PHILIPPINE -.!, 

PACIFIC PLATE 
MAP OF THE SAN ANDREAS FAULT SYSTEM, CALIFORNU7 

Maps of northern and southern California printed on flyleaf inside front 

cover and on adjacent pages show faults that have had displacement withil 

the past 2 million years. Those that have had displacement within hiiorica 

time are shown in red. Bands of red tint emphasize zones of historical dis 

placement; bands of orange tint emphasize major faults that have had Qua 

ternary displacement before historical time. Faults are dashed where uncer 

tain, dotted where covered by sedimentary deposits, and queried when 

doubtful. Arrows indicate direction of relative movement; sawteeth 01 

upper plate of thrust fault. 

These maps are reproductions, in major part, of selected plates from th~ 

"Fault Map of California," published in 1975 by the California Division o 

Mines and Geology at a scale of 1:750,000; the State map was compilec 

and data interpreted by Charles W. Jennings. New data about faults, no1 

shown on the 1975 edition, required modest revisions, primarily additions 

however, most of the map was left unchanged because the Californii 

Division of Mines and Geology is currently engaged in a major revision anc 

update of the 1975 edition. Because of the reduced scale here, names o 

faults and places were redrafted or omitted. Faults added to the reducec 

map are not as precise as on the original State map, and the editor of tht 

volume selected certain faults and omitted others. Principal regions fo: 

which new information was added are the region north of the San Francisco 

Bay area and the offshore regions. 

Many people have contributed to the present map, but the editor is 

solely responsible for any errors and omissions. Among those contributing 

informally, but extensively, and the regions to which each contributed wen 

G.A. Carver, onland region north of lat 40Â N.; S.H. Clarke, offshore regta 

north of Cape Mendocino; R.J. McLaughlin, onland region between la' 

40'00' and 40'30' N. and long 123-30' and 124'30' W.; D.S. McCulloch 

offshore region between lat 35" and 40" N.; J.G. Vedder, offshore reglor 

south of lat 35O N.; and D.G. Herd, southern San Francisco Bay region. Th< 

Fault Evaluation Program of the California Division of Mines and Geology 

under the direction of E.W. Hart, provided much data about many faults 

Unpublished material about the Bartlett Springs fault zone that was gath 

ered by Geomatrix Consultants for the Pacific Gas and Electric Co. wag 

very useful. In addition, selected publications that provided invaluable dati 

include Bortugno (1982). Herd (1977). Herd and Helley (1977). Pampeyai 

and others (1981), and Yerkes and others (1980). Full references for thes 

publications are lied at the end of chapter 2. 

To all of the above, and to numerous others, the editor expresses hi: 

great appreciation. 



10 20 30 40 50 KILOMETERS I 

THE SAN ANDREAS FAULT SYSTEM, 

CALIFORNIA 
The San Andreas fault in central California. Striking linearity of the trace of the fault is characteristic of 

strike-dip faults. Carrizo Plain to left; Temblor Range to right. View northwest. 
The San Andreas Fault System, 

California 

ROBERT E. WALLACE, Editor 

U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 

An overview of the history, geology, geomorphology, 

geophysics, and seismology of the most well known 

plate-tectonic boundary in the world 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1990 
U.S. DEPARTMENT OF THE INTERIOR 

MANUEL LUJAN. JR, Secrrfaly 

US. GEOLOGICAL SURVEY 

DALLAS L PECK. Director 

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

is for descriptive purposes only and does not imply endorsement 

by the US. Government 

First printing: 1990 

Second printing: 1991 

The San Andreas fault system, California 1 Robert E. Wallace, editor. 

p. cm. - (U.S. Geological Survey professional paper : 1515) 

Includes bibliographical references (p. 

Supt. of Doo. no.: I 19.16:1515 

1. San Andreu Fault (Calif.). L Wallace, B.E. (Robert Earl), 1916- . Series. 

QE606.5.U6S26 1990 90-13874 

551.8"r09794Ã‘dc2 CIP 

For sale by the Books and Open-File Reports Section, US. Geological Survey, 

Federal Center. Box 25425, Denver. CO 80225 
PREFACE 

With the increasing public concern about the potential 

for destructive earthquakes in California since the great 

Alaska earthquake of 1964, and the general acceptance of 

the concept of plate tectonics and sea-floor spreading by 

earth scientists in the late 1960's. the San Andreas fault 

has gained wide attention. The fault has long been 

recognized as the source of the destructive San Francisco 

earthquake of 1906 and of the similarly large Fort Tejon 

earthquake of 1867, as well as the smaller (M=7.1), but 

also destructive, Loma Prieta earthquake of 1989. Many 

textbooks in recent decades have included photographs, 

maps, and descriptions of the San Andreas fault, and so 

among earth scientists around the world, few geologic 

features have greater identity than the San Andreas 

fault. The fault-better designated a "fault system" 

because of its complexities-represents an exceptional 

example of a plate margin that can be seen and studied on 

land; many other plate margins are covered by the 

oceans. 

This volume is addressed to a varied audience, but 

especially to earth scientists who wish to gain a brief 

overview of the San Andreas fault system. We hope that 

the nontechnical reader also will find the volume interesting and useful. Each chapter has its own references 

that direct the investigator more to specific literature; 

suggestions for additional reading and sources of information are included at the end of the volume. 

The public-safety issues of earthquake-hazard mitigation and earthquake prediction are not covered in this 

volume except by implication. Fundamental concepts and 

understanding of how the solid Earth works are essential 

to the development of realistic and effective procedures 

for hazard mitigation. Without such understanding, mitigation measures may be ineffective and wasteful of 

resources. Earthquake prediction requires geologic and 

geophysical models and data that constitute valid bases 

for extrapolation into the future. We believe that these 

concerns are well served by the reviews in this volume. 

The 10 chapters that follow review geologic, geomorphic, geophysical, geodetic, and seismologic information 

about the San Andreas fault system. Although the 

material is intended to represent our current state of 

knowledge and understanding, many investigators will 

find their own specialities inadequately treated. A full 

discussion of the more rapidly changing and controversial 

ideas currently in play is impossible, given the constraints of this volume. 

Mwniscript approved for publication, May 4.1990. 

The need for such a volume has been recognized since 

the 1960's, but until recent years, data about the San 

Andreas fault system was so incomplete that a significant 

review seemed premature. Now, a general overview of 

many significant elements seems possible, but even so, 

the questions far outnumber the answers, and if history 

is any guide, many of the concepts put forth here will 

have changed markedly within the next few years. 

Indeed, one of the purposes of this volume is to assist and 

hasten the evolution of our understanding through the 

consolidation, under a single cover, of some of the current 

ideas and models. 

Reference to a few of the more outstanding problems 

concerning the San Andreas fault system may serve to 

suggest how little is yet known about the fault system 

and how much remains to be learned. In studying the San 

Andreas fault system, opportunities abound for learning 

how the Earth works in a general sense. 

A group of particularly significant problems can be 

collectively expressed under the question: "How does a 

fault system that has primarily strike slip terminate at its 

northwest and southeast ends, and how does it bottom 

out at depth?" 

At the northwest end of the fault system, the Mendocino triple junction represents an intriguing structural 

knot where the North American, Pacific, and Gorda 

plates join. A fourth block at depth, made up of material 

below the North American plate but east of the San 

Andreas fault and south of the Goda plate, also is 

juxtaposed with these three named plates. How are the 

diverse motions of these four plates or blocks accommodated where they join? Clearly, severe space problems 

must occur at detailed scales, even though the gross 

theoretical geometry of triple junctions has been fairly 

well described. How do these four structural blocks 

interact to influence the energetics of the fault system 

(see chap. lo)? How did the triple junction migrate over 

time, and how are the consequences of that migration 

recorded geologically? 

Crustal convergence also strongly influences the fault 

system. Where major bends occur, as in the Transverse 

Ranges region, structural complexities arise. There, 

major left-lateral faults splay from the main San Andreas 

fault, and dense clusters of earthquakes extend to depths 

of 20 km. Elsewhere, as in the Santa Cruz Mountains, 

some segments of the fault dip at steep, but not vertical, 

angles. 

Adjacent to the San Andreas fault on its east side lies 

the North American plate, at least in the upper few tens 

of kilometers; but below this plate, as indicated above, is v 
VI THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

a block of almost unknown characteristics left in the wake 

of the eastward-moving and subducting plate now represented by its remnants, the Gorda and Juan de Fuca 

plates. After pulling away from the Pacific plate in its 

eastward passage, did this plate leave remnants here and 

there under the North American plate? Did the spreading center spread continuously, or did it move eastward 

in one or more leaps? What sort of mantle material rose 

to fill in behind the stern of the eastward-moving plate? 

Was this newly emplaced material similar to that being 

generated at modern spreading centers, and to what 

extent did the overlying blanket of continental material 

alter both the geometry and thermal histories of the 

emplaced rocks? 

Some of the models mentioned above imply decoupling 

between the subducting plate and the North American 

plate, as well as a rather significant discontinuity between the North American plate and the underlying 

backfill behind the stem of the eastward-moving plate. 

Some of the problems of this ''window" behind the 

subducting plate are discussed in chapter 3. The San 

Andreas fault, indeed, may bottom in a zone of decoupling, either within the crust or below, possibly a 

low-angle thrust fault, as described in chapter 1, or 

perhaps involving gravitationally driven detachments. 

The characteristics of this decoupling are almost unknown. Much needs to be learned before the style of 

stress and strain propagation across such discontinuities 

can be addressed adequately. 

Present-day strain is demonstrable by geodetic techniques (see chap. 7), and longer-term strain is represented geologically by the pattern of folds, faults, and 

magmatic intrusion into the upper crust (see chaps. 3,4). 

Release of elastic strain, its timing and spatial distribution, is nicely displayed by seismicity, especially microseismicity (see chap. 5). Heat-flow measurements 

provide an important insight into the energetics of the 

San Andreas fault system (see chap. 10). Integration and 

comparison of these data sets, however, reveal numerous 

unresolved problems and apparent paradoxes. As reported in chapter 10, no sharp increase in heat flow is found 

directly over the San Andreas fault, even though heat 

would be expected to be generated in the narrow fault 

zone by the annual slip of several centimeters on the 

fault. Instead, the heat flow is distributed across a broad 

zone, further suggesting distributed slip on a subhorizontal plane that decouples the upper-crustal materials 

from those below. 

The detailed characteristics of the fault zone itself are 

far from fully understood. Low values of stress drop that 

occur during seismic events have long been known (see 

chaps. 5, 10). The dominance of right-lateral slip along 

the fault, despite evidence for fault-normal compression, 

together with the absence of a pronounced heat-flow 

anomaly, attests to general weakness of the fault. How 

the fault zone has grown to its present 0.5- to 1-km width, 

given this weakness, is also a puzzle. Furthermore, at 

many places along the fault, as in central California, the 

width of the fault zone is appropriately considered to be 

10 km or more wide; that is the width of highly sheared 

and deformed rocks which lie between relatively undeformed terrane to the northeast and southwest. What is 

the nature of asperities, or strong points, on the fault, 

and how do its stronger and weaker parts interact? To 

what extent does plastic-behaving fault gouge move 

within the fault zone to change the overall geometry? 

The role of water in the kinematics and dynamics of the 

fault system can scarcely be overemphasized, and yet 

almost nothing is known about the actual hydrotectonic 

relations. To what extent does elevated or reduced pore 

pressure modify the properties of the lithologic packages 

of rocks at different places along and across the fault? 

What geochemical changes in the fault zone are enhanced 

by the movement of fluids along the fault and through 

rocks adjacent to the fault? What mineralogic changes 

take place as a result of thermal changes related to 

friction, possibly accompanied by exsolution of water 

from minerals? Can localized, tectonically elevated pore 

pressure initiate slip, and once slip acceleration occurs, 

what role does friction play in the dynamics of slip? Why 

are so few volcanic rocks associated with the fault except 

for those related to the passing triple junction? 

What are the rates of fault slip, folding, and the overall 

budget of deformation among various forms of strain? 

How do erosional rates compare with tectonic rates in 

changing the landscape, and to what extent do these 

processes deviate from linearity? How do complex geomorphic processes interact among themselves, as well as 

with the tectonic processes? How can the ubiquitous 

landforms be interpreted to illuminate the younger 

history of the San Andreas fault? 

Strands and branches of the San Andreas fault system 

bound numerous exotic terranes, those aggregates of 

rocks which are so dissimilar that they could not have 

been born in their present relation to one another. The 

patterns of movement and distances traveled by these 

terranes may constitute the most significant characteristic and role of the San Andreas fault in the overall 

scheme of global tectonics. By whatever mechanisms 

these exotic blocks or terranes were transported, the 

western part of North America has been enlarged by the 

accretion of these "strangers," while at the same time 

other pieces of older continental material have been 

plucked away, eventually to join land somewhere to the 

northwest. 

This volume represents but a small punctuation mark 

in the early stage of our understanding of the San 

Andreas fault system and the tectonics that it highlights. 

Most of the story has yet to be learned. 
CONTENTS 

(Numbers indicate chapters, 

PW 

V 

1. General features----------------------------- 3 

Robert E. Wallace 

2. ~eo~~~hi~ expression 15 

Robert E. Wallace 

3. Geology and plate-tectonic development------------ 61 

William P. Irwin 

4. Quaternary deformation - -- 83 

Robert D. Brown, Jr. 

5. Seismicity 1980-86 ---- 115 

David P. Hill, Jerry P. Eaton, and Ludle M. Jones 

6. Earthquake history, 1769-1989------------------- 153 

William L. Ellsworth 

7. Present-day crustal movements and the mechanics of 

cyclic deformation-- 189 

Wayne Thatcher 

8. Lithospheric structure and tectonics from seismicrefraction and other data 207 

Gary S. Fuis and Walter D. Mooney 

9. Crustal and lithospheric structure from gravity and 

magnetic studies 239 

Andrew Griacom and Robert C. Jachens 

10. Stress and heat flow 261 

Arthur Lachenbmch and A. McGarr 

Supplement: Additional reading and source material 279 

Robert E. Wallace, compiler 



T he San Andrew fault system, a complex of faults that display 

predominantly large-scale strike slip, is part of an even more 

complex system offaults, isolated segments of the East Pacific Rise, 

and scraps of plates lying east of the East Pacific Rise that 

collectively separate the North Akan plate from the Pacific 

plate. 

CONTENTS 

1. GENERAL FEATURES 

INTRODUCTION 

This chapter briefly describes the San Andreas fault 

system, its setting along the Pacific Ocean margin of 

North America, its extent, and the patterns of faulting. 

Only selected characteristics are described, and many 

features are left for depictions on maps and figures. The 

other chapters in this volume elaborate on the history and 

evolution of the fault system, and the behavior of the 

Earth's crust and upper mantle within the fault system. 

Because of the extent and complexity of the San 

Andreas fault system, it is helpful to distinguish between 

the broad, complex feature as seen on a map of the 

Western United States and the individual faults on which 

displacements occur to produce single earthquakes. 

From larger to smaller features, the terms ''fault system," "fault zone," "fault," and "fault branches, splays, 

strands, and segments" are useful. 

The term "San Andreas fault system" refers to the 

network of faults with predominantly right-lateral strike 

slip that collectively accommodate most of the relative 

motion between the North American and Pacific plates. 

The boundaries of this fault system are poorly defined, 

but to separate the San Andreas fault system from other 

tectonic provinces and systems, it is useful to limit the 

term to the set of faults along the Pacific rim of North 

America, both on land and off shore. Accordingly, at the 

latitude of San Francisco, the system is approximately 80 

km wide, and at the latitude of San Diego approximately 

150 km wide (see fig. 1 and maps at front of book). 

The term "fault zone" refers to the complex zone of 

sheared rock that may be from 0.5 to more than 1 km 

wide and hundreds of kilometers long. The fault zone has 

developed over a period of millions of years while 

growing in width and complexity. The terms "fault," 

"fault branches," and "fault strands" refer to smaller 

elements and can be applied as needed. For example, 

surface rupture accompanying an earthquake commonly 

produces a complex pattern of fractures, and detailed 

elements can be discussed more effectively by using such 

terms as "fault branch, splay, or strand." The term "fault 

segment" recognizes that the fault is not completely 

continuous but is in sections or parts with poorly defmed 

boundaries, as discussed below in the subsection entitled 

"Segmentation." 

I FIGURE 1.1.-The San Andreas fault zone appears as a gash across the terrain in the Carrizo Plain area of south-central California. Streams are 

deflected or offset, as in foreground. Note that strata of Pleistocene age are tilted to steep angles on right (west) side of fault. View 

southeastward. 

3 

THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA; US. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
4 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

FIRST REFERENCES TO THE SAN ANDREAS FAULT 

The San Andreas fault first came into prominence only 

after it was fully understood by geologists as the cause of 

the great San Francisco earthquake of 1906. The name 

had been first used only 9 years previously by A.C. 

Lawson (1895) for a small segment of the fault on the San 

Francisco peninsula, where he reported that "a remarkably straight fault * * * has conditioned the San Andreas 

and Crystal Springs Valley" (p. 439). Lawson applied the 

name "San Andreas fault" almost incidentally in a discussion of "subsequent streams that flowed in the valleys." He suggested vertical displacement on the fault 

hut apparently was not convinced of that, inasmuch as he 

failed to show the fault or, with one exception, displaced 

strata on several cross sections in his report (for example, 

pi. 7). Clearly, neither the amount ofdisplacement on the 

fault nor its great regional and tectonic significance was 

appreciated at the time of Lawson's work. 

In one of the first reports about the 1906 earthquake, 

G.K. Gilbert (1907) accurately described the fault and its 

characteristic displacement, hut he did not use the name 

"San Andreas fault." In his field notes for April 28,1906, 

just 10 days after the great earthquake, Gilbert described 20 ft of right-lateral displacement of a road where 

it crosses the fault at the head of Tomales Bay. He had 

been following the surface rupture for several days, and 

on April 26 he recorded in his notes that along the west 

side of Bolinas Bay "some of the cracks were clearly 

secondary; others may have been primary." By the time 

the final report of the State Earthquake Investigation 

Commission was published (Lawson, 1908), the name 

"San Andreas" had been adopted, and its characteristics 

and role in causing the earthquake were clear. That 

report, which contains a remarkably extensive and accurate account, constitutes a major milestone in our understanding of the San Andreas fault and of strike-slip faults 

as a class. 

EXTENT AND PATTERNS OF FAULTING 

OVERVIEW OF THE PACIFIC MARGIN OF 

NORTH AMERICA 

The San Andreas fault system is part of a complex 

system of faults, isolated segments of the East Pacific 

Rise. and scraps of plates lying east of the East Pacific 

Rise that collectively separate the North American plate 

from the Pacific nlate (fie. 1.2). On a more eeneralized or 

global scale, theNorth American plate canbe thought of 

as lying across and partly covering the northern part of 

the Pacific system of plates. In simplified terms, the 

Pacific system of plates includes three elements: a 

westward expanding plate (the Pacific plate), an eastward-expanding plate (the Juan de Fuca plate), and a 

spreading center (the East Pacific Rise) from which the 

plates expand as new material is added. To the north, the 

Pacific plate is underriding, or being subducted under, 

the North American plate along the Aleutian thrust. 

Some investigators (Atwater, 1970; Atwater and Molnar, 1973) suggested that the North American plate has 

converged with and, indeed, slid over the Pacific system 

of plates, leaving only segments of the East Pacific Rise 

exposed, to which such names as "Juan de Fuca and 

Gloria Ridges" are applied. Similarly, related scraps of 

the eastward-expanding plate are the Juan de Fuca and 

Goria plates (fig. 1.2). Absolute plate motions derived by 

Minster and Jordan (1980). and Jordan and Minster 

(1988) are shown in figure 1.2, along with the platemotion vectors derived from the relative migration of 

mantle plumes or hotspots responsible for volcanic activity in Yellowstone National Park and the Hawaiian 

Islands. 

At its north end, the San Andreas fault joins the 

Mendodno Fracture Zone at a high angle, and there 

three plates are juxtaposed: one moving relatively northwestward, the second southeastward, and the third 

eastward and northeastward, to form a triple junction. 

At its southeast end, the San Andreas fault system 

merges more gradually with the set of transform faults 

underlying the Gulf of California. Just northwest of the 

area of merging, however, the trend of the San Andreas 

fault system changes to much more westerly, whereas a 

set of echelon faults accompanied by volcanism less than 

a million years old form a north-south-trending zone that 

extends northward across the Mojave Desert into Owens 

Valley of eastern California (Hill and others, 1985). This 

zone may be considered the East Pacific Rise overridden, 

and thus modified in pattern, by the North American 

plate. 

The San Andreas fault system may be viewed as 

forming the hypotenuse of a right triangle of which the 

northward extension of the East Pacific Rise and the 

eastward extension of the Mendocino Fracture Zone are 

the legs. The model of an overridden, suhducted oceanic 

plate within this triangle and underlying the North 

American plate (Dickinson and Snyder, 1979) presents 

signscant tectonic problems (see chap. 3). 

The San Andreas fault system has rearranged an 

assemblage of microplates, or terranes, some of which 

originated tens of degrees of longitude apart. During the 

fault's approximately 29-m.y. existence, an extremely 

complex pattern of rock distribution, has thus has been 

created (see chap. 3). 

MAJOR ELEMENTS OF THE SAN ANDREAS FAULT SYSTEM 

The San Andreas fault system consists primarily of the 

San Andreas fault and several major branches, such as 
1. GENERAL FEATURES 

PLATE 

\ 

EXPLANATION 

- L 

5.7 

- Strike-slip faultÃ‘Showin slip vectors and rate (cdyr) +-- Subduction zoneÃ‘Showin slip vectors and rate (cmlyr) of convergence; 

sawteeth on upper plate 

Spreading CenterÃ‘Showin vectors and rate (cmlyr) of spreading 

Vector and rite (cdyr) of absolute plate nmllon 

Hat spot and v8-0~ of plate motion rehtive la It 

-- Fracture zone in Pacific plate 

A Fault in North American plateÃ‘Sawteet on upper plate of thrust fault 

FIGURE 1.2.-Northeastern Pacific Basin, showing relation of the San AndreaB fault as one element in the complex boundary between the North 

American and Pacific plates. Modified from Drunimond (1981). 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

the Hayward and Calaveras faults in central California 

and the San Jacinto and Elsinore faults in southern 

California (fig. 1.3). In addition, in southern California 

the San Andreas fault splits into northern and southern 

branches in the eastern Transverse Ranges east of Los 

Angeles. These major faults accommodate about twothirds of the right-slip motion between the North American and Pacific plates. 

