, these 

two low-velocity features lie along common P-wave raypaths 

with the shallower low-velocity feature, L, and so we may 

ask whether they are artifacts of that shallower low-velocity 

feature. We summed offdiagonal elements of the resolution 

matrix R to estimate this smearing, and found that smearing of low-velocity feature L can contribute only a minority 

fraction of low-velocity feature M2 and actually mutes the 

velocity contrast of low-velocity feature M1 slightly. Thus, 

low-velocity features M1 and M2 in figures 4C and 4E appear 

to represent real features, lying mainly between 45- and 70-

km depth.

Other features of our model, particularly around the edges, 

are poorly resolved. The high-velocity feature near the 

south edge of our principal model (for example, fig. 4B) is 

consistent with the results of Zandt (1981), who interpreted 

it as Salinian basement, which has higher velocities than 

surrounding Franciscan rocks. However, the depth of this 

high-velocity feature is very poorly resolved here, and so 

further interpretation is inappropriate.

TELESEISMIC TOMOGRAPHY OF THE LOMA PRIETA REGION: IMPLICATIONS FOR STRAIN PARTITIONING
E196 EARTHQUAKE OCCURRENCE

Figure 4.—Maps (A–C) and cross sections (D, E) of compressional-wave velocity perturbation in thinned, smoothed principal model. A–C, 

Blue lines, coast; black lines, faults; gray or white lines, velocity contours. Plan views of layers are centered at about 4-km depth (A), 35-km 

depth (B), and 58-km depth (C); datum is mean station elevation (290 m). Layer depths are slightly exaggerated by flat-earth transformation used. Small red dots, local earthquakes from January 1, 1990, through May 31, 1994, occurring within depth range of each layer; L, 

M1, and M2, low-velocity features discussed in text. D, E, Contour lines are black. Small red dots, earthquakes occurring within 5 km of 

cross-section plane.

STRIPPED MODELS:

UPPER-CRUSTAL STRUCTURE

In this subsection, we describe some of the differences 

between our four models, interpreting those that are of wellunderstood origin. More speculative issues are relegated to 

the section below entitled “Discussion.”

Differences between the stripped and unstripped models are 

minor throughout the upper mantle and, for the most part, in 

the lower crust. Thus, the effects of shallow crustal structure 

on our results are insignificant at such depths, including for 

low-velocity features L, M1, and M2 (figs. 4, 5). However, 

upper-crustal structure does vary greatly between the various 

models, and unsettlingly large anomalies remain in both our 

A- and B-stripped models.

Differences at the periphery of the modeled volumes are 

easily understood as the effects of poor resolution and differing styles of this weakness for the different data sets, 

local and teleseismic. These rays have differing incidence 

angles, and their numbers differ in given areas. The most 

noteworthy example of a mismatch at the periphery of the 

teleseismic model (most intense anomaly, fig. 4A) is at the 

southeast edge of our array. We have one station modeled at 

a much higher velocity than in either local-earthquake model. 

Indeed, the local-earthquake models indicate somewhat 

low velocities in this area, intensifying the high-velocity 

feature in our A- and B-stripped models. This feature may 

be related to Salinian rocks at depth, as inferred by Zandt 

(1981); local earthquakes presumably sense shallower lowvelocity material, with that near the San Andreas having an 

inordinate influence because of the concentration of sources 

there. In another example, the B-stripped principal model 

has a low-velocity area (−11 percent) in the west corner, an 

area of exposed Salinian igneous and metamorphic rocks. 

Unstripped A-Stripped

B-Stripped Iteratively Stripped

-14% Mean 14%

50 km

Depth: 0.0-7.19 km Contour Interval 4%

A

N

Unstripped

L

N

A-Stripped

L

B-Stripped

L

Iteratively Stripped

L

-3.5% Mean 3.5%

50 km

Depth: 33-37 km Contour Interval 1%

B
E197

Unstripped

M1

M2

A-Stripped

M1

M2

B-Stripped

M1

M2

Iteratively Stripped

M1

M2

-3.5% Mean 3.5%

50 km

Depth: 56-60 km Contour Interval 1%

N

C

Figure 4.—Continued

TELESEISMIC TOMOGRAPHY OF THE LOMA PRIETA REGION: IMPLICATIONS FOR STRAIN PARTITIONING

This mismatch is absent in the A-stripped principal model. 