Numerous smaller branches, and extensions of segments of the fault, include in northern California such 

faults as the Rodgets Creek and Maacama faults, which 

may he considered northward extensions of the Hayward 

fault. The Green Valley and Bartlett Springs fault zones 

extend the Calaveras fault northward in a complex way 

(see maps at front of book). At the south end of the 

system, the Imperial fault represents a transition from 

the more continuous San Andreas fault to a more nearly 

echelon pattern characteristic of the faults under the Gulf 

of California. The Superstition Hills and Coyote Creek 

faults similarly represent a transition from the San 

Jacintofault to a more segmented pattern to the south in 

Mexico. 

In this volume, the San Andreas fault system is 

considered to lie principally within a belt about 100 km 

wide by 1,300 km long, but this boundary is arbitrary. 

Indeed, part of the relative strike-slip motion between 

the North American and Pacific plates seems to be taken 

up as far as 1,000 km east of the coastline throughout the 

Great Basin province (Jordan and Minster, 1988). The 

name "San Andreas fault system," however, should be 

confined to the more limited belt with the highest 

concentration of right-lateral strike slip. 

Sacramento 

. . . . . . . . . ... ..,... . .... 

0 - 50 100 KILOMETERS 

FIGURE 1.3.-The San Andreas fault system in California. Arrows on San Andreas fault (red) indicate direction of relative movement. 
1. GENERAL FEATURES '1 

PRINCIPAL TRENDS 

With some notable exceptions, the fault trends about 

N. SV40Â°W (fig. 1.3). Inits central section between the 

latitudes of San Jose and Bakersfield, the fault is 

relatively simple and straight, but farther to the south 

and north several branches splay from the main active 

trace. Near San Jose, where the fault bends about N. 60' 

W., the Calaveras and Hayward faults splay to the east 

and trend between N. 20Â°an %OW. South of the latitude 

of Bakersfield, the main fault changes most sharply in 

strike, in what commonly is referred to as the Big Bend 

reach of the fault. For 120 km or more the fault strikes 

about N. 60Â°W. where it bounds the Mojave block on the 

south. This bend has significant tectonic implications (see 

chaps. 2 and 3, and maps at front of book). 

South of the latitude of Los Angeles, the Elsinore and 

San Jacinto faults splay to the southeast, forming, in a 

general way, a reversed image of the splays in the San 

Francisco Bay region, although both faults trend about 

N. 60 W. 

COMPLEXITIES OF THE FAULT SYSTEM 

OTHER RELATED FAULTS 

In addition to the right-lateral strike-slip faults that 

characterize the San Andreas fault system, faults displaying left-lateral strike slip, as well as thrust faults and 

reverse faults of many sizes, are present (see maps at 

front of book). Normal faults are less common but are 

present in some places, for example, in zones of extension 

4 "*9 4'^ f "A 

:: ... . . . . . . . . 

... ... .. ... \ , , , , . ^. 8 *,- 8 '^6 "a. 

Faults dotted where concealed; sawteeth on upper plate of thrusts. Mendmino Fracture Zone queried where uncertain. 
8 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

at the crest of folds associated with the major faults, ir 

the bordering ranges, and at jogs in the fault where lo& 

extension is to he found. 

Most conspicuous of the faults displaying left-latera 

slip is the Garlock fault which intersects the San Andrea 

fault at about lat 3.5" N. and extends northeastward and 

eastward from there for 240 km (fig. 1.3; see maps at 

front of book). On the west side of the San Andreas fault, 

the southwest- and east-west-trending Big Fine fault 

joins the San Andreas fault a few kilometers northwest of 

the point where the Garlock fault joins the San Andreas 

fault. 

The left-lateral Pinto Mountain fault zone joins the San 

Andreas fault on its east side at about lat 34' N. and 

extends northeastward, in a pattern not unlike that of the 

Garlock fault. The Blue Cut fault is another left-lateral 

fault in the same general area. 

On a broad regional scale, thrust faults and detachments that accommodate subduction of the lower crust 

are significant, and various interpretations and speculations have been offered (fig. 1.4; Weldon and Humphrey~, 1986; Namson and Davis, 1988). Intennediatescale thrust faults that border the Transverse Ranges on 

the south side are characterized by such faults as the San 

Fernando fault zone (Grantz, 1971) and the Cucamonga 

fault zone (fig. 1.6). At the north end of the San Andreas 

fault system, where it joins the Mendocino Fracture 

Zone, such thrust faults as the Kings Range fault 

FIGURE 1.4.-Some interpretations suet that the San Andreas fault 

bottoms at a detachment fault and that numerous thrust faults, as 

well as the San An* fault itself, are important elements which 

accommodate the relative displacements between the North American and Pacific plates (from Namson and Davis, 1988). Heavy lines, 

faults; arrows indicate direction of relative movement: A, away from 

observer; T. toward observer. 

similarly accommodate crustal shortening. Smaller 

thrust faults that flank the San Andreas fault and dip 

toward it are common along many parts of the fault zone. 

A range of fault types, and the complexities that typify 

much of the fault zone, are well illustrated near Cajon 

Pass and southeast of there (fig. 1.5). where the San 

Andrea? fault zone splits into a northern and a southern 

branch. The strike of the San Gomonio - Pass fault zone in 

the east-central part of figure 1.5 changes in several 

places, and depending on the trend of a given segment, 

strike or dip slip may predominate. In this same area, the 

San Andreas fault, which to the northwest is relatively 

continuous and linear, ends as a surface feature, and the 

style of deformation changes from strike slip to primarily 

dip slip at the surface. Left-lateral slip characterizes the 

Pinto Mountain fault in the eastern part of figure 1.5, and 

small normal faults can be found throughout this area. 

SEGMENTATION 

Different behavior patterns along different parts of the 

fault began to be recognized when Steinbrugge and 

Zacher (1960) found that continuous slip or "creep" 

occurred along the fault in central California. Alien (1968) 

delineated five different regions along the San Andreas 

fault, three displaying seismic activity and two displaying 

little or no current activity. Wallace (1970) described in 

more detail major differences in behavior along different 

segments. 

Large segments of the fault system that are believed to 

produce damaging earthquakes are illustrated in figure 

1.6, which also shows an evaluation of the probabilities of 

earthquakes of different magnitude along these major 

segments of the San Andreas fault and three of its 

branches, the Hayward, San Jadnto, and Imperial faults. 

Both historical seismicity and paleoseismic evidence of 

large earthquakes and slip rates that characterize these 

different segments have been used in this assessment 

(Working Group on California Earthquake Probabilities, 

1988). 

Segmentation at a scale of a few to several kilometers 

is shown in figure 1.7, and a cumulative plot of the 

segment lengths in figure 1.8. The maximum length of 

these segments is about 18 km, but a more significant 

upper range appears to be near 10 km. These mappable 

segments are based on relatively fresh geomorphic features considered to he "young," that is, probably less 

than several thousand years old. 

Both left (fig. 1.7)- and right-stepping echelon patterns 

are displayed, and combinations or transition zones are 

also present. Complex patterns are common. 
e.. Fault-Dotted where conceaIe6 queried whem - Thrust fault-Dotted where concealed. Saw. '+ Location of sllp~ate delerminatien nnd 

uncenein. Paired arrows show relalive strike teeth on upper plate amount (mm) slip; single arrow shows dip of fault. Bar and 

ball are on downdropped block 

FIGURE 1.5.-Complex branching and changes in the style of faulting characterize the eastern branch of the San Andreas fault system in southern California. Although right-lateral 

strike slip characterizes the main San Andreas fault and its major branches, thrust faults, normal faults, and left-lateral faults are also present (modified from Matti and others, 

1985). 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

EXPLANATION 

Fir.l'~r 1.6.-Seimenta of the San Andre- fault ayfttem display different behavior. Here arc sliown condilKinal 

probabilities for the #iccuirenc-c of major etulhquakeri aloni; ~t.pmcntt of the San Andrew (A) and the Hnyward, 

San Jacimo, and Imperial faults ti31 ffi-the ^O-par interval from l!W t('201b. From WorkineCroup on California 

Earthquake probabilities (1988). 
1. GENERAL FEATURES 11 

0 - 1 2 KILOMETERS 

A. Quail Lake 

ti 

\ - 

8. Yucaipa quadrangle A 

C. Gualala River 

- \ -- 

'\, Ã‡ ,;-^---- 

0. Palmdale Reservoir 

.? 

E. Panorama Hills + 

\ - \ v...... 

......... F Paicines + - 

Y 

\ 

v ....... 

G Bear Valley M 

FIGURE 1.7.-Individual branches and H. Frazier Park 

strands of surface trace ofthe San Andre- as fault are amneed in various patterns: 

+ 

A-C, left-stepping echelon &ment; 

D-F, right-stepping echelon arrangement; G, both left- and right-stepping 

arrangements; H-J, complex arrangements of individual strands. Note that 

individual strands may be at an angle of as 

much as 12Â to general trend of fault zone. I. Elkhorn Hills 

Sources: A, 5 (RoM, 19691; B (Hope, 

1969; see references in this chapter); C, J 

(Brown aid Wolfe, 1972); E, H. I (Vedder 

and Wallace, 1970); F, G (Brown, 1970, 

see references in chap. 21. 

J. Point Reyes peninsula 
12 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

1 

0 2 4 6 8 10 12 14 16 18 

SEGMENT LENGTHS, IN KILOMETERS 

FIGURE 1.8.-Census of lengths of individual segments of the 

Sail Andreas fault. Maximum lengths are in the range 10-18 

km. Line is fitted by eye to show trend of values. From 

Wallace (1973). 

REFERENCES CITED 

Allen, C.R., 1968, The tectonic environment of fleismieaUy active and 

inactive areas along the San Andreas fault nyatem. in Diekinson, 

W.R., and Grantz, Arthur, eds., Proceedings of conference on 

geologic problems of San Andreas fault system: Stanford. Calif., 

Stanford University Publications in the Geological Sciences, v. 11, 

p. 70-80. 

Atwater, Tanya, 1970, Implications of plate tectonics for the Cenozoic 

tectonic evolutionofwestern North America: Geological Society of 

America Bulletin, v. 81, no. 12, p. 3513-3635. 

Atwater, Tanya, and Mohiar, Peter, 1973, Relative motion of the 

Pacific and North American plates deduced from sea-floor spreading in the Atlantic, Indian and South Pacific Oceans, in Kovach, 

R.L., and Nur, Amos, eds., Proceedings of the conference on 

tectonic problems of the San Andreas fault system: Stanford, 

Calif., Stanford University Publications in the Geological Scienc8, V. 13, p. 1S6-148. 

Drummond, K.J., chairman, 1981, Pacific Baain aheet of Plate-tectonic 

map of the circum-Pacific region: Tulsa, Okla., American Association of Petroleum Geologists, scale 1:10,000,000. 

Dicldinon, W.R., and Snyder, W.S., 1979, Geometry oftriple junction 

related to the San Andreas transform: Journal of Geophysical 

Research, V. 84, no. B2, p. 561-572. 

- 1979, Geometry of subducted slabs related to the San Andreas 

transform: Journal of Geology, v. 87, no. 6, p. 60&-027. 

Gilbert, G.K., 1907, The earthquake as a natural phenomenon, in The 

San Francisco earthquake and fire of April 18, 1906, and their 

effects on structures and structural materials: U.S. Geological 

Survey Bulletin 324, p. 1-13. 

Grantz, Arthur, 1971, The San Fernando, California, earthquake of 

February 9, 1971, Introduction, in Grantz, Arthur, ed.. The San 

Fernando, California, earthquake of February 9, 1971: U.S. 

Geological Survey F'mfemional Paper 733, p. 14. 

Hill, D.P., Wallace, R.E., and Cockerham, R.S., 1985, Review of 

evidence on the potential for major earthquakes and volcanism in 

the Long Valley-Mono Craters-White Mountains regions of Eastem California: Earthquake Prediction Research, v. 3, no. 9-4, p. 

671-594. 

Hone. . . R.A.. 1969. Man shod recentlv active breaks alw the San 

And& andrelated faults between Cajon Pass and <lton Sea, 

California: U.S. Geological Survey open-file map. scale 124.000.2 

sheets, 

Jordan, T.H., and Minster, J.B., 1988. Measuring cruatal deformation 

in the American West: Scientific American, v. 259, no. 2, p. 48-68. 

Lawson, A.C., 1895, Sketch of the geology of the San Francisco 

peninsula: U.S. Geological Survey Annual Report 15, p. 439473. - chairman, 1908, The California earthquake of April 18, 1906: 

Report of the State Earthquake Investigation Commission: Carnegie Institution of Washington Publication 87, 2 v. 

Matti, J.C., Morton, D.M., and Cox, B.F. 1985, Distribution and 

geologic relations of fault systems in the vicinity of the central 

Transverse Ranges, southern California: U.S. Geological Survey 

Open-File Report 85-365,27 p., scale 1:250,000,2 sheets. 

Minster, J.B., and Jordan, T.H., 1980, Present-day plate motions: A 

summary, in Source mechanism and earthquake prediction (Coulomb volume): Paris, Centre National de la Recherche Scientifiaue. D. 109-124. 

~amfioi. J.s., and ~avis. T.L., 1988. Structuraltransectofthewestern 

Transverse Ranges, California: Implications for lithospheric kinematics and seismic risk evaluation: Geology, v. 16, no. 8, p. 

675-679. 

Steinbrugge, K.V., and Zacher, E.G., 1960, Creep on the San Andreas 

fault. Fault creep and property damage: Seismological Society of 

America Bulletin, v. 50, no. 3, p. 389-S96. 

Wallace, R.E., 1970, Earthquake recurrence intervals on the San 

Andreas fault: Geological Society of America Bulletin, v. 81, no. 

10, p. 2876-2889. 

- 1973, Surface fracture patterns along the San Andreas fault, in 

Kovach, R.L., and Nur, Amos, eds., Proceedings of the conference on tectonic ~mblema of the San Andreas fault avstem: 

Stanford, Calif., tanf ford University Publications in the &logical Sciences, v. 13, p. 248-250. - 1975, The San Andreas fault in the Canh Plain-Temblor Range 

reeion, California, in Crowell, J.C.. ed., San Andreas fault in 

southem California: A mide to San Andreas fault from Mexico to 

Carrizo Plain: califom; Division of Mines and Geology Special 

Report 118. p. 241-250. 

Weldon, R.J., and Humphreys, E.D., 1986, A kinematic model of 

Southern California: Tectonics, v. 6, no. 1, p. 33-48. 

Working Group on California Earthquake F'mbabilitiea, 1988, Probabilities of large earthquakes occurring in California on the San 

Andreas fault: US. Geological Survey Opal-File Report 88-398, 

62 P. 
1 

Sacramento 

T he San Andreas fault is marked in the landscape by a series of 

linear valleys and mountain fmnts, aligned lakes and bays, 

elongate ridges, and disrupted or offset stream channels. 

2. GEOMORPHIC EXPRESSION 

CONTENTS 

REGIONAL FEATURES 

On maps, aerial photographs, or satellite images, at 

almost any scale, the San Andreas fault zone appears as 

a linear scar across the landscape. At scales small enough 

to display the entire fault length, valleys, bays, chains of 

lakes and ponds, linear flanks of mountain ranges, and 

elongate ridges bounding one side or the other of the fault 

are the principal features that reveal its location. 

Erosion of the softer broken and sheared rocks in the 

several-hundred-meters- to 1-km-wide fault zone accounts for much of the valley-like expression of the fault, 

but differential vertical displacements also play a major 

role. The ratio of local horizontal to vertical displacement 

may be about 10 or 20 to 1. Differential erosion of the 

various rock types juxtaposed by faulting also influences 

the geomorphic expression of the fault. 

The San Andreas fault separates some mountain masses from adjacent broad regions of low relief, creating 

pronounced linear topographic discontinuities. Considering the large lateral displacements that have occurred, 

juxtaposition of some mountains against flatter, lower 

areas probably has come about by lateral slip. Some 

range-size blocks bounded by the fault, however, have 

risen or dropped hundreds of meters to create the lmear 

topographic features. Elongate blocks of the crust bounded by branches and subparallel strands of the fault have 

been created and shuffled one block against another by 

both upward, downward, and lateral differential displacement within the broad shear zone. Such movements 

are reflected in the topography as elongate ridges and 

depressions. Whether lateral or vertical block displacement, or warping or folding, has dominated in the 

development of a specific landform, and what role erosion 

has played, have yet to be well analyzed for most 

topographic features within the fault system. Evidently, 

a complex interaction of tectonic, erosional and depositional processes has influenced the development of each 

feature, and the result is a linearity of topographic 

features along and parallel to the San Andreas fault. 

From Point Arena southeastward to the vicinity of San 

Jose, Calif., the trace of the San Andreas fault is 

topographically conspicuous on regional maps and images 

as a series of lmear valleys (see fig. 2.1 and maps at front 

of book for locations). Aligned linear valleys also mark 

the fault trace throughout central California between San 

Jose and the Carrizo Plain, but along that reach another 

important characteristic is that the fault trace crosses 

mountain ranges and major ridges at a low angle. These 

FIGURE 2.1.-Shaded-relief map of California and western Nevada, Range province, are myriad block-faulted mountains separated by 

showing the San Andreas fault as a series of linear valleys and ridges intervening basins. Synthesized from digital terrain models of a 

passing through a mountainous region, the Coast Ranges, the trends l:250,000-scale map series originally prepared by the U.S. Defense 

of which are a low angle to the fault. To theeast, between the massive Mapping Agency, now maintained by the U.S. Geological Survey; 

Sierra Nevada and the Coast Ranges, lie the broad, flat valleys of the compiler, Ray Batson, U.S. Geological Survey. Illumination is from 

Sacramento and San Joaquin Rivers. Farther east, in the Basin and north at elevation of W. 

THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA: U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
16 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

mountain ranges and ridges, many of which are antiformal structures, trend from 5' to 10' more westerly than 

the strike of the fault. Thus, the altitude of the surface 

trace of the fault alternately rises and falls along strike. 

In the Carrizo Plain-Temblor Range area, the surface 

trace of the fault does not lie at the base of the range but 

more within the Carrizo Plain, where the surface expression of the fault is narrowest. clearest. and best defined. 

Offset streams are especially well preserved here; individual strands of the fault reach a maximum length, from 

9 to 18 km, anywhere along the fault. 

In the Big Bend area at the south end of the San 

Joaquin Valley, the fault trace rises to a high altitude as 

it passes through mountainous terrain. Along the Mojave 

segment to the northwest and southeast of Palmdale (see 

maps at front of book for locations), the fault trace is 

again marked by a distinct narrow, linear valley. In 

addition, a gross contrast between the high, rugged 

mountain masses of the Transverse Ranges and the 

relatively flat Mojave Desert block is apparent (fig. 2.1). 

The surface of the Mojave Desert itself stands 700 m or 

more above sea level and above the San Joaquin Valley. 

Clearly, the Mojave block has been uplifted, even though 

the adjacent mountain masses have risen more. 

To the southeast of the Mojave segment, the San 

Andreas fault crosses the Transverse Ranges at a low 

angle and separates the high San Gabriel Mountains from 

the San Bernardino Mountains. An extremely complex 

structural knot, formed by branching of the San Jacinto 

fault and numerous other faults (fig. 2.1; see fig. 1.5) 

is reflected as a complex topographic region surrounding 

the Cajon Pass area, through which the San Andreas 

fault passes. 

Southwest of the Cajon Pass area, the fault divides into 

a northern and a southern branch and numerous other 

smaller faults of different tectonic style. Each fault has 

its own distinctive geomorphic expression. 

Southeast of the zone of major branching, the fault 

again is less conspicuously marked by contrasts of large 

topographic features, but it is readily visible on aerial 

photographs at scales of 1:60,000 and larger (see section 

below entitled "A Photographic Album of Fault Features"). 

LOCAL GEOMORPHIC FEATURES 

WITHIN THE FAULT ZONE 

Within the fault zone, various geomorphic features are 

found that have their origin in both the lateral and 

vertical shuffling of fault-bounded slices, as well as in the 

persistent, large strike slip. These features include sag 

depressions and sag ponds, shutter ridges and medial 

ridges, offset and deflected stream channels, linear 

benches along valley walls, aligned notches and saddles 

on spurs, offset marine and river terraces, scarps, 

fault-controlled drainage, and folds and pressure ridges 

(fig. 2.2). 

Along its entire length, the fault zone exhibits peculiar, 

anomalous drainage patterns. In regions where tectonic 

activity is less pronounced, streams generally flow more 

or less perpendicular to mountain blocks and highlands, 

and grade more or less regularly to the lowlands; not so 

along an active fault like the San Andreas. When 

drainage flowing from highlands meets the San Andreas 

fault, it is diverted subparallel to the trends of the 

highlands or is interrupted or blocked completely. In less 

active areas, erosion generally is the dominant factor in 

carving geomorphic forms, but displacements are so 

rapid within the fault wne that tectonic effects overwhelm erosion, and so the geomorphic features directly 

express fault movement. 

Movement within the network of branching and anastomosing fault strands jostles the intervening blocks, 

compressing some, rotating some, or causing extension 

across others. Because the principal slip is horizontal and 

lateral, the blocks tend to he elongate parallel to the 

trend of the fault. Blocks under compression tend to be 

squeezed upward to form elongate ridges, whereas 

blocks under extension may drop downward to form sags, 

and laterally displaced slices or ridge spurs create 

shutters across drainage channels. 

The dominantly lateral slip across the fault zone and 

the rate of slip, from 1 cm to a few centimeters per year, 

make stream channels that are offset right laterally, a 

common and characteristic geomorphic feature. Stream 

channels can be completely beheaded or merely offset 

while maintaining continuity of flow. 

In addition to the effects of lateral slip, streams are 

extremely sensitive to vertical slip on faults and warping 

of the land surface. For example, only a small upward 

movement of a block on the downstream side of a fault 

crossing a stream may divert the stream either to the left 

or right, thus mimicking lateral slip on the fault. Similarly, warping of the land surface over folds adjacent to 

the fault or on pressure ridges within the fault zone can 

distort the patterns of streams. Combinations of these 

different tectonic processes can produce many unusual 

features. Both the tectonic and the erosional changes at 

times may occur almost instantaneously, and so the 

dominance of one or the other process suddenly may 

change. Between such periods of sudden change, very 

little may happen for decades or even centuries. The 

relative rates of erosional and tectonic processes, and the 

timing of sudden events, are critical to the landforms 

created. Some of the patterns of streams found in the 

Carriw Plain area are illustrated in figure 2.3, and an 

example is shown in figure 2.5. 
2. GEOMORPHIC EXPRESSION 17 

The geomorphic forms created represent the results oi 

a continuing contest between erosional changes and 

changes related to fault slip, folding, and warping. Where 

streams are large and rainfall is greater, only displacements of hundreds of meters or more are preserved for 

longer than a few centuries. In desert climates, however, 

as in the Carrizo Plain, the rate of fault slip outruns 

erosion, and the effects of only a few meters of fault slip 

may be preserved for hundreds of years, if not millennia, 

where small channels cross the fault (fig. 2.4). 

As an example of how erosion and sedimentation 

interact with the faultingprocess, a straight channel that 

formerly crossed the fault at right angles is shown after 

having been offset by right-lateral strike slip (fig. 2.5). 

The strike slip partly or temporarily dams the stream, 

causing upstream alluviation at C. A fresh fault scarp is 

formed in the vicinity of A, and successive offsets expose 

Linear valley Offset drainage channel I I 

new scarp areas to the left of A. The dam at B is eroded, 

and the alluvium deposited earlier at C is dissected. As 

offset progresses further, the channel segment along the 

fault trace, between B and A, continually elongates, thus 

lowering the channel gradient more and more. Because of 

this decreasing gradient, alluvium is deposited upstream 

from A to and beyond C, and eventually the stream, 

having difficulty maintaining a channel along that elongate course, spills across the fault trace and creates a new 

channel more nearly in alignment with the segment 

upstream from the fault. 

After fault movement has progressed sufficiently, the 

downstream segments of other channels are brought into 

alignment, or nearly so, with the original channel. For 

example, in the vicinity of D in figure 2.5, drainage 

flowing to the right in an adjacent channel would tend to 

erode headward toward C, and capture of the original 

Offset drainage channel 

< 

FIGURE 2.2.-Common landforms along the San Andm fault system (from Vedder and Wallace, 1970). 
18 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

0 

u 

400 METERS EXPLANATION 

Approximate scale - Streamchannel 

........ - Abandonsd channel 

Â¥M Ridge or spur 

Ã‘.......... Fault 

&RE 2.3.-Diagrammatic representation of patterns of fault-related 

stream channels found in the C& Plain area (from Wallace, 1975): 

A, misalignment of single channeb directly related ta amount of fault 

displacement and age of channel-no ridge ondownslope side of fault, 

beheading common; B, paired stream channels misaligmed; C, compound offsets of ridge spurs, and offset and deflection of channels, 

both right andleft deflection; D, trellisdrainage produced by multiple 

fault strands, sliver ridges, and shutter ridges; E, exaggerated or 

reversed apparent offset, caused by offset plus deflection by shutter 

ridge; F, "Z" pattern, caused bycapture by adjacent channel followed 

by right-lateral slip; G, false offset caused by differential uplift or 

warping; H, false otfeet caused by echelon fractures over fault zone, 

followed by subsequent streaniflow. 

~URE 2.4.-Topographic map of a segment of the San Andreas fault 

in the Carrizo Plain area, showing some characteristic small-scale 

mmorphic features. Markers (SAF) at the left and rkht meina 

indicatemain fault trace. Between300and500m(see .~~~~~ scale ~ at bottom) ~~ ~~ . is one of the best examples of a stream offset by right-lateral slip on the San Andreas fault (see fig. 2.20). To the northwest, between 100 

and 300 m, is an abandoned channel of the same stream. To the 

southeast, between 600 and 700 m, small streams are offset about 10 

m: a few of these streams record multiple offsets of about 8 to 10 m. 

The last offset presumably was du& the great Fort Tejon earthquake of 1857. Between 2.150 and and 2,250 m is a pair of all-eams 

that has been offset a few tensof meters. At 2,000 m, the do-tream 

segments of those two streams have been beheaded completely, and 

at 2,100 m is another possible beheaded segment of the streams (see 

fig. 2.21). Between 2,300 and 2,700 m is a sag depression about 7 m 

deep. multine {rom downdrm of a narrow block into the San 

~ndreas fault Lne. From Sieh Ad Wallace (1987); map preparedby 

U.S. Geological Survey from aerial photographs taken January 13, 
2. GEOMORPHIC EXPRESSION 19 

stream would take place. The gradient of this capturing 

segment, in which flow is to the right, progressively 

increases because right-lateral slip shortens the channel, 

thus accelerating erosion; at the same time, the channel 

flowing to the left elongates, gradient decreases, and 

erosion decelerates. An example of some geomorphic 

features that result is shown in figure 2.5. 

DETAILED MAPS OF THE FAULT SYSTEM 

The U.S. Geological Survey and the California Division 

of Mines and Geology have prepared numerous detailed 

maps of the faults within the San Andreas fault system. 

Maps prepared by the California Division of Mines and 

Geology address the problem of "active faultsn as defined 

under the Alquist-Priolo Special Studies Zones Act of 

1972. These maps, too numerous to list here, were 

indexed and described by Hart (1985); they constitute a 

rich data set about the San Andreas and other faults in 

California. Those readers interested in examining the 

features of the fault system in more detail or in the field 

are referred to the published "strip" maps and special 

fault maps, an index map of which is in figure 2.6. 

REFERENCES CITED 

Bortugno, E.J., 1982, Map showing recency of faulting, Santa Rosa 

Quadrangle, 1:260,0(10: California Division of Mines and Geology 

Regional Geologic Map Series, no. 2A, sheet 5. 

Brown, R.D., Jr., 1970, Map showing recently active breaksalong the 

San Andreas and related faults between the northern Gabilan 

Range and Cholame Valley, California: U.S. Geological Survey 

Miscellaneous Geologic Investigations Map 1-575, scale k62.500. 

- compiler, 1972, Active faults, probable active faults, and associated fracture zones, San Mateo County, California: U.S. Geolop 

FIGURE 2.5.-General featurea and conditions produced where a 

stream channel is offset by strike slip on a fault. See text for 

discussion of A-D. 

ical Survey, San Francisco Bay Region Environment and Resources Planning Study Basic Data Contribution 44 (Miscellaneous 

Field Studies Map MF-355) scale 1^2,600. 

Brown, R.D., Jr., and Wolfe, E.W., 1972, Map showingrecently active 

breaks along the San Andreaa fault between Point Delgada and 

Bolinaa Bay, California: U.S. Geological Survey Miscellaneous 

Geologic Investigations Map 1-692, scale 124,000, 2 sheets. 

Clark, M.M., 1973, Map showing recently active breaks along the 

Garlock and associated faults, California: U.S. Geological Survey. 

Miscellaneous Geologic Investigations Map 1-741, scale L24.000. 

- 1982. ~ ~ Man showinc- recentlv active breaks alone the Elainore and . r P - 

associated faults, California, between Lake Henshaw and Mexico: 

U.S. Geological Survey Miscellaneous Investigations Series Map 

1-1329, scale 1:24,000. 

- 1984, Map showing recently active breaks along the San Andreas 

fault and associated faults between Salton Sea and Whitewser 

River-Mission Creek, California: U.S. Geological Survey Miacellaneous Investigationn Series Map 1-1483, 6 p., scale 134,000,2 

sheets. 

Hart, E.W., 1985, Fault-rupture hazard zones in California: AlquistPriolo Special Studio Zone8 Act of 1972 with index to sped 

study zones maps: California Division of Mines and Geology 

Special Publication 42, 24 p. 

Herd, D.G., 1977, Geologic map of the Las Posits, Greenville, and 

Verona faults, eastern Alameda County, California: U.S. Geological Survey Open-File Report 77-689,25 p., scale 1:24,000. 

Herd, D.G., and Helley, EJ., 1977, Faults with quaternary displacement, northwestern San Francisco Bay region, California: U.S. 

Geological Survey Miscellaneous Field Studies Map MF-818, scale 

k126,oOo. 

Kahle, J.E., 1975, Recent fault features and related geology, Leona 

Valley area, southern California, in Crowell, J.C., ed., San 

Andreas fault in southern California: A guide to San Andreas fault 

from Mexico to Carrizo Plain: California Division of Mines and 

Geology Special Report 118, p. 203-207. 

Lawson, A.C., chairman, 1908, The California earthquake of April 18, 

1906: Report of the State Earthquake Invention Commission: 

Carnegie Institution of Washington Publication 87, 2 v. 

Matti, J.C., Morton, D.M., and Cox, B.F., 1985, Dirtribution and 

geologic relations of fault systems in the vicinity of the central 

Transverse Ranges, southern California: U.S. Geological Survey 

Open-File Report 85-366, 27 p., scale 1260,000,2 sheets. 

McCulloch, D.S., 1987, Regional geology and hydrocarbon potential of 

offshore central California, in Scholl, D.W., Grantz, Arthur, and 

Vedder, J.G., eds., Geology and resource potential of the continental ~ ~ marrin of western North America and adiacent ocean 

basins-~eaufort Sea to Baja California (Earth Science Series, v. 

6>' Houston, Tex., Circum-Pacific Council for Energy and Mineral 

Resourcet, p. 369-W2. 

McLaughlin, R.J., 1971, Geologic map of the Sargent fault zone in the 

vicinity of Mount Madonna. Santa Clara County. California: U.S. 

Geological Survey, San Francisco Bay ~egion~nvironment and 

Resources Planning Study Basic Data Contribution 13, scale 

1:12,000,2 sheets. 

Pampeyan, E.H.. Hamh, P.W., and CoaBey, J.M., 1981, Preliminary 

map showing recently active breaks along the Maacamafault zone 

between Hopland and Laytonville, Mendocino County, California: 

U.S. Geological Survey Miscellaneous Field Studies Map 

MF-1217, 9 p., scale 1:24,000, 2 sheets. 

Radbruch, D.H., 1967, Approximate locationoffault tracesand historic 

surface ruptures within the Hayward fault zone between San 

Pablo and Warm Springs, California: US. Geological Survey 

Minctllanecnu Geologic Invaitigations Map 1-622, scale 1:62,500. 

Radbruch-Hall, D.H., 1974, Map showing recently active breaks along 

the Hayward fault zone and the southern part of Calaveras fault 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

0 - 1 200 KILOMETERS 

Veddw and Wall 

dark. 1982 otters. 1982 

FIGURE 2.6.-Index map of California, showing locations of selected maps of surface traces of the San Andreas fault 

system. 

zone, California; U.S. Geological Survey Miscellaneous Geologic 

Investigations Map 1-813, scale 1-24,000,2 sheets. 

Ross, D.C., 1969, Map showing recently active breaks along the San 

Andreas fault between Tejon Pass and Cajon Pass, southern 

California: U.S. Geological Survey Miscellaneous Geologic Investigation8 Map 1-653, scale 124,000. 

Sarna-Wojcicki, A.M., Pampeyan, E.H., and Hall, N.T., 1975, Map 

showing recently active breaks along the San Andrean fault 

between the central Santa Cruz Mountains and the northern 

Gabilan Range, California: U.S. Geological Survey Miscellaneous 

Field Studies Man MF-6Sf). wata 1- 24. 000. 2 sheets. 

Sharp, R.V.. 1972, M& showing recently active breaksiong the San 

Jacinco fault zone between the San Bemadno area and Borrego 

Valley, California: U.S. Geological Survey Miscellaneous Geologic 

Investigations Map 1-675, scale 124,000, 8 sheets. - 1973, Map showing recent tectonic movement on the Concord 

fault, Contracostaand Solano Counties, California: U.S. Geological Survey, San Francisco Bay Region Environment and Besow 

ces Planning Study Basic Data Contribution 55, (Miscellaneous 

Field Studio Map MF-505). scale 1 24, 000. 

Sharp, R.V., Lienkaemper, J.J., Bonilla, M.G., Burke, D.B., Fox, 

B.F., Herd. D.G., Miller. D.M., Morton. D.M., Ponti, D.J., 

Rwner. .. M.J.. . Tinslev. J.C.. Yount. J.C.. Kahle. J.E.. and HÃ§rt .. . . . ,~ . ~ ~, 

E.W., 1982, Surface faultingin thecentral Imperial Valley, in The 

Imperial Valley, California. earthquake ofOctober 15, 1979: U.S. 

Geological Survey Professional Paper 12'34, p. 119-144. 
2. GEOMORPHIC EXPRESSION 21 

Sieh. . K.E.. and Jahm. R.H.. 19%. Holom activitv of the !Jan 

~n& fault st ~a~~ace ~rek,'ca~ifomia: WI~&~I %cietyef 

America Bulletin, v. 95, no. 8, p. 8%34%. 

Sieh, K.E., and Wallace, R.E., 1987. The San Andm fault at WaUw 

Creek, San Luis Obisp County, Califomix Wlogiml Sdety of 

America, CodUem &tion Centend Field Guide. D. 23%238. 

Steinhm~, K.v., and Zwher, E.G., 1950. Creep on tk sin Andma6 

fault (California)-fault mp and property damage: ~mological 

Society of Amerim RuUeth, v. 50, no. 3, p. 389-395. 

U.S. Geological Survey, lW, The Imprial Valley, California, earthquake of October 15.19W hfesaional Pap= 1254,451 

Vedder, J.G., and Wallace, R.E., 1970, Map &owing recently dve 

breaks along the SanMms and dakd faults between ~olm 

Valley and Tejon Pam Cnlifomia: US. Wlogical Smy Mi8cellaneous Geologic lnvestigatiom Map 1574, stale P94,OCO. 

Wallwe, R.E., 1975, The Sm Andreas fadt in the Carrim PlainTemblor Range region, California, in Cmwell, J.C., &., San 

Andrew fault in southern California: Aguideta San Andrem fault 

hm Mexico to Carrim Plain Calif-* Division of Mines ad 

Geolow Special Report 118, p. 241-250. 

Yerkes, R.F., Green, H.G., Timley, J.C., and L@.joie, K.R., lH0, 

Maps ahowing aeismotectanic stting of the Santa Barb ChanXI area, Californix U.S. Geological Smey MiieUanenus Field 

Staxlies Map MF-1169, !25 p., scale k%O,OOO. 

PHOTOGRAPHIC ALBUM OF 

FAULT FEATURES 

the lwtions of thephotapphs. All photographs are my own except m~ otherwise credited in the 

captions. 
THE SAN ANDREA5 FAULT SYSTEM, CALIFORNIA 

124- 120. 118. 

I I I I 
2. GEOMORPHK EXPRESSIOK 25 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

Fl~u~~2.9.-A scarp facing the Pacific- Ocean, with marine terrace above, marks trace of the San Andreas fault at Boilcga Bay, 70 km north 

of San Francisco. View northwestward. 
2. GEOMORPHIC EXPRESSION 27 

FIGL'HE 2.10.-Tomales Bay lies in an clongitti-' (l~-pr~wsion along the marshy sediment. The Point Reyes peninsula is to right (west); 

San Andreas fault. In 1906, amaximum strikeslip of fi m on the fault Mount Tamalpais is high-point in skyline at top left. View southeastwas reported at the head of Tomales Bay, mid-distance in view ward toward San Francisco, 55 km away. 

however, some of this slip may have been related to shifting of 
28 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE 2.11.-A broad valley marks trace of the San Andreas fault and aliened with line of view and with tape in foreground' fence on 

on the Point Reyes peninsula. During the great San Francisco near (east) side bows gradually to meet the fault. View southwestearthquake of 1906 , slip on one strand of the fault offset a fence ward along fence in southern part of the Strain Ranch (Lawson, 

about 2.5 m. Fence beyond break (west side) is relatively straight 1908, p. 70-71; Brown and Wolfe, 1972). 
2. GEOMORPHIC EXPRESSION 29 

FIGLIKE 2.12.-Along the San Andreas fault zone near San Francisco. ni.s]>laccrnent in 1906 was along linear trace on left (wc'st) side of the 

Crystal Springs Reservoir (foreground). To northwest, active trace of fault bends to east side of the San Andreas Reservoir, which lies just 

beyond the Crystal Springs Reservoir. View northwestward. 
30 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

--A -. . -- 

FIGURE 2.13.-The San Andreas fault at San Juan Bautista. In 1906, southeast end of surface rupture wasjust north of here. Northwest limit 

of creeping segment of fault is now at about the same point. Active trace of fault lies along scarp at northeast (right) edge of town, which 

is located on a tectonic upwarp along the fault. View northwestward. 
2. GEOMORPHIC EXPRESSION 31 

Ficin~ 2.14. -Fault creep on the San Andreas fault was first reported at the Paicines Winery south of San Juan Bamista in I960 by 

Steinbmgge and Zaclier (IWiO). Fault trace (between arrows) passes under winery building, and walk (if building ait; continually being 

displaced. View northwestward, 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE 2.16.-The San Andreas fault about 40 km south of San Juan Bautista. Here. the Calaveras and San Andreaa faults join The Gabilan 

Range is on left; the San Benito River meanders across fault zone. View northwestward. 
2. GEOMORPHIC EXPRESSION 33 

FIGL-RE 2.16.-Tho San Andreas fault about 70 krn wuth of San Juan Bautista. The Diablo Rangc is on right (cast); the Gabilan Range, 

displaying a broad, arched upland surface, is on skyline at upper left. View northwestward. 
34 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE~.~~.-F~O~ 65 to 80 km northwest of Cholamc, trace [if the San Andreas fault follows crest of a ridce. Fault trace is marked by a series 

of s poi, in seemingly anomalous position near crest of a major elongate topographic high. View southeastward from California Highway 198. 
2. GEOMORPHIC EXPRESSION 36 

FIGURE 2.18.-Trace of the San Andrea fault in the Palo %eta Pass, about 18 krn southeast of Cholame. A sag pond is in forepound. View 

southeastward. 
36 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

FIGURE 2.19.-A pben locally marks trace of the San Andreas fault near north end of the Carrizo Plain. The Caliente Range is on horizon 

in upper right. View southeastward. 
2. GEOMORPHIC EXPRESSION 37 
38 THE SAN ANDREA6 FAULT SYSTEM, CALIFORNIA 
2. GEOMORPHIC EXPRESSION 

--'-.... . 

FIGURE 2.22.-Stream channel offset by the San Andreas fault. Southwest (far) side is raised along a fault sliver, which tends to preserve 

incised position of stream. Mound topography is believed to be ancient modified ground-squirrel burrows. View southwestward toward 

the Carrizo Plain. 
40 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGL-RE 2.23-A signwidal pattern of stream channel iwultn from right-lateral offset, followed hy stream captuw producing left deflection, 

followed by further right-lateral offset. Fault extends from right to left; stream flows from bottom to top. View southwestward toward 

the Carrizo Plain. 

42 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 
2. GEOMORPHIC EXPRESSION 43 

Fi~u~~2.26.-The San Andrew fault on east edge ufthe Cai-rizo Plain. Rugged, dissected terrain west Oeft) of fault trace is the Elkhorn scarp. 

The Elkhorn Plain is to rieht in distance (see Vedder and Wallace, 1970). 
44 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE 2.27.-A complex expression of the San Andreas fault forms west flank of the Elkhorn Hills. IJpwaqring of hills, in combination with 

echelon fractures along the San Andreas fault zone, helped control drainage. View southeastward. 
2. GEOMORPHIC EXPRESSION 45 

FIGL-RE 2.B.-In the Biz Bend region of the San Andreas fault, the upper left corner of photograph. Light-coloi-"(1 mountain just west of 

Garlock fault, which has left-lateral dip, branches eastward. June- fault junction is Fraser Mountain, an isolated thrust block placed 

tion of these two great faults is near center (circle); trace of the there by compression across the Big Bend. U.S. Air Force photoGarlock fault extends to left (east), and that of the San Andreas fault graph, taken November 27, 19G7. 

(between arrows) extends east-southeastward from lower right to 
46 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE~.~P.-T~~ Palmdale Reservoirliw dong the San Andreas fault north ofthe Kan Gahriel Mountains. The Mojave Desert is to northeast 

(right), and a series of ridges lies between the fault and the desert basin. View northwestwad. 
2. GEOMORPHIC EXPRESSION 47 

FIGL-RE 2.20.-The San Andreas fault on northeast flank (if the San Gabriel Mountains. View northwestward from Jackson Lakc (foreground) 

toward Palmdale. 
48 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE 2.31.-Northern branch of the San Andreas fault in northern Coachella Valley. The San Gorgonio Pass is at left horizon, and the Sam 

Bemarriino Mountaim arc at center of horizon. View northwestward. 
2. GEOMORPHIC EXPRESSION 49 

FIGURE 2.32.-The San Andreas fault (between arrows) in northern Coachella Valley, disrupts drainage, offsetting and beheading many 

channels. Mount San Jacinto is in center distance; the San Gorgonio Pass, afault-controlled pass, is at right in thedistance. View westward, 

U.S. Geological Survey photograph, taken November 15, 19%. 
50 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE 2.33.-The San Andreas fault in badlands of the Mecca Hills, Coachclla Valley. Sympathetic slip (if as much as I+ cm occurred here 

during the Borrego Mountain and Imperial Valley earthquakes of 1W and 1979, respectively. 
2. GEOMORPHIC EXPRESSION 51 

FIGURE 2.34.-TheSan Andreas fault near the Salton Sea, Imperial Valley. Salt C'i-eekis inforeground; Bat CaveButtes aredark hills at left, 

and the Salton Sea is in distance. Y~ew sraithaaatward. 
52 THE SAX ANDREAS FAULT SYSTEM. CALIFORNIA 

- - - - - - -- - - - . . . . - . - -- - -- -a& 

FIGI~ 2.35.-The Coyote Lake Reservoir lies in a topographic trough along the Calaveras fault a few kilometers east of Morgan Hill, Calif. 

View eastward across fault, 
2. GEOMORPHIC EXPRESSION 53 

FIGCRE 2.36.-Creep of about 1 cndsr on the t'alavera.1- fault in Hollister. Calif.. has bowed walls, sidewalks, streets, and houses. View 

eastward at Sixth Street. 
54 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FIGURE 2.37. -Valleys and ponds mark trace of the Calaveras fault (first referred to at this location as the Paicines fault). View northwestwstrd 

toward village of Paicines (top). 
2. GEOMORPHIC EXPRESSION 55 

-- - - - - - - . . . - 

FIGCKE US.-The Elsinore fault just north of the United States-Mexican border is not as conspicuously displayed as the San Andrea.? and San 

Jacinto faults, however, this fault markedly influences the topopaphy. Here, a major vallcy is eroded in the general fault zone, but trss 

of late Quaternary displacement are not as obvious as farther northwest along fault. U.S. Geological Survey photograph, taken October IS, 

19%. 
56 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

Fmrri~S'.:W. -Imperial fault south of the Salton Sea, southern California. During the Imperial Valley earthquake of October 15, 1979, surface 

rupture occurred along fault trace (black-and-white line). Bull's-eyc denotes epicenter. From L.S. Geological Survey (1982, cover). K.S. 

Air Force photograph, taken July 17, 1968. 

I 

THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

PI~~~~2.41.-Right-lateral slip accompanyinganearthquake creates characteristic echelon fractures and zones of compression that produce 

a "mole track." This mole track formed along the Coyote Creek fault during- the Borrego Mountain earthquake of 1968. 