We infer a weakness in local-earthquake model B, probably 

owing to the low seismicity of that area and to our not having 

used data from temporary stations in the area (as we did with 

model A). In a similar example, the high-velocity feature at 

the northernmost station (blue dot nearest the Santa Clara 

Valley, fig. 4A) in the unstripped principal model is also present in the A- and B-stripped models, and even accentuated. 

The velocity anomalies in model A are northwest-southeast 

elongated (for example, Eberhart-Phillips and Michael, this 

chapter, fig. 5), presumably partly representing the resolution effects of high concentrations of sources and receivers 

along the San Andreas fault. Teleseismic features streak 

subvertically for equivalent reasons, thus at right angles to 

the streaking in crustal model A. Teleseisms probably are 

sensing deeper-crustal high-velocity ultramafic rocks of the 

Franciscan Formation.

Contradictions in the central parts of our model—for 

example, along the Sargent fault in crustal model B, along 

the Zayante fault in crustal model A, and near (northeast 

of) the intersection of the Sargent and San Andreas faults in 

both models—are interpreted below in the section entitled 

“Discussion.”

RESOLUTION

The diagonal elements of the resolution matrix R for the 

base models, as a general indicator of the spatial variation in 

resolution, are shown in figure 6. These diagonal elements 

and associated standard errors are typical of well-sampled 

Unstripped

Coast SAF Cal.

L

A-Stripped

Coast SAF Cal.

L

-3.5% Mean 3.5%

D (S43W) 50 km Contour Interval 1% (N43E) (N43E) D' D

B-Stripped

Coast SAF Cal.

L

Iteratively Stripped

Coast SAF Cal.

L

Moho

Moho

D

Unstripped

L

M2 M1

A-Stripped

L

M2 M1

-3.5% Mean 3.5%

G (S45E) 50 km Contour Interval 1% (N45W) G'

B-Stripped

L

M2 M1

Iteratively Stripped

L

M2 M1

Moho

Moho

E
E198 EARTHQUAKE OCCURRENCE

teleseismic models; the diagonal elements are as large as 

0.77 for the principal model and 0.69 for the detailed model 

in the best-resolved areas. Because the ray sets are essentially 

identical between unstripped and stripped models (with trivial 

variations due to first-layer velocity differences changing the 

raypaths slightly), figure 6 is representative of all four models. 

Worst-case standard errors in these areas are about 0.5 percent 

for the unstripped principal model and 1.2 percent for the 

unstripped detailed model. As expected, the resolution is fair 

to good at the centers of the modeled volumes and gradually 

decreases toward the peripheries.

To help readers visualize resolution effects in more detail, 

including the directions and lengths of smearing, volume 

metrics for selected blocks are plotted in perspective in figure 

7. Volume metrics are described above in the section entitled 

“Inversion Technique” but, succinctly, should be thought 

of as the “impulse responses” of our tomographic imaging 

process—the sizes and shapes of very small, intense, point 

anomalies once we have imperfectly imaged them.

Volume metrics for blocks near the centers of the modeled 

volumes generally are more compact than for blocks near the 

peripheries. Indeed, some edge blocks smear over a large part 

of the model (for example, fig. 7B) and are largely meaningless. The shape and orientation of smearing generally is radially away from the seismograph array, in the direction of the 

most plentiful (P wave) rays. These characteristics are also 

typical of well-sampled teleseismic-tomography models.

Unstripped A-Stripped

B-Stripped Iteratively Stripped

-28% Mean 28%

20 km

Depth: 0.0-3.38 km Contour Interval 8%

N

A

Unstripped

L

A-Stripped

L

B-Stripped

L

Iteratively Stripped

L

-3.5% Mean 3.5%

20 km

Depth: 27-29 km Contour Interval 2.5%

N

B

Figure 5.—Maps (A, B) and cross sections (C, D) of compressional-wave-velocity perturbation in thinned, smoothed detailed 

model. Same symbols as in figure 4. Datum is mean station elevation (476 m). Plan views of layers are centered at about 2-km 

depth (A) and 28-km depth (B).

Unstripped

SAF

L

A-Stripped

SAF

L

-3.5% Mean 3.5% 20 km D (S43W) Contour Interval 2.5% (N43E) D'

B-Stripped

SAF

L

Iteratively Stripped

SAF

L

Moho

Moho

C
E199

low-velocity body lies between the San Andreas and Zayante 

faults (unstripped model, fig. 5A). The A-stripped model 

effectively corrects for this low-velocity feature, except at 

a station at the intersection of the Sargent and San Andreas 

faults. The B-stripped model is less effective—no surprise, 

given that model’s regional scale. We conclude that crustal 

model A compensates well for the core of the low-velocity 

marine sedimentary rocks between these faults but that the 

A-stripped principal teleseismic model shows some mismatch of the Zayante fault boundary of this feature. If the 

low-velocity wedge dips southwest, teleseismic rays at the 

surface would have traversed some of it, making the feature 

look wider to us than to Eberhart-Phillips and Michael (this 

chapter). Similarly, our results are more subject to diffraction and evanescence than theirs, a difference that is likely to 

matter along a fault or other major contact, and so we defer 

to their model (A).

Leaving the upper crust, the likely cause of low-velocity 

feature L (figs. 4B, 4D, 4E, 5B–5D) is a crustal “root,” as 

suggested above and by Zandt (1981). Bürgmann and others 

(1994) evaluated fission-track dates, geomorphology, and 

geodesy in the Santa Cruz Mountains, concluding that the 

region northeast of the San Andreas fault has undergone rapid 

uplift of about 3 to 4 km over the past 4.6 m.y., whereas both 

the current uplift rate and total uplift since 10 Ma southwest 

of the San Andreas fault are much lower. Bürgmann and 

others also concluded that the locus of uplift has remained 

about fixed in the Loma Prieta region on the northeast side 

of the fault, whereas the southwest side was uplifted mainly 

as it passed the fault bend near Loma Prieta, leaving a trail 

of moderate uplift to the northwest. Therefore, a significant 

TELESEISMIC TOMOGRAPHY OF THE LOMA PRIETA REGION: IMPLICATIONS FOR STRAIN PARTITIONING

Unstripped

L

A-Stripped

L

-3.5% Mean 3.5% 20 km

G (S45E) (N45W) G'

B-Stripped Contour interval 2.5%

L

Iteratively Stripped

L

Moho

Moho

D

Figure 5.—Continued

DISCUSSION

Differences in modeled upper-crustal structure between 

figures 4A and 5A and among the various stripping methods 

shed light on what is “real” in the upper crust. Two low-velocity features (yellow dots) northeast of the Sargent fault in the 

A-stripped principal model imply that model A velocities are 

higher there than we modeled with teleseisms. Crustal model 

B leaves a similar but weaker feature (that is, is closer to our 

result). We suspect that complexity near the intersection of 

the Sargent and San Andreas faults and the steepness of thrust 

faults in this area are contributing to this mismatch, and that 

higher velocity material underlying a shallower velocity low 

may be present. (This area contains many slivers of ultramafic 

rock, commonly fault bounded.) For the uppermost crust, we 

consider the local-earthquake models more reliable.

Disagreement also exists between the local-earthquaketomography models of Eberhart-Phillips and others (1990) 

and Lees (1990) on the location of the high-velocity body 

near the Sargent fault. Their results and ours all contain 

strong velocity gradients near these faults (for example, 

±28.0-percent velocity perturbation in our detailed model), 

although the location of the responsible high-velocity body 

northeast of the San Andreas fault is poorly constrained. In 

particular, it is unclear on which side of the Sargent fault it 

lies in the uppermost crust. It appears to dip southwest and 

probably is an ophiolitic fragment bounded above or below 

by the Sargent fault.

Along the Zayante fault, the A-stripped principal model 

leaves systematically lower velocities than the unstripped or 

B-stripped models. In our detailed models, the core of the 
E200 EARTHQUAKE OCCURRENCE

crustal root northeast of the San Andreas fault at Loma Prieta 

and less or no root to the southwest are both plausible and 

consistent with our result.

The traveltime delay due to low-velocity feature L is about 

0.09 s, equivalent to crustal thickening of about 2 to 5 km 

(depending on the velocity contrast across the Moho, which 

is uncertain). In cross section (for example, figs. 4D, 5C), 

this body is triangular or diamond shaped—the expected 

pattern for a horizontal, lenticular object, and a likely shape 

for a crustal root (Evans and Achauer, 1993, figs. 13.7, 13.8). 

Therefore, our result is consistent with a crustal root.