PRESENT rn EXPLANATION 

approximately 

located. Arrows 

indicate direction 

of movement 

pate 

A - 

7 

FaunÃ‘Arrow indicate 

Lnr&ion ol reh 

movement 

PRESEN I ^ 

40 m.y. 

2400 krn 

EXPLANATION w x Â¥ass. 

app~-ly 

located. Arrows 

ind'kale dimtion 

of movement 

A 

I - 

7 

FaultÃ‘Arrow indicate 

armon of relative 

movement 
T he San Andreas fault is a trimsfom fault along the boundary 

between the Pacific and North American vlates. Bedrock alorw 

the fault includes va-n0us lithologic units that range in age from 

Precambrian to Tertiary and younger. Some bedrock units that can 

be matched across the fault suggest strike-slip displacement of as 

much as 560 km. 

3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 

CONTENTS 

INTRODUCTION 

The great scar across the land of California, extending 

from the Gulf of California to Point Arena on the north 

coast, was recognized as a major fault during early 

geologic study of the San Francisco peninsula; it was 

4 FIGURE 8.1.-SequentiaJ diagrams showim; interactions between the 

North American. Faratlon. and Pacific plates. a~auming a constant 

relative motion of 6 cm/yr parallel to the San Andrea8 fault (modified 

from Atwater, 1970). position of the North American plate in each 

time frame is relative to those of the Farallon and Pacific plates rather 

than to outlines of diagram. Lengthening interface between the North 

named for San Andreas Valley, which lies a few kilometers south of San Francisco. Interest in the San Andreas 

fault was heightened as a result of movement on the fault, 

in some places as much as 5 m of strike-slip displacement, 

that occurred during the great San Francisco earthquake 

of 1906. Because of this earthquake, a concerted study 

was carried out by several leading geologists of that time 

(Lawson, 1908) that dramatically increased our knowledge of the regional extent and general features of the 

fault. Much controversy ensued during the following 

decades as to the time of origin of the fault, the 

magnitude of cumulative displacement along it, and, 

indeed, even whether the displacement might be principally dip slip rather than strike slip. An early proponent 

of substantial lateral movement on the San Andreas fault 

was Levi Noble (1927) of the U.S. Geological Survey, 

who suggested a 38-km right-lateral displacement based 

on the similarity of Tertiary strata on opposite sides of 

the San Andreas fault at Cajon Pass and Rock Creek (lat 

34%' N., long llT50' W.). Substantial lateral offset of 

Quaternary terraces along the Mojave segment of the 

fault was recognized by R.E. Wallace (1949) while 

working on his Pb.D. thesis at the California Institute of 

Technology. He estimated a slip rate of 0.4 dyr, which 

he extrapolated to 120 km of right-lateral slip since 

mid-Tertiary time (approx 30 Ma). In 1953, a benchmark 

American and Pacific plates, shown in three upper diagram, repsents the& Andrea Worm fault. Captions for eachatep indicate 

amount of time and lateral movement necessary fw the North 

American plate to reach its present position relative to the Pacific 

plate. GS, Guaymas; MZ, Mazatlan; S, Seattle; SF, San Francisco. 

THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA; U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
62 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

paper by M.L. Hill and T.W. Dibblee, Jr., of the calculations of plate motions between North America, 

Richfield Oil Corp. cited various evidence for great Africa, India, Antarctica, and the Pacific, and many 

right-lateral offset along the San Andreas fault and land-based geologists of the time would have agreed that 

speculated that the total offset amounts to 560 km or it seemed "outrageous" to be "studying the San Andreas 

fault bv usine data that is no closer to California than 

FIGURE 3.t-Sketch map of part of 

the Pacific coast of North America. 

showing J. Tuzo Wileon's edy idea 

of the Sac Andreas fault (SAF) as a 

right-lateral mows) transform 

fault connecting the termination of 

two spreading centers represented 

by the East Pacific Rise and the Juan 

de Fuca Rid@ (inixlified Irnm WZÂ¥on 1965). 

more since Jurassic time. 

During middle and late 1960's, a time of great ferment 

of concepts regarding the plate-tectonic development of 

the planet Earth, the foundation was laid for much of the 

present view of the tectonics of California and the San 

Andreas fault. A highly significant breakthrough to our 

understanding of the development of the fault system 

was the brilliantly simple construction by J. Tuzo Wilson 

(1965, fig. 9), who showed the San Andreas fault as a 

transform fault connecting two spreading oceanic ridges 

(figs. 3.1.3.2). This view was soon modified by McKenzie 

7,000 km" (~iwater and Molnar, 1973). 

and Morgan (1969) and Atwater (1970) to account for the 

effects of migrating triple junctions and for the timing, 

rates, and vectors of plate movement. Their platetectonic analysis of the San Andreas fault was based on 

GEOLOGIC FORMATIONS 

, 

On its path through nearly the length of California, the 

San Andreas fault separates major crustal blocks (fig. 

3.3). In much of northern and central California, the fault 

is a southeast-trending boundary between the Salinim 

block of granitic and metamorphic rocks on the west and 

the Franciscan assemblage and overlying strata of the 

Great Valley sequence on the east. In its southerly course 

the fault abruptly curves eastward to cut diagonally 

across the Transverse Ranges, and then splays into 

several auxiliary faults before the main strand terminates near the Gulf of California. In southern California 1 the basement rocks cut by the San Andreas fault are 

mostly Precambrian and younger metamorphic and plutonic rocks, and the crustal blocks on either side of the 

fault generally do not show the distinctive lithologic 

contrast that is so striking in central and northern 

California. 

^ 

FIGURE 3.3.-Geologic sketch map of California, showing distribution 

of principal basement rocks. Cenozoic cover not shown except for the 

Mdoc Plateau, northeastern Sierra Nevada, Great Valley, and Santa 

Barbara-Ventura Basin. Based on U.S. Geological Survey (1966), 

Jennings and others (1977). and Ross (1984). Units: 1, Quaternary 

alluvium-shown only in Great Valley; 2, basement rocks concealed 

by thick Upper Cretaceous and Tertiary deposits in the Santa 

Barbara, Vent-. and Los Angeles Basins- 3. Cenozoic volcanic 

rocks of the Madoc Plateau; 4, Great Valley sequence-Lower 

Jurassic to Upper Cretaceousstrata, including Coast Range ophiolite 

at base; 5, Franciscan assemblage of Lower Jurassic to Tertiary 

oceanic rocks; 6, Sierra Nevada batholith-dominantly Cretaceous 

granitic rocks, 7. Sierra Nevada metamorphic belt-early Paleozoic 

to Late Jurassic rocks, including fracments of ophiolite~, island arcs, 

and melanges, intruded by Mesozoic plutons: 8, Klamath Momtains-early Paleozoic to Late Jurassic ophiolites, island axes, and 

melanges, intruded by early Paleozoic to Cretaceous pinions; 9, 

Salinian block-dominantly Cretaceous plutons intruding metarnorphic rocks of questionable age (Bairett Ridge slice [BRS], commonly 

shown as part of the Salinian block, is here shown as part of unit 11); 

10. southern California batholith-dominantly Cretaceous plutons 

intruding sedimentary and volcanic rocks of Jurassic age and metamorphic rocks of mostly unknown age; 11, mainly Precambrian 

metamorphic and plutonic rocks, in part overlain by Paleozoic 

continental shelf deposits and intruded by Mesozoic plutons, locally 

underthrust in southwestern part of region by schist (similar in 

Ethology to the PelonaSchist) of probable Cretaceous age. 
3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 63 

NORTHERN AND CENTRAL CALIFORNIA 

FRANCISCAN ROCKS 

Franciscan rocks form the east wall of the San Andreas 

fault for virtually its entire course through the Coast 

Ranges of central and northern California, although the 

Franciscan is concealed along some reaches of the fault by 

overlying rocks. The Franciscan is a heterogeneous 

assemblage that consists largely of dismembered sequences of pywacke, shale, and lesser amounts of mafic 

volcanic rocks, thin-bedded chert, and rare limestone. 

EXPLANATION 

- Contact-Dashed where 

approximately located 

- Fault-Dashed where 

approximately located; 

dotted where concealed 

0 

u 

100 200 KILOMETERS 
64 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

These rocks also occur with serpentinite and tectonic 

pods of blueschist in melange zones that are the locus of 

much shearing withim the Franciscan and that generally 

separate blocks of the more coherent sequences. The 

sedimentary and volcanic Franciscan rocks were formed 

in a marine environment, as attested by the abundance of 

foraminifera in the limestone and by radiolarians in the 

chert. Most of these rocks are probably Late Jurassic and 

Cretaceous in age (Bailey and others, 1964), but some of 

the chert and associated volcanic rocks are as old as Early 

Jurassic (Pliensbachii) (Irwin and others, 1977; Blome 

and Irwin, 1983). In the northern Coast Ranges, some of 

the rocks assigned to the coastal belt of the Franciscan 

assemblage are as young as late Tertiary and are thought 

to have accreted to North America during poet-middle 

Miocene time (McLaughlin and others, 1982). The age 

and origin of Franciscan melange is problematic. MidCretaceous limestone in melange near Laytonville in the 

northern Coast Ranges, 225 km northwest of San Francisco, has a paleomagnetic inclination that indicates an 

origin several thousand kilometers to the south (Alvarez 

and others, 1980). Similarly, Franciscan pillow basalt 

about 45 km northwest of San Francisco is thought to 

have moved northward 19Â of latitude (approx 2.000 km) 

from its site of origin (Gromm6, 1984). These and other 

features indicate that some, possibly much, of the Franciscan has been transported great distances northward 

along the Pacific margin relative to a stable North 

America. 

The Franciscan rocks are locally overlain structurally 

by the Coast Range ophiolite and the Great Valley 

sequence, and are separated from them by the Coast 

Range thrust (Bailey and others, 1970). The original 

extent of the Coast Range thrust is not clearly known 

because most of the ophiolite and Great Valley sequence 

that formed the upper plate of the thrust has been 

removed from the top of the Franciscan except in the 

general area of the Diablo antiform, which is marked by 

a line of windows from Mount Diablo to Parkfield, and 

along the west edge of the Great Valley (fie. 3.3,3.4). A 

few small outliers of upper-plate rocks are present 

elsewhere east of the San Andreas fault as far north as 

Pillsbury Lake, 35 km east of Wilits, and the Camp 

Meeker area, 17 km northeast of Bodega Head, and at 

several localities west of the Salinian block as far south as 

the Santa Ynez fault in the Transverse Ranges (see maps 

at front of book). Much of the serpentinite in Franciscan 

melange may well be sheared-in fragments of dismembered Coast Range ophiolite. 

COAST RANGE OPHIOLITE 

The Coast Range ophiolite represents oceanic crust on 

which much of the sedimentary rock of the Great Valley 

sequence was deposited. A complete ophiolite sequence 

consists of serpentiniied harzburgite tectonite at the 

base, overlain by cumulate ultramafic and gabbroic 

rocks, passing upward into noncumulate gabbroic and 

related plutonic rocks, then into diabase dikes, and finally 

into pillow lavas. The Coast Range ophiolite, however, 

generally is highly sheared, dismembered, thinned, and 

locally missing, presumably as a result of faulting, at 

many places along the fault contact between Franciscan 

and Great Valley rocks. Only in a few places is a nearly 

complete lithologic sequence of Coast Range ophiolite 

preserved, and there the total stratigraphic thickness of 

the ovhiolite is about 3 to 5 km (Honson and others. 1981). 

 so ti pic ages ranging from about 166 to 153 Ma (HO~SOII 

and others, 1981) indicate that the Coast Range ophiolite 

is Middle and Late Jurassic in age. Paleontologic and 

paleomagnetic evidence suggests that the Coast Range 

ophiolite formed in an equatorial setting and was transported great distances northward before being accreted 

to North America and overlain by the Great Valley 

sequence (Pessagno and others, 1984; Hopson and others, 

1986; McLaughlin and others, 1988). 

GREAT VALLEY SEQUENCE 

The Great Valley sequence consists of interbedded 

marine mudstone, sandstone, and conglomerate that 

range from Late Jurassic to Cretaceous in age (Bailey 

and others, 1964). It crops out as thick, monotonously 

bedded sections of strata that generally are markedly 

less deformed and more coherent than sedimentary 

sections of the Franciscan and also have greater lateral 

continuity. Where most fully developed, such as along 

the west side of the northern Great Valley, the aggregate 

stratigraphic thickness of Great Valley sequence is at 

least 12 km. The strata normally lie depositionally on 

Coast Range ophiolite except where disrupted by faults, 

but at the north end and alone the east side of the Great ~ ~ 

Valley they onlap the ~evidan and older basement 

terranes of the Klamath Mountaii and Sierra Nevada. 

This enormous thickness of clastic detrital material 

probably represents submarine fans and turbidity deposits that formed as a result of rapid erosion of the ancestral 

Klamath Mountains and Sierra Nevada. 

Ãˆ 

FIGTOE 3.4.-Schematic block diagram of part of the Coast Ranges of 

California, showing gross structural relations between principal 

tithotectonicunits cut by the San Andreas fault (modifiedfrom Irwin, 

1977). The Coast Range ophiolite lies at base of the Great Valley 

sequence and is separated from underlying Franciscan rocks by the 

Coast Range thrust. Blueechist-facie6 metasedimentary and metavolcanic rockB that form border zone in the Franciscan immediately 

below the thrust are not shown. View southeastward along the San 

Andreas fault 
3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 6Ei 

COAST RANGE THRL'ST 

The upper plate of the Coast Range thrust, consisting 

of Great Valley sequence with Coast Range ophiolite at 

the base, is cut by the San Andreas fault only along the 

west side of the Diablo Range in central California. 

There, the upper plate of the Coast Range thrust forms 

a broad blanket over the Franciscan rocks except in 
66 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

several places along the length of the Diablo Range 

where windows, or piercement structures, in the upper 

plate expose Franciscan rocks of the lower plate (fig. 3.4): 

at Mount Diablo. Pacheco Pass. New Idria, and Parkfield. This deformed antiformal structure (Bailey and 

others, 1964) is truncated on the west by the Hayward 

and Calaveras faults in its northern part and by the San 

Andreas fault in its southern part, and narrows to the 

southeast as the San Andreas fault converges with the 

south end of the Great Valley. The ophiolitic rocks are 

generally thinned and highly discontinuous along the 

faults that bound the windows except for the New Idria 

window, which is occupied mostly by serpentinite. South 

of the Pacheco Pass window, the structure is complicated 

by a series of west-northwest-trending synforms that 

cross the axis of the Diablo Range antiform at a low 

angle, and by multiple strands of the San Andreas fault. 

In the Parkfield area, the western part of the antiform 

has been virtually destroyed by its proximity to the San 

Andreas fault. The Franciscan rocks of the Parkfield 

window, and the associated serpentinite and Great Valley sequence of the highly dissected upper plate, now are 

mostly elongate fault slices that form part of the San 

Andreas fault zone (see Dibblee, 1980). 

Most of the serpentinite of the Coast Ranges of 

California is related to the ultramafic parts of the Coast 

Range ophiolite. Serpentinite is strikingly absent along 

the San Andreas fault north of San Francisco, a situation 

that reflects the absence of the upper plate of the Coast 

Range thrust in that area. Although many elongate 

fault-bounded bodies of serpentmized ultrama6c rocks in 

the Coast Ranges traditionally have been mapped and 

described as part of the Franciscan, most are probably 

dismembered and dislocated parts of the Coast Range 

ophiolite or equivalent. Recognition that the ultramafic 

rocks in most places belong to the Coast Range ophiolite 

rather than to the Franciscan assemblage is important to 

a tectonic analysis of the region. The common presence 

and possible seismotectonic significance of serpentinite 

along creeping segments of the San Andreas and related 

faults were described by Alien (1968) and Irwin and 

Barnes (1975). 

SALINIAN BLOCK 

The west wall of the San Andreas fault consists mainly 

of rocks of the Saliinian block from the Transverse Ranges 

northward to Bodega Head. At the latitude of the San 

Francisco peninsula, the Salinian block is separated from 

the San Andreas fault by a narrow fault dice of Franciscan rocks (fig. 3.3). Northward, from just beyond 

Bodega Head to Point Arena, the rocks that crop out 

along the fault are chiefly Upper Cretaceous and Tertiary 

sedimentary strata, but because of the presence of a 

relatively small exposure of spilitic volcanic rocks that 

may be Franciscan, it is questionable whether these 

sedimentary rocks overlie granitic and metamorphic 

rocks of the Salinian block or whether they are separated 

from the Salinian block by another fault just off shore 

(Wentworth, 1968). The north end of the Salinian block 

off shore is thought to be at about the latitude of Point 

Arena (see McCulloch, 1987). At Point Delgada (see maps 

at front of book), which many workers consider to be the 

location of the northernmost onland trace of the San 

Andreas fault, the rocks on both sides of the fault are 

Franciscan. 

The Salinian block is composite, consisting of central, 

western, and northern belts, and commonly is considered 

to include the Barrett Ridge slice (fig. 3.3; Ross, 1984). 

The basement rocks of the Barrett Ridge slice, though 

poorly exposed, are thought likely to be a northward 

continuation of the rocks of the San Gabriel Mountain 

area because they include similar-appearing metamorphic and granitic rocks, as well as schist similar to the 

Pelona Schist, along a possible exposure of the Vincent(?) 

thrust (Ross, 1984). The principal formations of these 

central, western, and northern belts are granitic and 

metamorphic rocks, locally overlain by Upper Cretaceous 

and younger strata. The metamorphic rocks, which 

commonly are moderate- to high-grade gneiss, granofels, 

impure quartzite, and minor schist and marble, probably 

represent a metamorphosed thin-bedded sequence of 

siltstone and sandstone, with lesser amounts of shaly, 

marly, and calcareous strata (Ross, 1978). The metamorphic rocks of the western belt are higher in metamorphic 

grade than those of the central and northern helts. The 

stratigraphic age of the protoliths of the metamorphic 

rocks is not known. The plutonic rocks are mostly granite 

and tonalite, but they range in composition to gabbro. 

U-Pb-isotopic measurements on zircon in the plutonic 

rocks indicate that plutonic activity began about 120-105 

Ma in the northwestern part of the Salinian block and 

migrated southeastward over a period of 40 my., with 

the youngest plutons intruding the Barrett Ridge slice 

about 80-75 Ma (Mattinson and James, 1985). 

The basement rocks of much of the Salinian block do 

not clearly differ from those of the Sierra Nevada, and so 

many workers have speculated that the Salinian block 

may be a displaced part of the Sierra Nevada (for 

example King, 1959; Page, 1981). According to Ross 

(1989, who compared the two terranes in considerable 

detail, the similarities are so great that strong data would 

be required to support an alternative origin. Paleomagnetic data, however, indicate that the Saliian block may 

have been displaced 2,500 km northward since Cretaceous time (Champion and others, 1984) and that it may 

have originated near the latitude of Central America or 

Mexico in the axial part of the Cordilleran Cretaceous 

plutonic arc (Page, 1982). 
3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 67 

DISPLACEMENT OF PRE-QUATERNARY ROCKS BY 

THE SAN ANDREAS FAULT 

The largest offset on the San Andreas fault postulated 

by Hill and Dibblee (1953) is 560 km, on the basis of the 

speculation that a contact between Sierran basement and 

Franciscan rocks near the south end of the Great Valley 

was cut bv the fault and that the contact on the west side 

of the fait was offset northward to a point at sea north 

of Point Arena. This concept, however, is questionable 

not only because of the uncertain correlation of the 

Salinian block with rocks of the Sierra Nevada but also 

because the northernmost exposure of granitic rocks 

along the west side of the fault is at Bodega Head, about 

90 km southeast of Point Arena, and neither the presence 

nor the type of a contact between Sierran-type basement 

and Franciscan rocks is known off shore north of Point 

Arena (see McCuUoch, 1987). Nor is the significance 

dearly known of the possibly substantial lateral offset 

along the San Gregorio fault, which intersects the San 

Andreas at a low angle just west of the entrance to the 

San Francisco Bay. However, Upper Cretaceous strata 

that crop out near Gualala, on the west side of the San 

Andreas fault between Bodega Head and Point Arena, 

were thought by Ross and others (1973) also to have been 

offset 560 km, consistent with Hill and Dibhlee's concept. 

The strata near Gualala include quartz-plagioclase arkose 

and a conglomerate characterized by unusual quartzbearing mafic clasts, including volcanic rocks, diahase, 

and diorite to gabbro. Paleocurrent features indicate a 

source to the east of the fault. The source of the unusual 

gahhroic clasts is thought to he the Eagle Rest Peak area 

in the San Emigdio Mountains, just east of the San 

Andreas fault in the Transverse Ranges (fig. 3.5). This 

area may also be the source of relatively small fault 

slivers of similar mafic rocks that now lie 160 km 

northwest along the San Andreas fault at Gold HiU and 

320 km northwest at Logan (Ross and others, 1973). 

Other features that suggested substantial right-lateral 

offset to Hill and Dibblee (1953) include (1) the south limit 

of Cretaceous strata near Fort Ross, offset 512 kin from 

FIGURE 3.5.-Locations of some pre-Quatemary features offset by the 

San Andreas fault in the California Coast Ranges. The La Honda 

Basin is an offset segment of the San Joaquin sedimentary baain 

(Stanley, 1987); boundary dashed where approximately located. 

Circled x's. bedrcckescosma of unusual auartz-bearinc' mafic rocks 

at Lw, &ld Hill, id Eagle Rest peak (ROBS and others, 1973); 

squares, locations of the Pinnacles and Neenach Volcanic Formations ^r 

(see fig. 3.6) (Matthewt, 1976). Numbered dots: 1, Butano Sandstone 

and Point of Rocks Sandstone Member (of Kreyenhagen Formation), 

representing offset parts of an Eocene deep-sea fan (Clarke and 

Nilaen. 19731: 2. ateenlv southwest-din~inc . . .. alones , of the San Joaauin .$ 

and L; ~ond'a Basins during late &monk time (Stanley, 1987): 3. 

deepest pans oft he San J-in and La Honda Basins adjacent 10 the 

San Andreas fault dunne Saucesian time (Stanley. 1WJ; 4, unusual 

clasts in upper Miocene conglomerate in the Temblor Range east of 

fault and their postulated source area in the Gabilan Range west of 

fault (Huffman, 1972). 

Bodega Head # 

0 100 KILOMETERS 
68 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

their south limit in the Temblor Range; (2) Eocene 

formations in the Santa Cruz Mountains, offset 360 km 

from lithologically and faunally similar formations in the 

Temblor-San Emigdio Mountains; (3) sequences of lower 

Miocene volcanic rocks, red beds, and Oligocene and 

lower Miocene marine deposits of the Gabilan Range, 

offset 280 km from similar sequences in the San Emigdio 

Mountains; (4) a facies transition from marine to continental middle and upper Miocene beds of the Carrizo 

Plain, offset 104 km from a similar transition projected 

from the south end of the Great Valley; and (5) 16 km of 

offset based on the juxtaposition of two facies of Pleistocene gravel south of the Temblor Range. 