If low-velocity feature L is caused by crustal thickening, then a gravity low might also be anticipated; however, 

no clear signature of such a body is evident in the gravity 

field mapped by Roberts and Jachens (1993). Therefore, to 

test this interpretation, we calculated the expected gravity 

anomaly caused by a simple crustal-root model, using the 

method of Singh (1977) for disk-shaped bodies. The root is 

approximated by two thin, concentric disks at Moho depth, 

the smaller below the larger. The combined volume of these 

disks equals that of a “root” of 3-km maximum thickness, 

with a cosine-bell shape and a diameter of 28 km. Assuming 

that the density contrast across the Moho is 0.4 g/cm3 (Donald 

Plouff, oral commun., 1994), the resulting gravity anomaly 

peaks at about −2 mGal (fig. 8). This anomaly is broad as 

Depth: 0.0-7.19 km Depth: 33-37 km

Depth: 56-60 km 50 km

Resolution

(Diagonal Elememts)

Contour Interval 0.1

0.0 A

0.5

0.8

N

D (S43W) (N43E) D'

Coast SAF Cal.

G (S46E) (N46W) G'

B

50 km

Resolution

(Diagonal Elememts)

Contour Interval 0.1

0.0

0.5

0.8

Depth: 0.0-3.38 km Depth: 26-31 km

D (S43W) SAF (N43E) D' G (S45E) (N45W) G'

0.0 0.5 0.8

Contour Interval 0.1 20 km

C

N

Figure 6.—Diagonal elements of resolution matrix of base models. A, Plan views of principal model. B, Cross sections of principal model. 

C, Plan and cross-sectional views of detailed model. Plan views are for layers containing thinned layers shown in compressional-wave velocity-perturbation maps (figs. 4, 5); that is, layer shown here is subdivided into three layers, one of which is shown in figures 4 and 5. Cross 

sections match those in figures 4 and 5.
E201

If we assume that low-velocity features M1 and M2 were 

contiguous when formed, then, regardless of their origin, they 

constitute a marker for strain partitioning—a piercing point. 

Although this assumption is certainly not required, it is plausible. If correct, however, it has important implications.

The northward passage of the Mendocino Triple Junction 

and consequent inception of the transform boundary at this 

part of the North American plate occurred from as early 

as about 8 Ma (Atwater, 1989) to as late as about 6.5–5.5 

Ma (Riddihough, 1984). At this part of the Pacific plate, it 

occurred much earlier, at 30–20 Ma (Atwater, 1989). The 

position of the plate boundary also has moved laterally over 

time, but it probably has always been diffuse and generally 

moving northeastward, inland (Atwater, 1989). Some evidence 

exists that it has also moved southwestward at least once, 

with east-bay faults dominating from 8 to 6 Ma, then primary 

shear returning to the San Andreas fault and its adjacent predecessor, the Pilarcitos fault (McLaughlin and others, 1996). 

Because we do not know what low-velocity features M1 and 

M2 are, we do not know with which plate to associate them, 

and so we do not know which time interval is appropriate for 

estimating San Andreas fault shearing of a piercing point at 

low-velocity features M1 and M2. However, even using the 

shortest time period (6 m.y.), motion across the San Andreas/

Pilarcitos ranges from about 150 to 200 km, possibly more 

(fig. 9; Dickinson and others, 1972; Sedlock and Hamilton, 

1991; R.J. McLaughlin, oral commun., 1997). Thus, the 45-km 

offset between low-velocity features M1 and M2 (fig. 4C) is 

much too small to have been caused by a narrow, persistent 

shear locus beneath the San Andreas fault. Therefore, if lowvelocity features M1 and M2 are a valid piercing point, then 

this discrepancy implies that the San Andreas fault was a deep, 

narrow shear locus for only a small part of that time.

As much has been proposed by Furlong and others (1989), 

who inferred a midcrustal or lower-crustal detachment fault 

connecting the San Andreas fault with the Hayward/Calaveras 

fault system, and a deeper, broad lithospheric boundary now 

centered beneath the Hayward/Calaveras fault system. Beneath 

the initial plate boundary, they modeled uppermost-mantle 

shearing as a transient pattern, frozen into place within a few 

million years after passage of the Mendocino Triple Junction. This model is consistent with the limited offset that we 

observe between low-velocity features M1 and M2 if shearing 

at their depth “froze in” within 2 to 3 m.y. of the inception of 

faulting (fig. 9). This consistency is the second major result 

or our study.

An alternative interpretation of the 45-km offset between 

low-velocity features M1 and M2, still assuming that they 

originated as a single feature, is that mantle shear has always 

been broadly distributed at this depth, rather than constrained 

to a narrow locus. A broad shear locus might be similar to 

the modern surface displacements measured by Lisowski 

and Savage (1992), a fault-parallel band about 85 km across. 

If so, the fraction of the plate motion taken up between M1 

and M2 (centered ~17 km apart in the northeast direction) 

TELESEISMIC TOMOGRAPHY OF THE LOMA PRIETA REGION: IMPLICATIONS FOR STRAIN PARTITIONING

well as weak (~ 42-km half-width) and so it would be difficult 

to identify among anomalies from shallower features in this 

structurally complex region.

Many other explanations are possible for this velocity 

low, including complexities in the last-subducted material 

that may underlie the region. Because the low-velocity body 

in figures 4B and 5B is elongate parallel to the Sargent and 

San Andreas faults, we should also consider possible local 

upper-mantle shear heating. Few heat-flow data are available 

in this region (Lachenbruch and Sass, 1980), and any surface 

heat-flow anomaly from such shear heating would still be far 

below equilibrium values (for example, Lachenbruch and 

others, 1976). Therefore, the shear-heating model cannot 

be dismissed. However, (1) the absence of the low-velocity 

anomaly beneath other sections of the San Andreas fault 

and (2) the great nucleation depth of the 1989 Loma Prieta 

earthquake (suggesting that the adjacent lower crust is cold) 

argue against shear heating.

The presence of velocity gradients along the San Andreas 

fault in the upper mantle suggests that the fault was at least 

at one time a deeply penetrating feature present in the upper 

mantle as a narrow shear locus. As mentioned above, simple 

coincidence tests suggest that the San Andreas fault is more 

closely correlated with such gradients than most arbitrary 

lines that can be drawn through the centers of figures 4 and 

5, or equivalent plots at other depths. The inference of a deep, 

narrow shear locus beneath the fault during some interval is 

the first major result of our study. We may ask when this locus 

was active, and we address that question next in reference to 

low-velocity features M1 and M2 (fig. 4C).

With the complexity of subducted material in the Loma 

Prieta region, low-velocity features M1 and M2 may not 

have originated as contiguous. Certainly, we can say little 

about what they are. R.J. McLaughlin (oral commun., 1997) 

suggested that low-velocity feature M2 could be the source 

volume for volcanism and its associated hydrothermal alteration and mineralization at New Almaden, whereas low-velocity feature M1 could be the source for the Mindego Volcanics 

in the La Honda Basin. He noted that these thermal events 

are of different ages and suggested that low-velocity features 

M1 and M2 may also be of different age. In our view, the 

complexity of crustal tectonics in the region and its likely 

control of volcanism suggests that being the heat sources 

for different-age volcanic rocks does not imply that the two 

mantle anomalies are of different ages. We also note that the 

preservation of low-velocity upper-mantle anomalies beneath 

volcanoes (presumably, regions of basaltic partial melt) for 

more than 10 m.y. is, at least, rare. The oldest such anomaly 

known to us is the ≥6-Ma anomaly beneath the eastern Snake 

River Plain (Evans, 1982). Beneath the Miocene southwestern Nevada Volcanic Field is a deep high-velocity anomaly 

interpreted as the cooled and (or) depleted mafic rock of a 

dead magma source (Evans and Smith, 1995). Although the 

correlation of low-velocity features M1 and M2 with Neogene 

volcanism is intriguing, it is uncertain.
E202 EARTHQUAKE OCCURRENCE

we accounted for the fraction of plate motion taken up elsewhere in the Western United States. However, the resulting 

shape of low-velocity features M1 and M2 would probably 

be elongate and still contiguous. Therefore, we prefer the 

frozen-in narrow-shear-locus model described first, or some 

composite model.

Finally, we emphasize that low-velocity features M1 and 

M2 might simply be unrelated. They could have originated 

in any spatial configuration, from their current one to a large, 

fortuitous, left-lateral offset now erased and reversed by 

motion on the San Andreas fault; they may also be of different 

age. We cannot rule out any of these interpretations without 

further study.