Offset of the Eocene formations was described in more 

detail by Clarke and Nilsen (1973), who considered the 

lower to upper Eocene sedimentary sequences to represent parts of a single deep-sea fan that have been offset 

305 km dong the San Andreas fault. These sequences are 

(1) the Twobar Shale Member of the San Lorenzo 

Formation and the Butano Sandstone, exposed in the 

Santa Cruz Mountains on the west side of the fault; and 

(2) the shale member and Point of Rocks Sandstone 

Member of the Kreyenhagen Formation, exposed along 

the southwest boundary of the Great Valley on the east 

side of the fault (fig. 3.5). Both sandstone units, which 

consist chiefly of detritus from the Salinian block, are 

thought to have been deposited as a submarine fan on the 

west side of a deep, northwest-trending offshore basin. 

The Eocene fan deposits on the east side of the fault are 

in a structurally complex area at the south end of the 

Diablo antiform. If they overlie Franciscan basement 

there and Salinian basement on the west side of the fault, 

they indicate that the Franciscan and Salinian basements 

had already been juxtaposed by Eocene time when the 

fan was deposited, as noted by Page and Engebretson 

(1984). 

The strongest evidence for measuring the large 

amount of right-lateral displacement along the trace of 

the San Andreas fault through the Coast Ranges may be 

the correlation of the Pinnacles and Neenach Volcanic 

Formations (Matthew, 1976), which are exposed on 

opposite sides of the fault about 315 km apart (fig. 3.5). 

The Pinnacles Volcanic Formation is in the central Coast 

Ranges, on the west side of the Chalone fault, a parallel 

strand that is several kilometers west of the San Andreas. The Neenach Volcanic Formation is in the Mojave 

Desert, adjacent to the San Andreas fault on the northeast side, about 20 km southeast of its intersection with 

the Garlock fault. The volcanic rocks at both exposures 

rest on granitic basement rocks that are petrographically 

and chemically similar (Ross, 1984). As described by 

Matthew (1976), the volcanic rocks form stratigraphic 

sections that are remarkably similar in composition, 

lithologic sequence, and age; they consist of calc-alkaline 

andesite, dacite, and rhyolite flows interbedded with 

pyroclastic and volcaniclastic rocks (fig. 3.6). K-Arisotopic analyses of the volcanic rocks indicate an early 

Miocene age (23.5 Ma; Turner, 1970). The similarity of 

the 315-km offset of the volcanic formations and the 

305-km offset ascribed to the previously mentioned 

Eocene fan deposits suggests that little movement occurred along that section of the San Andreas fault during 

intervening Oligocene time (Clarke and Nilsen, 1973). 

Paleobathymetric contour maps of middle Tertiary 

(late Oligocene to early Miocene) topographic features of 

the San Joaquin sedimentary basin match similar features across the fault in the La Honda Basin (fig. 3.5), 

according to Stanley (1987). The paleobathymetry is 

based on studies of the distribution of fossil foraminifers. 

In both basins, (1) the southwestern margins were 

bounded by shelf areas and steep north-facing slopes, (2) 

maximum water depth was about 2,000 m, (3) the deepest 

part was truncated by the San Andreas fault, and (4) the 

paleobathymetric contours generally trend at large angles to the fault and are truncated by it. These features 

indicate a post-late Zemoman (late Oligocene to early 

Miocene) displacement of 325 to 330 km and a postSaucesian (early Miocene) displacement of 320 to 325 km. 

Large offset along the San Andreas fault is also 

indicated by an anomalous distribution of upper Miocene 

sedimentary denosits that occur on both sides of the fault 

for more than300 km in the central Coast Ranges. 

According to Huffman (1972), the clast composition of 

upper Miocene sedimentary rocks of the Temblor Range 

lying east of the fault in the southern Coast Ranges 

indicates that they were deposited adjacent to the 

Gabilan Range, which lies to the northwest on the 

opposite side of the fault, and that they subsequently 

have been displaced approximately 240 km (fig. 3.5). 

The various aforementioned features that have been 

used to measure lateral offsets along the San Andreas 

fault are all situated near the present-day trace of the 

fault. They do not measure the substantial displacements 

that probably occurred along presently inactive older 

faults of the system, nor do they measure the large 

lateral movement (described below) that is thought to 

have occurred along the interface between continental 

and oceanic crust during the early plate-tectonic development of the fault system. 

RELATION OF GEOLOGIC STRUCTURE TO 

SEISMIC BEHAVIOR 

From the preceding descriptions, the San Andreas 

fault evidently cuts through many different types of 

rocks and regional structures along its traverse from the 

Transverse Ranges to the Mendocino triple junction, and 

the patterns of seismicity differ strikingly along various 
3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 69 

segments of the fault. These differences in seismic 

behavior coincide so closely with certain geologic situations along the San Andreas fault in central California as 

to suggest a causal relation. 

Certain segments of the San Andreas fault in central 

California are characterized by frequent small-magnitude 

earthquakes and aseismic slippage (creep); creep also 

occurs along the San Andreas fault in Coachella Valley, 

the Imperial fault, and the Superstition Hills fault. Other 

segments, said to be "locked," are characterized by 

infrequent earthquakes, some of which have been hiitorically of large magnitude, and by an absence of creep. The 

occurrence of creep has been described by Allen (1968), 

Wallace (1970), Nason and Tocher (1970). and Thatcher 

(see chap. 7). among others. The creep is mainly on faults 

along the west side of the Diiblo antiform. The creeping 

segment of the San Andreas fault extends from Cholame 

to near San Juan Bautista. Other faults known to creep 

include segments of the Calaveras, Hayward, Concord, 

Green Valley, and Sargent faults (fig. 3.7). The locked 

segments of the San Andreas fault in central California 

extend northward from near San Juan Bautista and 

southward from Cholame. Though of no recognized 

significance, the creeping segment of the San Andreas 

fault terminates near the north end of the Barrett Ridge 

slice of the Salinian block. 

The creeping segments in central California occur 

where the faults regionally cut the upper plate of the 

Coast Range thrust (figs. 3.4, 3.7; Irwin and Barnes, 

1975). This position accounts for the common presence of 

serpentinite along the creeping segments of the San 

Andreas fault, because the serpentinifemus Coast Range 

ophiolite in the upper plate of the thrust is at the fault 

interface along these segments. The lower-plate hciscan rocks form a geochemical province characterized 

by carbon dioxide-rich springs. Where the upper plate of 

the thrust is present, it may suck as a hydraulic cap that 

helps to maintain high pore pressure caused by carbon 

dioxide in the underlying Franciscan rocks, and to direct 

fluid flow into the fault (Irwin and Barnes, 1975). The 

importance of pore pressure in relation to creep is its 

ability to reduce the frictional strength of rocks by 

lowering the effective confining pressure, as demonstrated by Byerlee and Brace (1972). The tectonic relations 

between carbon dioxide springs and seismicity were 

described by Irwin and Barnes (1980). 

-- Vim Mli lull 

"."/\ / \/ - '"- I\/- \\.\\ 

, I\/ \I \ erarntic,b=y?' -,\ ^ /\/\ 

PINNACLES VOLCANIC FORMATION 

0.5 KILOMETER 

OO 

f L 

1 KILOMETER 

.-. --- --- --- --- - - . . - . . - . . - - - - - - - - - - . - - . - - . - - - - - - - - - - - -- - - - . - - . - - - - - - - - - - - - - - . - - . - - - - - - - - - Vitric Ik tuft 

NEENACH VOLCANIC FORMATION 

FIGURE 3.6.-Idealized cross sections showinc - strati-hit ~-. and litholocic similaritv of the Neenach Volcanic Formation. on northeast side of the 

San Andreas fault in southernCalifornia. to lower half of northern& of the Pinnacles ~olcanic Formation severalkilometers nca of fault 

in central California (modified from Matthewrt, 1976) (sec fig. 3.5 for kicatiom). qi, ugglimierate; pr, porphyritic rhyolite. 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

SOUTHERN CALIFORNIA 

TRANSVERSE RANGES AND THE SALTON TROUGH 

Where the San Andreas fault bends eastward, near the 

south end of the Great Valley, it forms the boundary 

between the southern "tail" of the Sierra Nevada and 

rocks of the Transverse Ranges (fig. 3.3). Continuing 

southeastward along the Transverse Ranges, the fault 

trends along the north slope of the San Gabriel Mountains 

and the edge of the Mojave Desert for about 100 km, and 

then cuts diagonally between the San Gabriel and San 

Bernardino Mountains, where a major strand (San Jacinto fault) splits off to the south. The rocks traversed by the 

San Andreas fault in the Transverse Ranges are mainly 

Mesozoic plutons and their Precambrian metamorphic 

and plutonic host rocks. 

South of the Transverse Ranges, the bedrock along the 

fault is concealed by Quaternary deposits of the Salton 

Trough for a length of more than 200 km. The Salton 

Trough, which includes the Coachella and Imperial Valleys, widens toward the southeast, and the number of 

FIGURE 8.7.-Known creeping segments of the San Andreas and 

related faults in central and northern California. Data from N- 

(19731, Frizzell and Brown (1976), Presaitt and Burford (1976), 

Burford and Slum) (19821, and Harsh and Burford (1982). Fa& C, 

Calaveras; CD, Concord; GV, Green Valley; H, Haywari; RC, 

Rodgem Creek; S, Sargent. Compare creeping segments with 

concentratiom of small earthquakes shown in figures 6.4 and 5.5. 

faults and complexity of the zone increase. The east wall 

of the trough is at the San Andreas fault, where it 

consists of Precambrian rocks and Mesozoic plutons and 

of schists like those exposed in the Orocopia and Chocolate Mountains (fig. 3.8). The west wall consists of 

Cretaceous plutonic rocks of the Southern California 

batholith and their metamorphic host rocks, similar to 

rocks exposed in the nearby San Jacinto Mountains. As 

shown by Fuis and others (1983, the trough is a gap in 

the crystalline basement that is filled with Quaternary 

and older Cenozoic sedimentary rocks. The gap increases 

irregularly in width from 20 km at the north end of the 

Salton Sea to 60 km at the United States-Mexican border. 

The enormous thickness of the sedimentary fill is indicated by a drill hole that bottoms in Pleistocene(?) 

sedimentary rocks at a depth of about 4 km (Muffler and 

Doe, 1968). Seismic-refraction studies show an interface 

with "basement" rocks at a depth of 5 to 6 km, and the 

"basement" rocks below the 5- to 6-kmdeep interface are 

thought actually to be metamorphosed Cenozoic fill (Fuis 

and others, 1982). Near the south end of the Salton Sea, 

the San Andreas appears to terminate as a transform 

fault at a spreading center, or pullapart zone. that is the 

most northerly ini series of spreading centers distributed along the length of the Gulf of California which form 

part of the East Pacific Rift. The proximity of this 

pullapart zone accounts for the abundant young volcanic 

and geothermal features in the area (Elders and others, 

1972). 

The Precambrian rocks and associated Mesozoic plutons that constitute much of the crystalline basement cut 

by the San Andreas in southern California are locally 

seen to lie in thrust-fault contact on relatively younger 

metamorphic rocks. In the San Gabriel Mountains, on the 

southwest side of the San Andreas fault, the principal 

country rocks are divided into two plates by the Vincent 

thrust (lat 34-19' N., long 117-45' W.). As described by 

Ehlig (1981). the upper plate of this thrust is a Precambrian gneiss-amphiholite-granite complex (U-Ph age, 

approx 1,700 Ma; Silver, 1966) intruded by a Precambrian anorthosite-syenite-gabbro complex (U-Pb age, 1,220 

Ma; Silver, 1971), all of which are intruded by the Late 

Triassic Lowe Granodiorite (U-Pb age, 220 Ma; Silver, 

1971), by mid-Mesozoic rhyolitic to basaltic dikes, and, 

finally, by granitic plutons of probable Late Cretaceous 

age (U-Pb age, 80 Ma; Carter and Silver, 1971). The 

oldest rocks in the upper -. - plate are thought to be a 

remnant of a Precambrian craton. ~ortheastof the fault, 

in the Mojave Desert and San Bernardino Mountains, 

some Precambrian rocks are unconfonnably overlain by 

lower Paleozoic miogeoclinal strata that are thought to 

represent an essentially autochthonous part of the North 

American craton (Burchfiel and Davis, 1981; Ehlig, 
3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 71 

The lower plate of the Vincent thrust consists of the 

Pelom Schist, which is largely a sedimentary section of 

arkosic sandstone, siltstone, and shale that has been 

metamorphosed to white mica-quartz-albite schist and 

locally includes metavolcanic rocks, metachert, marble, 

and serpentinite (Ehlig, 1981). The section in the San 

Gabriel Mountains has an exposed thickness of 3.5 km. 

The metamorphism decreases downward, away from the 

fault, and so sedimentary structures, including graded 

bedding, are well preserved in the lowest 1 km of section 

This "upside down" metamorphism and other features 

indicate that the metamorphism of the lower-plate rocks 

occurred during thrusting. Pelona-type schist also occurs 

beneath Precambrian rocks of the Barrett Ridge slice 

west of the Tejon Pass and crops out in places along the 

Garlock fault (fie. 3.9). On the northeast side of the San 

Andreas fault in the Orocopia-Chocolate Mountains area, 

just east of the Salton Sea, the schist, there known as the 

Orocopia Schist, lies below the Orocopia thrust. Scattered exposures of Pelom-type schist and the VicentOrocopia thrust continue into southwesternmost Arizona 

(Haxel and Dillon, 1978). 

The protoliths of the Pelom-type schist are thought to 

be deep-marine sedimentary rocks deposited on oceanic 

crust, possibly representing mostly the distal parts of 

turhidite fans. They probably were Jurassic or Cretaceous in age and are thought to have been metamorphosed by the thrusting that probably occurred no later 

than Late Cretaceous time (Haxel and others, 1985). The 

tectonics of the Pelona and Qrocopia Schists is controversial. The Vincent-Orocopia thrust may have dipped 

southwest, and the Precambrian and other rocks of the 

upper plate been thrust northeastward over backarcbasin protoliths of the Pelona and Orocopia Schists 

(Haxel and Dillon, 1978). Conversely, the thrust relation 

may represent a gently north-northeast dipping subducalong the west edge of the Klamath Mountains and the 

west side of the Great Valley, where they form the lower 

plate of the Coat Range thrust (Blake and others, 1967). 

. . 

tion zone in which the protolith of the&& were 1 

trench deposits similar to the Franciscan rocks and were 

thrust north-northeastward under the sialic North American olate (Burchfiel and Davis. 1981: Crowell. 1981). The 1 

~eloL-type schist is reminiscent of the Franciacanderived South Fork Mountain Schist and related semischists of northern California that form a narrow, 

virtually continuous selvage for hundreds of kilometers 

FIGURE 3.8.-Salton Trough and north end of the Gulf of California, 

showing major spreadinzcenters and termination of the San Anh 

fault (modified from Fuis and others, 1982). Spreading centers: BZ, 

Brawley seismic zone; CP, Cem Prieto dermal area; W, Wagner 

Basin. Migor transform faults: CPF, Cem Weto; IF, Imperial; 

SAT, San Andreas. Other major faults: E, Elsinom; EH, East 

Highland Canal seismicity lineament: LS, Lamina Salada; SJ, San 

Jacinto. 

EXPLANATION 

Spreading centerÃ‘Dashe where approximately 

located. Arrows show dimion of spread 

Inferred extent of Precambrian and related basement 

rocks 

IÃ‘Ã Inferred extent of Southern California batholith and 

related basement rocks 

Submarine depression (Wagner basin) 

?--- Approximate contactÃ‘Querie where uncertain 

7- 7 Fault-Dashed where approximately located; queried 

where uncertain. Arrows indicate direction of 

relative movement 

------ Sdsmidty lineament 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

DISPLACEMENT OF BASEMENT ROCKS BY 

THE SAN ANDREAS FAULT 

The distribution of the Felons-type schist (fig. 3.9) has 

played an important role in many attempts to measure 

the slip on the San Andreas fault in southern California. 

In their classic report, Hill and Dibblee (1953) recognized 

the similarity of the Pelona Schist of the San Gabriel 

Mountains area to the Orocopia Schist east of the Salton 

Sea and postulated that these schists are offset 257 km 

from one another by the San Andreas fault. Crowell 

(1962) noted that the Precambrian rocks and Pelona-type 

schist of the Tejon area are separated from those of the 

San Gabriel Mountains area by the San Gabriel fault and 

that the Tejon, San Gabriel Mountains (Soledad), and 

Orocopia Mountains areas contain not only similar Precambrian rocks and Pelona-type schist but also similar 

Oligocene and other Tertiary strata. The similar lithologies and geologic histories of the rocks of these three 

areas indicated to him that these rocks once formed an 

east-west-trending belt which has been segmented and 

displaced by right-lateral slip of approximately 50 km on 

the San Gabriel fault and 210 km on the San Andreas 

fault. However, the validity of the concept that the 

Tertiary rocks of the San Gabriel (Soledad) and Orocopia 

Mountains areas once were parts of a single basin was 

questioned by Spittler and Arthur (1973), who believed 

that the Tertiary strata of these two areas were deposited in separate basins, consist of distinctly different flow 

rocks, and are dissimilar in age. 

More recent schemes for measuring offset along the 

San Andreas fault in southern California have been 

proposed. Among them, Powell (1981) postulated that all 

major exposures of the Pelons-type schist are in the core 

of an antiionnal fold and that they once formed a nearly 

continuous structure which subsequently has been disrupted by Cenozoic strike-slip faults (fig. 3.10). An 

important part of Powell's palinspasti reconstruction of 

the antiform is an old east-west-trending fault that is a 

composite of the Sam Fmcisquito, Fenner, and Clemens 

Well faults. This fault is thought to have had 80 km of 

right-lateral slip, cutting the antiform, and later was cut 

diagonally by the San Andreas fault. The axis of the 

antiformal structure is a linear feature that can be used 

as a "piercing point" in measuring offset where the axis 

appears on opposite sides of afault. On this basis, the axis 

in the San Gabriel Mountains (Sierra Pelona) area is 

offset 220 km along the San Andreas fault from the axis 

in the Orocopia Mountains area (Powell, 1981). The 

Transverse Ranges segment of the San Andreas fault 

was described by~attiand others (1986) as consisting of 

several old strands (Wilson Creek, Mission Creek, and 

FIGURE 3.9.-DiBtribtition of the Pelona Schist and lithdogically similar schists in southern California and southwestern 

Arizona (modified hrn Haxel and iron, 1978). Northemmad occurrence along the Sm AndreaB fault is in Bamtt Ride 

slice of the Saliman block. 
3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 73 

Mill Creek faults) and a young strand (San Bernardino 

strand). The totaldisplacement across all of these strands 

is thought to be 160Â±1 km (Matti and others, 1983, on 

the basis of exposures of a distinctive Triassic megaphyric monzogranite. The exposures of this monzogranite, 

which are on the northeast side of the San Andreas fault 

in the San Bemardino Mountain. but on the opposite side 

of the fault in Liebre Mountain Cat 34'43' N., long 11820' 

W.), are thought to represent displaced parts of the same 

pluton. 

PLATE-TECTONIC DEVELOPMENT OF 

THE SAN ANDREAS FAULT 

The Pacific coast of North America is a highly mobile 

zone of tectonic interaction between continental crust on 

the east and oceanic crust on the west. At some times and 

places, the relative movements of these contrasting 

crustal domains have been sufficiently convergent that 

the oceanic crust has underthrust the continental crust 

and swept island arcs and other crustal materials into the 

zone of interaction (Hamilton, 1969). However, much of 

the movement along this zone of interaction seems to 

have been oblique or lateral in an overall northwestsoutheast direction, and so as the two types of crust 

moved past one another, fragments of all sizes were 

sliced or pulled from them and carried various distances 

away from their source. By these processes of mnvergence and lateral translation, large crustal fragments 

referred to as "terranes" (Irwin, 1972, Coney and others. 

1980; Schermer and others, 1984) are juxtaposed against 

others that may differ strikingly in lithology, age, genetic 

environment, stratigraphy, metamorphic facies, plutonism, and mineral deposits. 

The general directions and rates of movement of the 

major crustal domains that converged along the western 

margin of North America are amenable to explanation by 

the theory of global plate tectonics as far back in time to 

EXPLANATION 

Ã‘Ã‘t ~~Bfc"" 

? - Fault-Dashed where approximately 

located: queried where uncertain 

FIGUEE 3.10.-Sequential diagrams showing evolution of the San 

Aadreaa fault system in southern California as proposed by Powell 

(1981). A, Approximately 30-16 Ma-formation of regional antifonn 

cored by ~chiet lithologicaflv similar to the Pelona Schist. B, 

AppBBrniately 20-10 Ifa-dkplacBlXillt of antBonn by the Sat 

Fmdtquitii-FennerOemerai Well fault. C. Appxhately 10-2 

Ma-inception of San Andreas-San Gabriel strand ofthe San Andreas 

fault, displaying approximately 60 km of right-lateral displacement. 

D, bximately 5-0 Ma-left-lateral bending of San Gabriel 

segment and possible inception of the Garlock fault. E, Approximately 2-0 Ma-inception of modern southern-California segment of 

the Sail Andreas fault, displaying 220 km of right-lateral dieplacemerit F, Approximately 2-0 Ma-left-lateral bending of Transverse 

Ranges segment of the San Andreas fault and development of the San 

Jacinto fault. Faults; CW. Clemem Welt F. Fenner, SA., San 

Andreas; SF, San Francisquito; SG, San Gabriel. Antifonn aegmentB: 

CM, Chocolate MountalllB; ESG, Eastern San Gabriel Mountnine; 

LSB, Little San Bemardino Mountains; 0, Orocopia Mountains; SP, 

Sierra Pelona; SS, Sierra de SaUnas. 
74 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

at least the Mesozoic (Engebretson and others, 1986). 

According to this theory, the crust of Earth is a mosaic of 

interacting rigid plates. Boundaries between the plates 

are spreading ridges where the plates pull apart, oceanic 

trenches above subduction zones where the plates converge, and transform faults where the plates slide 

laterally past one another. These boundaries are the loci 

of seismic activity. The corollaries of the theory of global 

plate tectonics have been enormously valuable as an aid 

in recognizing the genetic environments of the various 

terranes, whether oceanic or continental crust, volcanic 

island are, or oceanic trench. The presence of multiple 

ophiolite and blueschist belts along parts of the western 

margin of North America (Irwin, 1977) indicates that 

some of these terranes were subduction related during 

Paleozoic and Mesozoic time. Although much has been 

learned of the relative motions between the oceanic and 

continental rocks for the past 180 m.y. (for example, 

Engebretson and others, 1985). the tectonics of the zone 

of allochthonous temes is so complex, and the sites of 

origin of the various terranes generally so obscure, that 

much painstaking research remains to be done before the 

Paleozoic and Mesozoic margins of western North America can be palispastically reassembled. 