CONCLUSIONS

Our teleseismic-tomography images of the San Andreas 

and related faults in the vicinity of the Loma Prieta earthquake rupture generally agree with previous local-earthquake 

tomography for the upper crust and with earlier, less detailed 

teleseismic tomography. Beneath the Loma Prieta highlands, 

we model a large, low-velocity body near the Moho. This 

anomaly is most likely due to a crustal root of about 2- to 

5-km thickness beneath those highlands. Alternatively, it 

may be caused by complexities in late-subducted material or 

by laterally discontinuous shear heating in the upper mantle 

beneath the San Andreas fault.

We find many velocity gradients coincident with the San 

Andreas fault at all depths. These coincidences are more 

common than for most other, arbitrary lines drawn through 

the center of the model, and so we consider the coincidence 

meaningful. In addition, we find a right-lateral offset between 

two low-velocity bodies in the upper mantle beneath the San 

Andreas. If these features were contiguous before the initiation of strike-slip faulting in the Loma Prieta region, then they 

are markers for strain partitioning since that time—a piercing 

60 40 20 0 20 40

50

0

50

-40 A

-20

0

-60

B

50

0

50

60 40 20 0 20 40

-40

-20

0

-60

C

20

0

20

30 20 10 0 10 20

-20

-10

0

-30

Depth (km) Depth (km) Depth (km)

NW (km) SE

NW (km) SE

NW (km) SE

SW (km) NE

SW (km) NE

SW (km) NE

Figure 7.—Perspective views of representative volume metrics from 

the thinned principal model (A, B) and the thinned detailed model 

(C). View is northward from an elevation angle of 40°. Equivalent 

diagonal-element resolution is 0.75 in figures 7A and 7C but 0.60 

in figure 7B. Metrics are plotted together to save space, each being 

a more or less contiguous set of model blocks of one color (except 

diagonal-element block, which is magenta in each figure). In figure 

7A, red is for a block near center of low-velocity feature L, green 

for low-velocity feature M1, cyan for low-velocity feature M2, 

and yellow for a fairly well resolved block near northwest edge of 

model (see fig. 4 for locations). Figure 7B shows example of a very 

poorly resolved edge block—it cannot be interpreted meaningfully. 

In figure 7C, cyan is for center of low-velocity feature L, and yellow is for a typical edge block. For perspective, center of each magenta diagonal-element block is connected to surface by a magenta 

vertical line. Coastline (blue) and the San Andreas, Zayante, and 

Sargent faults (red) are shown at surface.

would be less than or equal to about 17/85, or 20 percent of the 

total plate motion since 8 Ma. That total motion is ~370 km 

(Atwater and Molnar, 1973), yielding an expected separation 

of as much as ~74 km between low-velocity features M1 and 

M2 in the northwest direction. The match would be closer if 

RADIUS, IN KILOMETERS

GRAVITY ANOMALY,

IN MILLIGALS

0 5 10 15 20 25 30

0.0 0.5 1.0 1.5 2.0 2.5

Moho

-0.4 g/cm3

-0.4 g/cm3

Anomaly

Sketch of root used:

Sketch of "ideal" root:

Moho

-0.4 g/cm3

Figure 8.—Estimated gravity anomaly due to a cylindrically symmetric lenticular “crustal root” comparable to low-velocity feature 

L (see figs. 4 and 5 for location). Root is approximated by two thin 

disks centered at 0-km radius, each with 0.4-g/cm3 density contrast. 

Roughness of gravity anomaly is numerical noise.
E203

point. In particular, the small size of the offset would imply 

that the San Andreas/Pilarcitos fault could not have been a 

narrow upper-mantle shear locus for the entire period since 6 

Ma—indeed, not for more than a few million years. This result 

is consistent with the detachment-fault model of Furlong and 

others (1989) if upper-mantle shear in this region “froze in” 

within a few million years after passage of the Mendocino 

Triple Junction, as predicted by their model.

ACKNOWLEDGMENTS

We thank Art Lachenbruch, Bob Jachens, and Don Plouff for 

their assistance with heat-flow and gravity data. Don Reed and 

Richard Sedlock of San Jose State University gave invaluable 

feedback and support at every stage. Reviews by Harley Benz, 

Phil Dawson, and Bob McLaughlin significantly improved the 

manuscript. Hitoshi Saito provided both moral and practical 

support to the first author. This research was partly supported 

by the Earthquake Hazards Reduction Program. The first 

author was partly supported by a San Jose State University 

foreign-student tuition waiver, by an Association for Women 

Geoscientists’ Chrysalis Scholarship, by a scholarship from 

the Geology Department of San Jose State University, and 

by the USGS.

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