The tectonic setting of the continental margin during 

late Mesozoic time, before development of the San 

Andreas fault, is highly controversial. Some geologists 

(for example Dickinson, 1981) favor a model of highly 

convergent plate interaction to develop a continental 

margin of the Andean type (fig. 3.11). Others, however, 

interpret certain paleontologic and paleomagnetic evidence to indicate that some late Mesozoic rocks of the 

Coast Ranges were translated great distances northward 

from equatorial sites of origin during Late Jurassic time, 

before accretion to North America and deposition of the 

Great Valley sequence (Hopson and others, 1986). This 

movement is thought to have been followed in Late 

Franciscan 

subduction 

Volcanic chain 

*s c.2- 

Great Vi 

saquenc. ... ,forearc basin 

Oceanic crust7 

FIGURE 3.11.-Schematic diagram showing California as an Andeantype continental margin during late Mesozoic time (modified from 

Dickmaon, 1981). 

Cretaceous and early Tertiary time by a second episode 

of dextral translation during which part of the Franciscan 

assemblage was accreted and parts of the Coast Range 

ophiolite and overlying Great Valley sequence moved 

northward for distances as great as 1,120 km (McLaughlin and others, 1988). 

The rates and relative directions of motion of the 

principal tectonic plates are based mainly on patterns of 

magnetic anomalies in the oceanic crust. These patterns 

indicate that a subduction-related trench lay offshore of 

western North America during early Tertiary time 

because of the convergence of the Farallon plate (see fig. 

3.1), and that strike-slip movement on faults of the Sari 

Andreas system began no earlier than approximately 30 

Ma (late Oligwene), when the Pacific plate first impinged 

on the North American plate (McKenzie and Morgan, 

1969; Atwater, 1970). Triple junctions formed at the 

point of contact of the Pacific plate with the North 

American plate and migrated to the northwest and 

southeast as subduction of the Farallon plate continued. 

These triple junctions are now approximately 2,500 km 

apart: The Mendocino triple junction is off the coast of 

northern California, and the Rivera triple junction is at 

the mouth of the Gulf of California. The relative motion 

along the transform fault that formed the lengthening 

boundary between the Pacific and North American plates 

was right lateral. This early transform movement probably was not along the modern trace of the San Andreas 

fault hut must have been along other faults of the system 

that now lie mostly to the west and at the edge of the 

continent (fig. 3.12). The modern San Andreas fault 

apparently did not come into being in southern California 

until the opening of the Gulf of California during Pliocene 

time, about 4 Ma, since which time Baja California has 

moved 260 km away from mainland Mexico (Larson and 

others, 1968). The San Andreas fault is commonly refed to as the boundary between the Pacific and North 

American plates, which is true in the sense that the rocks 

on the west side of the fault are moving somewhat in 

concert with the Pacific plate, although those rocks 

actually are displaced fragments that once were part of 

the North American plate (fig. 3.13). 

During the early development of the San Andreas fault 

system, the principal movement must have been along a 

transform fault that formed the boundary between rocks 

of the North American plate and newly formed oceanic 

crust of the Pacific plate as the triple junction migrated 

southward. At some point during southward migration of 

this triple junction, the transform apparently jumped 

eastward one or more times to positions within the North 

American plate, to become the northern section of the 

modern San Andreas fault. The modem trace of the San 

Andreas fault in central California probably had only 

minor slip until about 12.5-10 Ma and probably was not 
3. GEOLOGY AMD PLATE-TECTONIC DEVELOPMENT 

NORTH AMERICAN PLATE , -- 

EXPLANATION 

/\/ Spwingcenter 

- Subduction zone-Dashed where approx- imately located. Sawteetii on upper plate 

A 7 Fault-Dashed whereapproximately located. A- indicatedirectionof relative 

movement 

OR Rivera tripleJunctiM 

presently most &ve element, the present-day Sun Andreas fault, is now in rocks of North American 

nlated. IneariierdiacTanÃˆ.oartialoutlineoftheRalfdCalifornia.whichdidnotexistbefore 
76 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

the strand of dominant slip before 7.5-5 Ma (Dicldnson San Andreas fault was formed by similar eastward jumps 

and Snyder, 1979). The southern section of the modem 1 of the transform fault, resulting in opening of the Gulf of 

Â¥XPLANATIO 

Captured part of North 

American plate 

Remnantsof Farallon plate 

SpreadingcenterÃ‘Arrow 

indicate direction of 

indicate direction of relative 

movement 

MOLOKAI FRACTURE ZONE 

CIAR10N FRACTURE ZONE 
3. GEOLOGY AND PLATE-TECTONIC DEVELOPMENT 

California. Some of the major faults that lie between the 

San Andreas fault and the continental margin may 

represent these earlier positions of the transform, and 

some of these faults are still active. Possible candidates 

for such early intermediate faults include, among others, 

the San Gregorio-Hosgri fault in central and northern 

California and the Elsinore and offshore faults in southern California. With our present state of knowledge, 

however, it is unclear which of the many faults in the 

Coast Ranges are early faults of the San Andreas system 

and which may have formed before mid-Tertiary initiation of the San Andreas system. 

The models of McKenzie and Morgan (1969) and 

Atwater (1970) have been widely used in relating Pacific 

plate interactions to the tectonics and geology of California and other parts of western North America (fig. 3.1). 

A more actualistic developmental sequence of diagrams 

was drawn by Dickinson (1981), in which the position of 

an early San Andreas fault is shown before the opening of 

the Gulf of California (fig. 3.12). In figures 3.1 and 3.12, 

the boundary between the Pacific and North American 

plates is shown as a transform fault that was formed by 

the passage of migrating triple junctions. However, 

marine geologic studies offshore of California between 

Point Conception and Cape Mendocino indicate that this 

old interface between the Pacific and North American 

plates is an inactive eastdipping low-angle fault which is 

interpreted to be a fossil subduction zone (McCulloch, 

1987). There, magnetic stripes of the oceanic (Pacific) 

plate can be traced some distance eastward-locally as 

much as 30 km-beneath the leading edge of the upper 

(North American) plate, and both plates are covered 

along the suture by a thin veneer of undeformed Miocene 

strata. The presence of a Miocene or older subduction 

zone rather than a transform fault at the ocean-continent 

interface is difficult to reconcile with a strict interpretation of McKenzie and Morgan's (1969) and related models, 

although it is interpreted to indicate that the transform 

motion was accompanied by oblique convergence (McCulloch, 1987). 

The present overall rate of relative movement between 

the Pacific and North American plates, earlier thought to 

be about 6 cmlyr (Atwater, 1970) or 5.6 cdyr (Minster 

4 

FIGURE 3.13.-Western margin of North America, showing tectonic 

capture of part of the North American plate by the Pacific plate 

(plate-tectonic base modified from Moore, 1981). Captured part of the 

North American nlate is now movinc- - northweaterlv ~~~. with the Pacific ~ ~~ ~~~~ 

plate and is commonly considered to be part ofthe Pacific plate. West 

boundary of captured pan of the North American plate is inactive 

early tmfnrm fault with component of nuhduction (open aawirt-th; 

see text); interfaces of the North American plate with remnants of 

the Farallon plate are active subduction zones (solid sawteeth). 

and Jordan, 1978), is now thought to be more likely about 

4.8 cdyr (DeMets and others, 1987). This rate probably 

has varied over time (Atwater and Molnar, 1973); however, the rate of relative motion between the two plates 

is substantially greater than the slip rates based on 

measured offsets of geologic features along the San 

Andreas fault. For example, the 315-km offset of the 

lower Miocene Pinnacles and Neenach Volcanic Formations indicates a minimum overall slip rate of 1.3 to 1.4 

cdyr (Matthews, 1976), and the offset of the channel of 

Wallace Creek by the San Andreas fault in central 

California indicates a slip rate of about 3.4 cmlyr for the 

past 3,700 yr and of 3.6 cm/yr for the past 13,250 yr (Sieh 

and Jahns, 1984). Geodetic-survey measurements indicate that slip rates during the past 90 yr in central 

California average 2.9 cdyr for the upper 15 km of crust 

but 3.7 cdyr below 15 km (Thatcher, 1977). The discrep 

ancy between the rate of relative motion between the 

Pacific and North American plates and the much smaller 

slip rate on the San Andreas fault has been noted by 

many workers (for example, Atwater, 1970; Minster and 

Jordan, 1978; Weldon and Humphreys, 1986). Part of the 

total slip along the boundary between the Pacific and 

North American plates probably is occurring in small 

increments along other faults in a broad zone of interaction that may extend from the continental margin eastward even as far as the Basin and Range province. 

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southern California: Tectonics, v. 5, no. 1, p. 33-48. 

Wentworth, C.M., 1968, Upper Cretaceous and Lower Tertiary strata 

near Gualala, California, and inferred large right slip on the San 

Andreas fault, in Diekinson, W.R, and Grantz, Arthur, eds., 

Proceedings of conference on geologic problems of San Andrea 

fault system: Stanford, Calif., Stanford University Pliblicatiom in 

the Geological Sciences, v. 11, p. 130-143. 

Wilson, J.T., 1965, Anew classoffaults and theirbearing oncontinental 

drift: Nature, v. 207, no. 4996. p. 843-847. 


D isplaced or deformed rock units and landfarms record the past 

2 m.y. of faulting, folding, uplift, and subsidence in California. Properly interpreted, such evidence prwides a quantitative 

basis far predicting future earthquake activity and far relating 

many diverse structures and landforms to the 5 cmlyr of horizontal 

motion at the boundary between the North American and Pacific 

plates. 

4. QUATERNARY DEFORMATION 

CONTENTS 

INTRODUCTION 

Modern techniques of geologic dating and expanded 

research on earthquake hazards have greatly improved 

our knowledge of the San Andreas fault system. Much of 

this new knowledge has been gained since 1965, and that 

part which concerns crustal deformation during the past 

2 m.y. is briefly summarized here. 

In emphasizing recent work, I cite only a few of those 

earlier investigators who first recorded the extent and 

timing of deformation in upper Cenozoic deposits. Their 

contributions deserve a more complete accounting, and in 

several works cited herein such credit is given, most 

thoroughly by Wahrhaftig and Binnan (1965). Also 

especially important in understanding the framework of 

the fault system are the 1:250,000-scale, lo by 2' sheets of 

the "Geologic Atlas of California" published by the 

California Division of Mines and Geology, the 1:750,000- 

scale "Geologic Map of California" (Jennings, 1977), and 

the l:750,000-scale "Fault Map of Californian (Jennings, 

1975). 

The Quaternary deformation processes-chiefly faulting, folding, and uplift-represent crustal changes that in 

many places continue today but at barely perceptible 

rates. The perspective of 2 m.y. of geologic time permits 

us to detect and measure these processes, and where 

historical deformation is evident-as shown by seismology, geodetic studies, or geologic investigations of recent 

earthquakes-the Quaternary record provides an independent check on the reliability of our observations and 

measurements. More importantly, we also use the Quaternary viewpoint to sketch the outlines of the currently 

active fault system, its intact crustal elements and major 

active faults, and how these faults propagate and change 

over time. From such evidence we can better understand 

4 FIGURE 4.1.-Strata within a few hundred meters of the San Andreas fault are highly deformed. Here, folded and faulted beds of Pliocene age 

on the north side of the San Andreas fault are exposed in a roadcut along California Highway 14 just south of Pahndale along the southern 

margin of the Mojave Desert. Photograph by R.E. Wallace, U.S. Geological Survey. 83 

THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA: US. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
84 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

the pre-Quaternary history of the system and build the 

predictive kinematic models needed for earthquake hazard assessment. 

Evidence of Quaternary deformation comes chiefly 

from the observed displacement of strata or geomorphic 

features (fig. 4.1). Where this displacement can be dated 

or. at least, bracketed in time. the aee of deformation can 

be established, and its rateÃ‘ffen&ll an average-can 

be determined. In this report, most deformation rates 

are converted to units of centimeters per year, in part to 

facilitate comparison of different data sets but also 

because the dominant process, strike slip on major faults, 

is conveniently expressed in such units. Both the amount 

of displacement and the materials used to determine the 

age of displaced geologic markers require careful geologic 

analysis to ensure that field relations are unequivocal and 

correctly interpreted: the most reliable deformation rates 

depend on highly detailed analyses of local stratigraphy, 

which provide multiple measures of displacement over 

broad timespans, typically thousands to tens of thousands of years. 

Deposits of Quaternary age can be dated by several 

methods, most of which ultimately depend on geochemical analyses that require highly sophisticated laboratory 

practices and expertise, as well as carefully chosen 

samples. The principal methods of dating Pleistocene and 

Holocene deposits within the San Andreas fault system 

(fig. 4.2) include radiocarbon, soil chronostratigraphy, 

correlation with standard sea-level stages (chiefly used 

for uplifted marine terraces), tephrochronology, and 

mammalian and invertebrate paleontology; supplementary techniques include paleornagnetism, uranium-series 

analysis, and amino-acid racemization. The underlying 

theory and limitations of these methods (for example, 

Pierce, 1986) are beyond the scope of this review. Each 

method, however, differs in resolving power, the timespan over which it is effective, and applicability to 

different rock types. 

Difficulties in dating deformed Quaternary deposits 

are not the only deterrent to reconstructing the history of 

faulting. Large areas within the fault system are nearly 

devoid of Quaternary deposits; others are masked by 

landslide deposits, which conceal Quaternary folds and 

faults. Few published geologic maps differentiate bePleistocene, but the view much beyond the past 100 ka is 

still poorly resolved. Despite the problems, the results of 

the past 2 decades of Quaternary research have brought 

tween faults with QuaGemovement and those that 1 

have long been inactive, and for even the best known 

faults our knowledge of Quaternary faulting is still 

incomplete. Quaternary reverse and thrust faults, more 

difficult to identify and map than their strike-slip counterparts, are surely underrepresented on published geologic maps and in the literature. 

Our knowledge of Quaternary deformation is thus 

FIGURE 4.2.-Dating methods used in studies of the Quaternary 

deposits of the San Andreas fault system. Shading shows principal 

methods and timespans over which they have been most effectively 

used; open bars indicate potential range of the methods. Other 

available methods of Quaternary dating, some of which are not 

shown, either are littleused or have produced few or no reported data 

on Quaternary deformation within the fault system. Paleomagnetic 

mlarities: B. Brunhes Normal-Polaritv Chn: J. Jaramillo Normalincomplete and biased. We have learned much, and are bolnrity subchron; M, Matuyama Revolted-polarity chron; o, 

rapidly learning more, about the Holocene and latest Olduvai Normal-Polarity Subchron. 
4. QUATERNARY DEFORMATION 

new insights into the mechanics of the fault system and submarine relief, and sea-floor deposits of young sedithe promise of more discoveries yet to come. ment (Silver, 1971; McCulloch and others, 1985) obscure 

details of its Quaternary history. 

NORTHERN SECTION OF 

THE SAN ANDREAS FAULT SYSTEM 

SAN ANDREAS FAULT 

The northern section of the San Andreas fault system, 

as defined here, extends from Punta Gorda, 40 km 

southeast of the Mendocino triple junction, to the northem Gabilan Range (fig. 4.3). This belt of Quaternaryand earlier-faulting and folding, about 120 km wide by 

500 km long, includes the San Andreas, at least four 

major related faults, and several smaller faults, 5 to 20 

km long. All of these faults displace Quaternary deposits 

or landforms, and all are seismically active. From west to 

east, major faults within the system are the San Gregono, San Andreas, Hayward (and its northern, rightstepping extensions, the Rodgem Creek-Healdsburg and 

Maacama faults), Calaveras (and its northern, rightstepping extensions, the Concord and Green Valley 

faults), and an illdefined, possibly discontinuous fault 

zone along the eastern margin of the Coast Ranges that 

includes, from north to south, the Stony Creek, GreenviUe, and Ortigalita faults. 

The offshore counterpart of the main San Andreas 

fault, which trends seaward and westward from Punta 

Gorda to the Mendocino triple junction, represents the 

latest segment of the northward-growing transform. This 

east-west-trending, offshore segment of the fault is 

located chiefly on seismic and bathmetric evidence; it 

follows along the north face of the Gorda ~scarpment 

(along lat 40%' N. and west of long 124"39' W.) and the 

south wall of Mattole Canyon (from the east end of the 

Gorda Escarpment southeast to lat 40Â¡17 N., long 

WW W.), but water depths of as much as 2,000 m, high 

EXPLANATION 

Alluvial and estuarine deposits (QuaternarybChiefly basin 

fill; may also include some Pliocene deposits 

Clear Lake Vdcanics (Quaternary and TertiarybFlows, tuffs. 

aÃ 

and breccias of dacite and rhvolite and less abundant basalt. Axe i{? 

Bedrock (Tertiary and olderJ-Varied rock types 

- Fault exhibiting evidence of Quaternary displacement-Doned 

where concealed by water 

SAN 

-ANDREAS 

FAULT 

FIGURE 4.3.-Geologic sketch map of the northern Coast Ranges, 

central California, showing faults with Quaternary activity and basin 

deposits in northern section of the San Andreae fault system. Fault 

patterns are generalized, and only major faults are shown. Several 

Quaternary basins are fault bounded and aligned parallel to strikeslip faults, a relation most apparent along the Hayward-Rodgem 

Creek-Maacama fault trend. 
86 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

From Punta Gorda southward to Point Arena the fault 

is almost entirely off shore; its location is known from 

sea-floor features, coastal geomorphology, and relations 

at Point Delgada, where it locally intersects the coastline. A line of coastal bluffs, 350 to 850 m high and sloping 

W4", trends N. 50Â W. for 20 km northwest of Point 

Delgada. At the base of the bluffline, Lajoie and others 

(1982) measured Holocene uplift rates of about 0.5 cmlyr 

from dated molluscan fossils in beach ridges 30 m above 

present sea level. Seaward of the bluffs, the Spanish and 

Delgada Submarine Canyons head near the present 

coast, 2 to 3 km landward of the 100-m bathymetric 

contour, which approximately defines the late Wisconsin 

sea level. Shoreward of this contour, both canyons are 

linear, and neither they nor the intervening sea floor 

exhibits evidence of vertical or horizontal displacement. 

These relations are consistent with a near-shore San 

Andreas fault, located in or near the surf zone and dose 

to the base of the coastal bluffs. The unusually large 

component of vertical slip (0.5 dyr, northeast side up) 

for this segment of the fault may result from its more 

westerly trend than that of segments farther south, 

which exhibit almost pure strike slip; or, as suggested by 

Merritts and Bull (1989), it may result from the late 

Quaternary passage of the Mendocino triple junction 

along this part of the coast. 

The fault intersects the coastline about 6 km northwest 

of Point Delgada, cuts across the point, and locally 

separates upper Pleistocene marine terrace gravel and 

younger alluvium (McLaughlin and others, 1983) along 

the coast from intensely deformed bedrock inland. Other 

locations have been suggested (Curry and Nason, 1967; 

Beutner and others, 1980, McLaughlin and others, 1983) 

for the fault here, but several lines of independent 

evidence support this trend, including fault slip during 

the 1906 earthquake (Lawson and others, 1908, p. 64-68), 

the alignment of scarps, sags, and saddles along the 1906 

fault trend, bedrock faulting colinear with 1906 faulting 

(Brown and Wolfe, 19721, and continuity to the north and 

south with offshore and nearshore evidence of recent 

faulting. 

From Point Delgada southward to Point Arena, a 

distance of 125 km, the fault is defined by aligned, chiefly 

west-facing scarps on the sea floor, the right-lateral 

deflection of Noyo Submarine Canyon, straight boundaries of irregular submarine topography, and a 2- to 

5-km-wide zone of disturbed reflective layers evident in 

continuous seismic-reflection profiles (Curray and Nason, 1967). The shapes and dimensions of sea-floor 

features, as shown by seismic-reflection profiles and 

modem bathymetric maps, resemble fault-formed features on land and provide similar evidence for rightlateral strike slip. The offshore data, however, are 

insufficient to determine the amount of Quaternary 

displacement. 

Onland segments of the fault between Point Arena and 

the Gabii Range exhibit abundant geomorphic and 

geologic evidence of Quaternary displacement, some of 

which can be translated into slip rates. Trunk streams, 

systematically displaced from their headwaters by at 

least 10 km, record cumulative strike-slip faulting that 

must span much of Quaternary time. For example, the 

fault-deflected, northwest-trending courses of the Garcia 

and Gualala Rivers follow the fault zone for 50 km 

between Point Arena and Fort Ross; their present 

alignment (fig. 4.4) documents a long and complex history 

of channel extension and stream piracy. Three streams 

east of the fault-the Garcia River and the Wheaffield 

and South Forks of the Gualala River-exhibit similar 

offsets of about 13.5 km from possible former outlets to 

the sea west of the fault: the Garcia River from its 

present mouth near Point Arena, the Wheaffield Fork 

from the present mouth of the Gualala River near 

Gualala, and the South Fork from a low gap (110-m 

elevation) in the coastal ridge near the Sea Ranch. The 

average rate of post-Pliocene slip across the boundary 

between the North American and Pacific plates, derived 

from spreading rates at the mouth of the Gulf of 

California (DeMets and others, 1987, p. 912). is estimated 

at 5.1 cmtyr. In northern California, this slip is distributed chiefly on the San Andreas and other active faults to 

the east. On the San Andreas fault at Point Arena, 

Prentice (1989) obtained a maximum slip rate of 2.620.3 

cm/yr from a faulted stream channel "C dated at 2,350 to 

2,710yr. Together with the assumed 13.5-kmoffsets, this 

rate implies that the postulated drainage diversions 

required about 519 ka. Other lies of evidence indicate 

much greater antiquity for this and nearby parallel 

strands of the fault. 

Despite abundant evidence of Quaternary faulting 

along the northern section of the San Andreas fault, other 

examples of measurable fault displacement of dated 

upper Pliocene and Quaternary deposits are sparse. 

Much of the fault lies off shore, and most onshore 

segments cut older, highly deformed rocks of the Franciscan complex, from which the upper Cenozoic deposits 

have been eroded. Some reliable indicators of Quaternary 

slip, however, come from geologic and geomorphic relations observed from San Francisco southward. 

On the San Francisco peninsula and farther south in 

the Santa Cruz Mountains, estimates of Quaternary slip 

on the Ban Andreas fault have been derived from offset 

biofacies and lithofacies relations. The fault offsets are in 

discontinuously exposed upper Cenozoic strata, which 
4. QUATERNARY DEFORMATION 87 

FIGURE 4.4.-Sketch map of part of northern California, showing 

coastal drainages south of Point Arena, systematically diverted by 

right-lateral slip on the San Andreas fault (dashed and shaded line). 

Possible offsets of 13.5 km are indicated for the Garria River (A-A'), 

the Wheatfield Forkof the Gualala River (B-B'), and the South Fork 

of the Gualala River (C-C'). 

have been assigned by different investigators to the 

marine Merced, or Mered(?), Formation and the nonmarine Santa Clara Formation. East of the San Andreas 

fault, both the Merced and Santa Clara Formations 

contain, in their upper parts, the 400-ka-old Rockland ash 

bed of Sam-Wojcicki and others (1985). thus establishing at least a partial age equivalence of these two units. 

The upper, Rockland ash-bearing part of the Merced 

Formation (Clifton and others, 1988) is fault bounded on 

the southwest and juxtaposed against a ridge of older 

rock for 11 km southeast of San Francisco (fig. 4.5). The 

minimum offset of 11 km from any western counterpart of 

the Merced outcrop belt and the 0.4- to 2.0-m.y. age of 

the faulted Merced strata require a minimum slip rate of 

1.7Â±1. cmlyr for this section of the San Andreas fault. 

A similar rate can he derived from the present distribution of upper Pleistocene marine deposits, which 

Addicott (1969) interpreted as the faulted remnants of a 

narrow marine embayment that once extended southeastward across the San Andreas. The fault separation of 

35 to 40 km (Addicott, 1969, fig. 2) and an assumed 2- to 

3-m.y. age for the embayment yield a geologic slip rate of 

1.6k0.4 cdyr. 

Additional evidence for Quaternary strike slip (Dibblee, 1966; Cummings, 1968) comes from the southern 

Santa Cruz Mountains south and west of San Jose. 

Distinctive reworked conglomerate clasts in the Code 

Madera facies of the Santa Clara Formation west of the 

San Andreas fault are displaced 28 km from their source, 

an Upper Cretaceous pebble conglomerate lens on the 

east side of the fault. The age of the Corte Madera fades 

is poorly constrained, but Cummings (1968) assumed a 

probable age range of 1 to 3 m.y., which yields slip rates 

of about 1 to 3 cmlyr. 

These slip rates carry uncertainties as great as Â± 

cmlyr, but because they integrate slip over millions of 

years of geologic time, short-term changes in fault 

activity influence them only slightly. For comparison, a 

late Holocene slip rate is available from a site on the San 

Francisco peninsula, 11 km south of San Francisco. 

There, displaced and abandoned stream channels and 

relict alluvial deposits (fig. 4.6) were explored and 

described by Hall (1984; N.T. Hall, written commun., 

1986). The alluvium, "C dated at 1,130Â±16 yr, accumulated along a northwest-flowing stream and behind a 

northwest-trending shutter ridge along the main fault 

trace. Alluviation ceased when fault slip, unusually high 

seasonal streamflow, or stream piracy caused the stream 

to abandon its fault-extended course and to carve a 

shorter, steeper, southwest-flowing channel across the 

fault and directly down the slope. This younger channel 

now exhibits 13.5 m of right-lateral offset where it is 
88 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

crossed by the fault; about 2.7 m of this total offset can b 

attributed to fault displacement during the 1906 earti 

quake. Hall's (1984) analysis of this sequence of depos 

tional, erosional, and faulting events yields a rninimui 

slip rate for the past 1,200 yr of about 1.25 cmlyr. 

San Jose 

MOUNTAINS 

0 

u 

10 20KILOMETERS Bay 

IGL'RE 4.5.-Fault-displaced geologic units on the San Francisc 

peninsula provide a basis for estimating the rate of strike slip on thi 

sectionofthe San Andreas fault-about 1.5 cmlw forthe Quaternan 

Critical field relations indude: A. <iutcrop belt of the marine Mere 

Furmation. 2.0-0.4 Ma, truncated and faulted qsinrit older ixick.i fc 

11 km on the southwest; EB', faultdisplaced parts of a 3.0- t 

2.0-Ma, late Pleistocene marine embayment (Addicott, 1969); C-C 

3.0- to 1.0-Ma "onmarine conglomerate of the Carte Madera facies i 

the Santa Clara Formation west of the fault and its source area eas 

of the fault (Cumming~, IfXVi,; I). *it" where 1.2-ka (late Ilul~cen< 

slip rate of abuut 1.25 cm/y (Hall, IYblJ \\'as mcaaurcd. Quatcrnar 

dip ha.. akooccurred on the San Grcgorio, Ilavwiird, and Calaven 

faults. Faults dotted where concealed by water. 

F1,:rRF. 4.6.-InterpreHitive block diagrams, ba^d on rplationh dcwribed by Hall(lW1, forai-iU-onthe% Andreaafault 11 kmyouth 

ofsnii Francisco. Rate of fault slip, about 1.25 cm'yr, derived from 

observed 13.5-m stream offset,which postdates alluviation along 

original (precaptm) streamcome. A, Stream, flowing obliquely 

downslope from S to S', is partly diverted by ridge (R) where stream 

crosses the San Andreas fault (F-F'). Gully (G) heads southwest of 

fault, flowing toward observer. B, Displacements (Dl, caused by fault 

slip, offset stream and lengthen its channel; gully erodes headwad 

across fault. C, Low gradient of stream channel causes alluviation 

upstream from fault, facilitating capture of its headwaters by gully 

and ~ abandonment ~~~ ~ of streamcourae S-S'. Date of alluviation. 

l,l:idr.ll%,yr R.P., from dt-lrital fharciffd dawri by "C nit-thoflb. D, 

Renewed fault di5placements (D') &set captured Aream (S-S"I by 

13.5 m, an amcunt that includes 2.7 m of slip in 1906. 
4. QUATERNARY DEFORMATION 89 

OTHER FAULTS 

The rate of Quaternary strike slip along the northern 

section of the San Andreas fault south of San Francisco, 

about 1.5 cmlyr, is less than either the 5.1 cdyr 

attributed to the plate boundary or the 2.5 to 3.5 

cdyr observed along more southerly segments of the 

fault. Much of this slip deficit can be accounted for by 

distributed slip on several other northern California 

faults that nearly parallel the main trace of the San 

Andreas. 

One of these faults, the San Gregorio, is the northernmost of a 400-kn-long set of coastal faults lying southwest of the main trace of the San Andreas. The San 

Gregorio fault extends northward from Monterey Bay, 

joining the San Andreas about 20 km northwest of San 

Francisco, near Bolinas Bay. Where it cuts the coastline 

near Point Atio Nuevo, the San Gregorio is a complex, 3- 

to 5-km-wide zone of near-vertical strike-slip and northeast-dipping reverse faults. Faults in this zone offset 

stream channels incised into marine and nonmarine 

strata that were deposited on a 105-ka-old wave-cut 

platform; the faults also offset the shoreline angle of this 

and at least one older wave-cut platform (fig. 4.7). 

Evidence from offset shoreline angles and faulted strata 

indicates a late Pleistocene, right-lateral slip rate of 0.6 to 

1.1 cm/yr across the fault zone (Weber and Cotton, 1981, 

p. 16, 72-75); Hamilton and others (1979), however, 

argued for rates of only 0.1 cdyr in the same area. 

If the higher slip rate on the San Gregorio fault and the 

rates estimated for the San Andreas fault south of San 

Francisco are correct, Quaternary slip on the San Andrea north of its junction with the San Gregorio must be 

about 2.6 cmlyr-the sum of the rates on the two 

separate faults. This rate agrees closely with that previously cited for the San Andreas near Point Arena 

(Prentice, 1989). 

Additional slip occurs along faults northeast of the San 

Andreas, as shown by geomorphic evidence, alignment of 

earthquake epicenters, and, on some faults, gradual fault 

creep amounting to as much as 1 cm/yr. Direct geologic 

evidence of Quaternary slip on such faults as the Hayward, Calaveras, and Green Valley is abundant, but the 

age and crossfault correlation of displaced stratigraphic 

markers are uncertain, and Quaternary slip rates are 

only weakly constrained. These rates, however, can be 

estimated from such other criteria as modern creep rates, 

geologic slip rates for displaced pre-Quaternary markers, 

and geomorphic expression relative to that along the San 

Andreas fault. 

On such evidence, the most active fault trend, with 

Quaternary slip rates estimated to range from 0.5 to 1.0 

cmlyr, is that defined by the southern section of the 

Calaveras fault and by the Hayward, Rodgers Creekiealdsburg, and Maacarna faults (fig. 4.3). Together, 

,hese faults extend for 375 km as a series of chiefly 

ight-stepping breaks that exhibit abundant geomorphic 

:violence of recent movement. The seismically active 

keen Valley fault, locally as well defined as the Rodgers 

;reek and 27 km farther east, offsets cultural features 

,hat record right-lateral creep of 0.5 cdyr between 1922 

nd 1974 (Frizzell and Brown, 1976). 

The east boundary of the northern section of the Sau 

indreas fault system follows a series of faults that trend 

lorth-northwest, exhibit recognizable components of 

ight-lateral strike slip, and approximately separate the 

;oat Ranges from the Great Valley. This discontinuous 

WCRE 4.7.-Sketch map of Point Ano Nuevo area, northern Califora, showing faultdisplaced shoreline angles in the San Gregnrio fault 

zone. Broken segments of shoreline angles indicate possible alternative locations. This figure, simplified from Weber and Cotton (1981), 

shows the difficulty of accurately measuring displacement where 

faults intersect a relict shoreline at acute angles (see fig. 4.8A). 
90 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

and somewhat diffuse belt of faulting includes, from 

north to south, the Stony Creek, Greenville, and Ortigalita faults. No obvious surface faulting bridges the 

100-km interval between the Stony Creek and Greenville 

faults, but major earthquakes in 1892 at Vacaville (lat 

38'12.5' N., long 12135' W) and Winters (lat 38' 18.5' N., 

long 121Â°34 W.) may signify fault continuity at depth on 

this trend. 

The historically aseismic Stony Creek fault follows and 

locally truncates a major detachment surface (the Coast 

Range thrust) that has been rotated to vertical or 

overturned. It now exhibits a steep east-facing scarp 

along the front of the Coast Ranges and separates a 

mountainous Franciscan terrane on the west from less 

deformed rocks on the east. Stream-terrace and fan 

deposits of Quaternary age east of the fault terminate at 

its scarp and are gently tilted and dissected by erosion. 

Sparse geologic evidence of renewed Quaternary faulting 

favors dominantly vertical slip, up to the west, but a 

right-lateral strike-slip component is indicated by asymmetric cusps and bays in the mapped trace ofthe fault and 

by nearby southeasterly-plunging folds that deform rocks 

at least as young as late Pliocene (Fox, 1983, p. 22). 

System-boundary faults to the south exhibit both local 

seismic activity and evidence of Quaternary slip. The 

Greenville fault, here (fig. 4.3) combined with the Marsh 

Creek and Clayton faults, displaces Quaternary streamterrace deposits and soils, as well as fold axes in middle 

to upper Tertiary marine deposits. Right-lateral slip 

rates on these faults range from 0.01 to 0.30 cdyr, but 

most Quaternary rates are less than 0.1 cm/yr (Sweeney, 

1982, Wright and others, 1982). These relatively low 

rates are consistent with weak geomorphic expression 

along this fault trend. 

FOLDING, UPLIFT, AND VOLCANISM 

Northwest- or east-west-trendmg folds deform much 

of the San Andreas fault system except for the relatively 

undeformed Sebastopol block (Fox, 1983) between the 

San Andreas fault and the fault trend defined by the 

Hayward, Rodgers Creek, and Maacama faults. The 

relative structural simplicity of this block is most evident 

between San Francisco Bay and about lat 3930' N. (fig. 

4.3), where flat-lying upper Miocene and Pliocene strata 

of the Wilson Grove (Fox, 1983) and Ohlson Ranch 

(Higgins, 1960) Formations extend across much of its 

surface. The erosional surface of this part of the block, 

the Mendocino Plateau (Lawson, 1894; Wahrhaftig and 

Birman, 1966, p. 323, Fox, 1983, p. 22-24), is a planar or 

gently warped upland with accordant ridge crests at 

elevations of 300 to 600 m. Fox (1983) contrasted the 

apparent stability of his Sebastopol block with a more 

typical, highly deformed region to the east, his Santa 

Rosa block. 

In the Santa Rosa block and in other more easterly 

parts of the Coast Ranges, Quaternary folding and 

faulting have left uplifted or anticlinal ridges and downwarped or downfaulted basins. Some of the larger basins, 

now filled with locally derived sediment of Pliocene and 

Quaternary age, are the southern arm of the San 

Francisco Bay, Napa and Livermore Valleys (lat 3820' 

N., long 12220' W., and !at 37.40' N.,long 121W W.), 

and the topographic basin surrounding Clear Lake (lat 

39-01' N., long 12230' W.). Smaller aligned or linear 

basins that follow northwest-trending faults or synclinal 

folds (fig. 4.3) clearly are structurally controlled. 

Rates of folding and uplift are best known for coastal 

regions. Marine terraces between San Francisco and 

Monterey Bay indicate general uplift of 120 to 180 m 

during the past 0.5 to 1 m.y. (Helley and others, 1979, p. 

18)-an uplift rate of about 0.02 cdyr. Local variation in 

uplift is shown by the deformed shoreline angle and 

wave-cut platform of the youngest (82 ka) terrace (fig. 

4.8); near Half Moon Bay, the wave-cut platform has been 

warped by northwest-trending folds into a surface that 

exhibits 60 m of structural relief over a distance of about 

11 km (Lajoie, 1986, fig. 6.21). About 10 krn west of Santa 

Cruz, this terrace is about 14 m above modern sea level 

(Bradley, 1957; Bradley and Griggs, 1976; Hanks and 

others, 1984, p. 5776-5777), indicating a local uplift rate 

of 0.04 crnlyr. 

Coastal uplift south of San Francisco contrasts with 

Quaternary crustal subsidence 25 to 30 krn inland and 

east of the San Andreas fault (figs. 4.3,4.5). Estuarine, 

stream-laid, and freshwater-swamp deposits of Quaternary age underlie the northwest-trending structural 

basin containing the San Francisco Bay and the Santa 

Clara Valley (1st 37-10' N., long 121Â°40 W.) to depths 

greater than 200 m below sea level. This prolonged 

Quaternary downwarping (Wahrhaftig and Birman, 

1965) locally attained subsidence rates of 0.02 to 0.04 

cm&r (Atwater and others, 1977) during post-Sangamon 

time. 

Despite such local variations in the amount and orientation of crustal processes, Quaternary uplift (fig. 4.9) 

prevails throughout the northern Coast Ranges. Regional uplift, at least partly of Quaternary age, ismatest and 

most evident east of the IIa\-wanl-Rode'ei's Crock-Maacama fault trend. Summit elevations north of Clear Lake 

exceed 1,500 m (Wahrhaftig and Birman, 1965, fig. 9), 

and those in the Diablo Range, south of the Livermore 

Valley, range from 600 to 1,200 m (Christensen, 1965, pi. 

1). Both of these colinear uplands exhibit high relief, 

evidence of rapid downcutting, and cores of the emergent 

Franciscan complex flanked by outward-dipping strata of 

Pliocene and Quaternary age. 
4. QUATERNARY DEFORMATION 91 

Volcanic rocks of late Cenozoic age, widely distributed 

between the San Pablo Bay and Clear Lake, define a 

triangular, northwest-trending outcrop area, 120 km long 

by 35 km wide, with its meet acute, north apex at Clear 

Lake. These rocks are assigned to two geographically 

separated extrusive seauences: the late Miocene and 

~iocene Sonoma ~olcaics to the south, and the younger, more areally restricted Clear Lake Volcanics (fig. 4.3) 

to the north. The Clear Lake Volcanics is almost wholly 

A 

4 KILOMETERS 

OI Plunging syncline 

PACIFIC 3 OCEAN 

FIGURE 4.8.-Local variability in coastal uplift shown by deformed 

shoreline angle and wave-cut platform. A, Coastal geomorphk 

features used to measure Quaternary deformation. SLA, shoreline 

angle carved by marine erosion at present sealevel, WCP, wave-cut 

plah, WCP' and SLA', elevated surface and Shoreline am, 

respectively, of an older wave-cut platform now mantled by marim 

terrace deposits and terrestrial sediment (fight stippling). Normal 

seaward slope of wave-cut platform, exaggerated in diagram, is about 

lO-2". B, Folding and right-lateral faulting along the Seal Cove 

branch ofthe San G&o fault zone near Half Moon Bay (HMB), a 

shown by warped surface of an 82-ka wave-cut platform anc 

deformation of its landward edge, the shoreline angle (heavy line). 

Structure contours (long dash on land, short dash off shore) or platform are derived, from shallow boreholes and BeiBtnic-redata. Arrows on fault (dotted where concealed) indicate direction 01 

relative movement: U, upthrown side; D, downthrown side. Modiflet 

from Lajoie (1986). C, Platform pmfiles on fault plane (dotted lines 

E, east, W, west), showing vertical separation across fault ant 

deformation relative to shoreline angle (heavy line). Modified fron 

LKl(iie (1986). 

f Quaternary age. Flow surfaces, volcanoes, cinder 

ones, obsidian domes, and craters retain their original 

1 Sea level 

u 

EXPLANATION 

2400 rn 

CO 

FIGI'RE 4.9.-Generained surfaw on sumtiut altitudes tGipfelflur) in 

the Coast Ranees north of San Frandaco. showing malor faults with 

Quaternary activity intheSan Andreas faultt-yatem-that Dart ofthe 

&I southof ~aoi Mendodno and west of the Sacramento Valley. 

C&t7ur appro&tely defines refions of Quaternary uplift and 

~tal blocks bounded by strike- and dip& faulting (dotted whew 

coded). Gewralufd from Wahrhaftic and Birman (1966). 
THE SAN ANDREA8 FAULT SYSTEM, CALIFORNIA 

constructional form, and the Geysers Hot Springs area on 

the west side of the main volcanic field attests to 

continuing geothermal activity. Lava flows range in 

composition from olivine basalt to rhyolitic obsidian, and 

pyrdatic deposits are nearly as varied. The oldest 

welldated rocks in the Clear Lake Volcanics yield K-AT 

ages near 2.1 Ma; the youngest-ash beds in lake 

sediment beneath Clear Lake-yield 14C ages of about 11 

ka (Hearn and others, 1976; Sims and Rymer, 1976). 

The sustained late Cenozoic volcanic episode recorded 

by the Sonoma and Clear Lake Volcanics is unique within 

the San Andrea fault system, although similar but less 

extensive Quaternary volcanism marks the northern end 

of the Gulf of California ride-transform svstem in the 

Salton Trough Oat 33-19, N., long 115'50' w.). 1 

CENTRAL SECTION OF 

THE SAN ANDREAS FAULT SYSTEM 

SAN ANDREAS FAULT 

From the Gabilan Range southward to the Transverse 

Ranges, the San Andreas fault itself dominates the fault 

system. Except for the San Gregorio-Hosgri set of 

coastal faults, branch faults are few and short, and 

exhibit less evidence of Quaternary activity than those 

farther north. This relatively simple pattern changes at 

the northern margin of the Transverse Ranges, where 

major folds and faults trend more east-westerly and 

incorporate large components of reverse and thrust 

motion. In this chapter, faults in the western Transverse 

Ranges and southern Coast Ranges are grouped together 

as components of the central section of the &I Andreas 

fault system. This combination of twodierent structural 

and physiographic provinces stresses that the transition 

from compressional structures in the Transverse Ranges 

to strike-slip structures in the Coast Ranges is more 

gradual and less sharply defined than the physiographic 

boundary. 

This section of the fault system, 340 krn long by 140 km 

wide, extends from the northern Gabilan Range to the 

southern Transverse Ranges, near lat 34' N. (fig. 4.10). 

Besides the San Andreas, faults with Quaternary slip 

include the Hosgri, Sail Simeon, and Rinconada in the 

north; the Big Pine, Ozena, South Cuyama, and Modes 

to the south and in the Coast Ranges; and the eastwest-trending reverse and thrust faults in the Transverse Ranges. Several faults in the Transverse Ranges 

continue offshore into the Santa Barbara Basin. 

The San Andrea fault trends S. 40' E. from the 

northern Gabi Range to the Carrizo Plain, there, it 

bends eastward and ultimately attains a S. 75' E. trend 

in the Transverse Ranges. North of the bend, it follows 

a series of nearly aligned stream valleys, which separate 

the Diablo Range and parts of the Temblor Range on the 

s,. 

EXPLANATION 

Alluvial md estuarine deposits (Quaternary)Ã‘Chiefl fill: may also include some deoosiu of ,_____, Pliocene age 

Bedrock (Tertiary and older)-Varied rock types 

Fault exhibiting evidence of Quaternary displacement 

-Sawteeth on upchrown block of reverse or thrust 

fault; dotted where concealed by water 

FIGURE 4.10.-Sketch mso of southern Coast ram^ and western 

Transverse Ranges, southern California, showingfaulte with Quaternary activity and bin deposits in the central section of the San 

Andreas fault system. Fault patterns aregeneralized, and only major 

faults are shown. 
4. QUATERNARY DEFORMATION 

northeast from the Gabii Range, Cholame Hills (lat 

SSW N., long 120Â°18 W.), and Caliente Range (la1 

35-02' N., long 119'46' W.) on the southwest (Brown, 

1970; Vedder and Wallace, 1970). An actively creeping 

length of the fault, with historical slip rates as high as 3.4 

crnlyr, extends 160 km southward from the northern 

Gabilan Range to the south end of the Diablo Range. 

To the south, along the boundary between the Transverse Ranges and the Mojave Desert, the fault continues 

as a series of closely aligned echelon breaks, in a narrow 

(0.5 km wide) linear trench. This segment, which also 

displays abundant geomorphic evidence of Quaternary 

slip (Ross, 1969), is currently locked and seismically 

quiet. 

Hill and Dibblee (1963, p. 446), in proposing large 

sustained strike slip on the San Andreas fault, described 

fault-juxtaposed gravels of similar age but different 

lithologies in the area between the Temblor Range and 

the Transverse Ranges (fig. 4.11). Pebble gravel in the 

Paso Robles Formation, of Pleistocene and Pliocene age, 

consists chiefly of porcellanite and siliceous shale clasts 

derived from the Miocene Monterey Formation of the 

western Coast Ranges and the Temblor Range. These 

distinctive Paso Robles gravel beds are most extensive 

southwest of the San Andreas fault, but northeast of the 

fault they extend as a narrow band to about lat 36Â N. For 

several kilometers near this latitude, they are faulted 

against Pleistocene gravel containing clasts of plutonic 

and metamorphic rock and indurated sandstone, derived 

from the San Emigdio Mountains to the south. According 

to Hill and Dibblee (1953, fig. 2), the depositional contact 

between the two gravel lithologies is displaced about 16 

km from its prefaulting location. 

Farther north, between lat 3630' and 36"46' N., 

Galehouse (1967, p. 974-976) described different chat 

compositions and heavy-mineral assemblages in Paso 

Robles strata on opposite sides of the fault. These 

differences are reasonably and consistently resolved if 40 

km of right-lateral fault slip is removed, restoring the 

offset fades to their original positions. 

These displaced stratigraphic relations imply Quaternary slip of 16 to 40 km, but the uncertain age of the 

faulted strata makes estimates of the slip rate unreliable. 

More accurate measurements of fault slip have been 

obtained at three localities north of the Transverse 

Ranges: two in the creeping segment and one in the more 

southerly, currently locked segment that produced the 

1857 earthquake. 

At the northernmost locality (lat 3634.9' N., long 

121Â°10.4 W.), along the San Benit0 River near the east 

boundary of the Gabiian Range, the fault displaces the 

risers of river terraces (fig. 4.12), exhibiting64 m of right 

dip. Trenches on the terrace near the fault disclosed a 

fluvial scarp and a distinctive sand bed, both offset about 

17.7 m by faulting during the past 790Â±14 14C yr 

TRANSVEMSE RANGES 

na,hara 

EXPLANATION 

[Distribution of rock units shown only 

for region adjacent to San Andreas fault1 

Coast Ranges 

Monterey Formation lMiocene)Ã‘Well-bedde siliceous shale, 

porcellanite. and sandstone. Source unit for the type Paso 

Robles Formation of the southern Coast Ranges 

San Emigdio Mountdm 

mtamorphi~rocks and indurated sandstone. ~apped by 

some authors as Paso Robleb Fornation 

Plutonic, metamorphic, and sedimentary rocks IPaleogene 

and pre-TertiarybSource rocks for conglomerate near 

the San Emigdio Mountains 

?IGUEE 4.11.-Midway between Bakerafield and Santa Barbara, 

Calif., the San Andreas fault locally juxtaposes conglomerates of 

similar age but different litholcgies. Porcellanite and ailiceow shale 

clasts immediatelv northeast of the fault were derived chieflv from 

the folded and uplifted Monterey Formation in the southern~oast 

Range% plutonic, metamorphic, and nedimenm-j clasts across the 

fault lu the southwest werederived from the San Emigdiu Mouniaina 

of the northern Transverse Ranges. Total displacement and precise 

ace ofmnirlomerate units are unknown. but ma~d offset of at least 

16 kmreq&ssustained ~uaternary strike slip,L recognized by ill 

and Dibbled (1953). 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

(Perkins and others, 1989). The slip rate calculated from 

these data, about 2.2 dyr, agrees closely with creep 

rates currently observed along this section of the fault. 

In Bitterwater Valley Oat 36-23.9' N., long 1W58.9' 

W.), 27 km southeast of the San Benito River locality, 

fault-displaced buried channels in an alluvial fan of 

Quaternary age record a slip rate over the past 1 ka of 

about 2.8 dyr (Cotton and others, 1986), but the slip 

rate, as well as its distribution within the fault zone, has 

varied over time. The historical slip rate, 3.4 dyr, 

represents chiefly gradual fault creep and is based on 

offsets observed in fences built in 1908 and in channels 

that presumably date from about 1885. According to 

cotton and others (1986), at least one episode of slower 

movement is evident, and, during the past 1 ka, slip 

across the 20-m-wide fault zone has followed different 

breaks. 

Sieh and Jahns (1984) measured and dated late Quaternaiy displacement on the San Andreas fault at Wallace Creek (lat 35'16.3' N., long 11949.7' W.), near the 

southwestern margin of the Temblor Range. Wallace 

Creek," an informal name (Sieh and Jahns, 1984, p. 896) 

for a previously unnamed drainage, does not appear on 

published U.S. Geological Survey topographic maps. At 

this locality, dated fan and channel deposits, stratigraphic and geomorphic relations, and fault offsets of both 

abandoned and existing stream channels document 

128Â± m of strike-slip displacement (fig. 4.13). This 

record of cumulative fault slip began with incision of the 

existing channel of Wallace Creek sometime after 

3.680k155 vr aso and ended with the nearlv 10 m of fault " - 

sip that accompanied the 1857 earthquake. Because the 

3,680-yr 14C age is measured relative to A.D. 1960, the 

appropriate timespan for the measured slip is 

3,680-(1,950-1,857), or 3,587 yr. The minimum rate is 

FIGTOE 4.12.-Fault displacement (dl of stream-tei~ace mere CI"I') 

postdates old terrace deposits (OTD), which underlie old terrace 

surface (GT3). Dating of terrace deposits gives a maximum age for 

disolacement and thus a minimum din rate for the fault farrows 

indicatedirection of relative movement:fault dotted *Â¥her concealed 

by Btreaw). The udilly eple relations shown in this sketch 

eliminate the effects of multiple terraces, changing rates of lateral 

erosion by the stream, and other processes that, under normal field 

exposures, make analysis much more difficult. 

thus 128 m/3,587 yr, or 3.57k0.2 dyr; the maximum 

late, based on the stratigraphy and geometry of the 

abandoned channel, is 3.53Â±0. dyr. When these 

maximum and minimum values are combined and adjusted to incorporate uncertainty in the state of strain at the 

beginning of the time interval, a preferred slip rate (Sieh 

and Jahns, 1984, p. 891) of 3.39 Â 0.3 cmlyr is obtained. 

A similar but more complex analysis of 13,250-yr-old fan 

deposits displaced 475 m from their source gullies yields 

a slip rate of 3.58k0.5 cdyr (Sieh and Jahns, 1984, p. 

891-892) for the past 13,260 yr. The dose agreement of 

these rates suggests that the average rate of late 

EXPLANATION 

hannel alluvium 1-3 ka) 

5 L- High channel alluvium (11) lo 3 ka) 

Contact 

-- fault-Arrows show direction of relative movement 

-2250- Contour, in feet-Omitted near geologic contact 

FIGURE 4.13.-Slip on the San Andreas fault at Wallace Creek (lat 

35-16.3' N., long 119-49.7' W.) has displaced daud alluvial dewits, 

as well as stream channels and other geomorphic features. Trenches 

excavated across and parallel to fault at this locality disclose details 

of stratigraphy, origin of stratigraphic units, and their age as determined from detrital charcoal. Displacements of different ages 

yield aunSar average dip rates of about 3.4 crn/yr. Original source 

drainages for older fan alluvium are near or southeast of right margin 

of future. Simplified from Sieh and Jahns (1984, Be. 2). 
4. QUATERNARY DEFORMATION 95 

Quaternary fault slip has been relative uniform at this 

locality. 

Horizontal slip rates on the San Andreas fault are less 

well constrained in the western Transverse Ranges. 

Davis (1983, p. 368) and Rust (1982) independently 

estimated slip rates in the range 2-5 cdyr for areas 70 

km apart and on opposite sides of the junction of the 

Garlock fault with the San Andreas. These estimates 

involve substantial uncertainty in the amount and (or) 

timing of slip; they also apply to different parts of the 

Quaternary-no more than the past 400 ka for displaced 

gravel deposits in the San Emigdio Mountains (Davis, 

1983), and no more than the past 2 ka for offset landslide 

deposits 32 km southeast of the Garlock-San Andreas 

fault junction (Rust, 1982). These rates, however, are 

broadly consistent with those measured to the north. 

Moreover, they indicate that much Quaternary strike slip 

in the western Transverse Ranges has followed the main 

trace of the San Andreas-a departure from pre-Quaternary history, when the San Gabriel fault to the southwest 

accommodated much of the fault slip (Crowell, 1975). 

OTHER FAULTS 

Other faults with known or probable Quaternary slip in 

the central Coast Ranges include the Rinconada, Big 

Pine, Ozena, South Cuyama, Morales, and a coastal fault 

defined by its aligned but separately named segments: 

Hosgri, San Simeon, Sur (lat 36"lO.E' N., long 12132' 

W.), Monterey Bay (lat 36'40' N., long 122%' W.), and 

San Gregorio. For most of these faults, the evidence of 

Quaternary activity is sparse, discontinuous along strike, 

and in places arguable. It includes weakly expressed 

scarps and drainage offsets, displaced deposits of late 

Pliocene and Quaternary age, and historical seismicity 

along some fault trends. The geologic and geophysical 

evidence for Quaternary activity is somewhat more 

convincing for the Hosgri fault and its northern, colinear 

counterparts than for faults farther inland. 

The Hosgri and similarly aligned coastal faults trend 

north-northwesterly from near Point Conception (Steritz 

and others, 1986) to the latitude of Monterey Bay (Silver 

and Normark, 1978), where they continue northward as 

the San Gregorio fault. For most of its length, this coastal 

fault zone is a complex, multistrand break a few kilometers off shore; it is recognized and mapped from juxtaposed stratal units interpreted from seismic-reflection 

profiles and truncated magnetic anomalies. Locally, seafloor scarps and displaced Holocene deposits, interpreted 

from the seismic-reflection profiles, confirm its Quaternary activity (Wagner, 1974; Leslie, 1981). 

The San Simeon fault (fig. 4.14), which may be an 

onshore segment of the Hosgri (Leslie, 19811, is also a 

multistrand break; it cuts and displaces late Pleistocene 

marine terraces and Holocene dune sands near San 

Sieon Point (Weber, 1983). Although faulted shorelineangle geometry and the correlation of wave-cut platforms 

acrossthis fault are subject to some uncertainty, Weber 

119s:f. 1). 56-59) and Wanson and other? 11987) sumrested . ,A . . -- 

a horizontal right-slip component of about 0.5 cndyr. 

Earthquake focal mechanisms along the San Simeon fault 

indicate a dip of 55' E. and nearly equal components of 

reverse and strike slip (Eaton, 1984). Right-oblique slip 

at a rate of about 0.5 cndyr on the Hosgri fault is 

indicated by earthquake focal mechanisms, sea-floor 

scarps, the distribution of late Quaternary wave-cut 

platforms, and faulted seismic-reflection horizons of late 

Cenozoic age. Much of this evidence is documented from 

investigations near the Diablo Canyon Powerplant (Pacific Gas and Electric Co., 1988). where it has been 

interpreted as pure strike slip. 

In the western Transverse Ranges, faults with Quaternary displacement include some, like the San Andreas 

and San Gabriel, that exhibit chiefly strike slip and 

others, trending east-west or southwest, that exhibit 

reverse or oblique slip. The left-lateral Santa Ynez fault 

is an exception to the more general pattern of reverse and 

thrust faults. The reverse- and oblique-slip faults in the 

Transverse Ranges are complicated by splays, offsets, 

and changes in strike, hut from north to south they define 

at least four major continuous, or nearly continuous, fault 

zones of different lengths: the Santa Ynez (130 km); the 

More Ranch-Mission Ridge-San Cayetano (110 km), from 

the coast near Santa Barbara, eastward; the Mission 

Hills-San Fernando-Sierra Madre (56 km), along the 

southwest side of the San Gabriel Mountains; and the 

South Frontal (250 km). Many of these faults exhibit 

evidence of Holocene movement, and the San Fernando 

fault produced the surface rupture and destructive M= 

6.6 earthquake of February 1971. Some of these faults, 

and others with east-westward trends, extend offshore 

into the Santa Barbara Channel (1st W09' N., long 

11933' W.), where they also cut Pleistocene or Holocene 

deposits (Clarke and others, 1985). Ziony and Yerkes 

(1985, p. 43-60) provided more detailed descriptions and 

maps of the late Quaternary faults in this and nearby 

areas; they also noted (p. 44) a profound increase in 

deformation rates in and near the Transverse Ranges 

during the past 750 ha. This pulse of orogenic activitymanifested by the growth of folds, uplift, and accelerated 

slio on reverse and thrust faults-coincides with middle 

~iistocene and younger deformation observed to the 

south (Woodford and others, 1954, p. 77-78) in the Los 

Angeles Basin. 

Maximum and minimum slip rates for many of the 

faults in the Transverse Ranges are deduced from offsets 

observed in drainages, fan surfaces, river terraces, and 

wave-cut platforms. Reported by numerous investiga- 
96 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

per year, but dip-slip components as great as 0.9 cdyr 

are reported (Rockwell, 1988) for the San Cayetano fault. 

Observed slip rates may represent only a fraction of the 

total structural growth of the ranges because much of the 

late Quaternary deformation is by folding and uplift. 

as near the northboundary of the Transverse Ranges 

(fig. 4.10). Neither of these faults exhibits a clearly 

defined junction with the San Andreas, but together they 

tors (Clark and others, 1984). these rates of chiefly 

oblique slip cluster around a few tenths of a centimeter 

FIGURE 4.14. -Sketch maps and columnar section of San Simmn area, 

central California, showing selected examples of Quaternary faulting 

on the San Simmn fault zone. A, Displacement of shoreline angle 

(hachured line) of 124-ka wave-cut platform; other younger and older 

marine terraces, mapped by Weber (1983). are not shown. B, 3-m 

offset (d, and rip1 of stream channel near circled locality in figure 

4.14A. C, Faulted paleosola and possibly faulted modem soil in 

Although most Quaternary faulting south of the Gabilan Range is confined to structural blocks west and south 

of the San Andreas, two major faults, the Garlock and 

White Wolf. diverse northeastward from the San Andresw Ground surface NE 

\\ - .0 

- 

% 

- 

_f 

- 

- 

I 

/ 

EXPLANATION 

Sand and gravel oa 

Silty clay 

Black clay 

Sandy clay 

?- -- Fault-Dashed where approximately located; 

dotted where concealed: queried where 

- 

uncertain 

Soil boundary 

trench exposure near circled locality in figure 4.144 letters on right 

designate soil classifications. Quaternary strike-slip rate for this fault 

at the surface is about 0.5 cmlvr. but earthmahe focal mechanisms 

(Eaton, 19841 indicate right-reverse slip at 6.6-!m depth. From 

Pacific Gas and Electric Co. (unpub. data, 198% see Pacific Gas and 

Electric Co., 19881. 
4. QUATERNARY DEFORMATION 97 

bound the horstlike block of Sierran basement that makes 

up the Tehachapi Mountains. The White Wolf, a southeast-dipping reverse fault that was the source of the Kern 

County M=7.7 earthquake of July 1962, exhibits geomorphic evidence of recent slip at the range front but is 

concealed to the west beneath the southern San Joaquin 

Valley. The Garlock, devoid of major historical seismicity, exhibits clear evidence of left-lateral Quaternary slip 

(dark, 1973) that increases eastward from 0.2 dyr at a 

point 55 ton east of the San Andreas to 1.1 cdyr at a 

point 76 km to the east (Clark and others, 1984). 

POLDING AND UPLIFT 

Typically, Quaternary folds are more intense and 

better defined in thick Tertiary marine sequences than 

near or above such basement blocks as the granitic and 

metamorphic complexes of the Santa Lucia, La Panza, 

and Gabiian Ranges or the Franciscan core of the Diablo 

Range. Thus, some of the best expressed Quaternary 

folds are in the western Transverse Ranges (Bailey and 

South 

KILOMETERS N. 10Â W. 

Jahns, 1964, p. 96-98), the west side of the San Joaquin 

Valley (Harding, 19761, the Santa Maria Basin (Woodring 

and Bramlette, 1950, p. 109-116), and the southeast flank 

of the Santa Lucia Range (Compton, 1966). Harding 

(1976) used subsurface geologic data from oil fields 

northeast of the San Andreas fault to demonstrate that 

folding has moved outward from the San Andreas over 

time-the more northeasterly folds along the west side of 

the San Joaquin Valley being the youngest. 

Rates of Quaternary folding in the westernmost Transverse Ranges exceed those elsewhere in the fault system. 

On the coast near Venture (lat 34W N., long 119"18' 

W.), Pleistocene basin deposits more than 4 ton thick 

have been intensely folded and faulted during the past 

200 ka, attaining dips of 32Â¡-45 in the southern limb of 

the Venture Avenue anticline (fig. 4.15); uplift rates of 

about 1 dyr (Yeats, 1977; Lajoie and others, 1982) 

characterize the axis and southern limb of this fold. Uplift 

and folding elsewhere in the western Transverse Ranges 

may be slower but probably still exceeds deformation 

rates elsewhere in the fault system. 

VENTURA AVENUE 

ANTICLINE 

FIGURE 4.15.-Many drill holm penetrate welldated Quaternary cm\T at the fold emst during the past 200 ka. Nearly all of this uplift 

strata near the Ventura Avenue anticline. Besides the provincial is Quaternaryinage: most ie late Quaternary, as this highly generalized 

microfaunal control shown here (3 email circles), Lajoie and others ficure shows. Modified from Yeats (1982, Gg. 7). Contacts and faults 

(1982b) employed amarine terrace (0.1-0.08 Ma) and 3 ash beds (0.6, dashed whereapproximately located; arrowson faults indicatedirection 

0.7, and 1.2 Ma) in the "San Pedro," "Santa Barbara." and "Pim" of relative movement. C.H., corehole; LP, lower Pliocene; UP, 

Formations of Yeats (1982) to estimate an average uplift rate of 1.26 upper Pliocene; 1 and 6, provincial microfaunal markers. 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

Evidence for Quaternary uplift in this part of the San 

Andreas fault system is abundant and varied. Highstanding ranges (Diablo, Santa Lucia, La Panza, and 

Transverse) emerge from a cover of their own eroded 

debris, represented by the Paso Robles Formation of the 

Coast Ranges, the Tulare Formation of the San Joaquin 

Valley and northern San Ernigdio Mountains, and the 

Saugus Formation (and equivalent marine deposits) of 

the western Transverse Ranges. These upper Tertiary 

and Quaternary strata consist chiefly of alluvial debris, 

which has been deposited in fans, channels, and flood 

basins. Although the stratified deposits have been tilted, 

uplifted, and eroded (Bailey and Jahns, 1954; Christensen, 1866; Compton, 1966; Yeats, 1977; Davis, 1983). the 

map pattern of their remnants documents regional patterns of uplift. 

Christensen (1966) employed such evidence (fig. 4.16) 

to document uplift of 600 m in the Diiblo, Santa Lucia, 

and La Panza Ranges and to establish general systemwide uplift, north of the Transverse Ranges, of about 300 

m. This broad pattern of late Pliocene and Quaternary 

uplift is reversed chiefly in the Salinas and Santa Maria 

Valleys Oat 36'09' N., long 121W W., and lat 34%' N., 

long 120'15' W.), Kettleman Plain (lat 35'55' N., long 

120W W.), and southern San Joaquin Valley, where 

elongate deep (100-900 m) but local basins have formed at 

the edge of basement blocks. 

Late Pliocene and Quaternary uplift of the northern 

San Emigdio Mountains totals about 1,800 m (Davis, 

1983, p. 299-300), the present average height of a 

Pliocene erosion surface preserved on accordant ridge 

crests near the summit of the mountains. Relatively 

rapid uplift of the western Transverse Ranges was also 

accompanied by local downwarping of narrow basii-for 

example, the offshore Santa Barbara Basin and the Santa 

Clara River valley (fig. 4.16) east of Ventum 

Evidence that uplift is still continuing is found in 

raised, paired stream terraces and youthful drainage 

patterns in the Santa Lucia, Diiblo, and western Transverse Ranges, in elevated Holocene marine terraces near 

Venture (Lajoie and others, 1982), and in the radially 

segmented slopes of Holocene alluvial fans that record 

progressive episodes of deformation along the east front 

of the Diablo Range (Bull, 1964). 

SOUTHERN SECTION OF 

THE SAN ANDREAS FAULT SYSTEM 

SAN ANDREAS FAULT 

South of the Transverse Ranges, the San Andreas fault 

svstem (fie. 4.17) extends to the latitude of the Salton Sea as a broad belt of northwest-trending strike-slip faults 

bounded on the northeast by the San Andreas and on the 

iouthwest by the offshore Santa Cruz-Catalina Ridge 

ault zone. Between the Salton Sea and the Gulf of 

Salifomia, the San Andreas fault system merges with a 

smplex pattern of active ridge segments and transform 

hulls that continues beneath the gulf for more than 1,200 

on (see chap. 3). Alluvium of the Colorado River, lake 

leposits, and the waters of the gulf obscure the transition 

7 Sea level 

FIGURE 4.16.-Generalized topographic map of central California 

between the San Francisco Bay (SFB) and the La Panza Range, 

showing major faults with Quaternary activity in the San Andreas 

fault system. Modified from Christensen (1965, pi. 1). Fadte (lotted 

where concealed by water; hachuree on contours indicate area of 

dosed tow. Compare with figure 4.9. 
4. QUATERNARY DEFORMATION 99 

from the San Andreas pattern of deformation, dominated 

by northwest-trending strike-slip faults, to the ridgetransform pattern of the gulf, but current practice (see 

chap. 8; Crowell, 1981, p. 596-597) places the south 

boundary of the fault system near the latitude of the 

Salton Sea. Thus defined, the southern section of the San 

Andreas fault system is about 200 km long by 150 to 175 

km wide. 

Late Quaternary faulting in this region follows several 

major breaks, of which the San Andreas is the most 

easterly and most active. The dominant role of the San 

Andreas fades south of Cajon Canyon (kit 34*18' N., long 

ll7Â¡23'W.) where the San Jadnto fault diverges from it, 

and farther south along the south front of the San 

Bernardino Mountains, where it merges into the Banning 

and Mission Creek reverse faults (lat 33%' N.. low 

116'47' W., and lat 34W N., long 1163.3' W.). Despite 

such complexities, continuity of Quaternary faulting on 

the main San Andreas can be traced at least as far south 

as the Salton Sea. 

Large increments of Quaternary strike slip on the 

segment of the $an Andreas fault along the San Gabriel 

Range front near Pahndale (kit 34%' N., long 118W 

W.) were first recognized by systematic right-lateral 

stream offsets and by Pleistocene fan and terrace deposits displaced 2 to 8 km from their source regions across 

the fault (Wallace, 1949, p. 799-802; Noble, 1954, p. 46). 

Underlain by distinctive bodies of plutonic and metamorphic rocks, the ranges have contributed eroded debris to 

fans that spread across the range-front faults. SubseSari ClementPACIFIC OCEAN 

"and '^Ifc, 0 I~OKILOMETEW 

, , san oiego 0 

EXPLANATION 

Alluvial and esluarine deposits (Quaternary)-Chiefly basin fill: may also include some deposits of Pliocene age 

0 Bedrock (Tertiary and older)-Varied rock types 

Fault, exhibiting evidence of Quaternary displacement-Sawteeth on upthrown block of reverse or thrust fault; dotted where 

concealed by water 

FIGURE 4.17. -Sketch map of the eastern Transverse Ranges. Mojave Dem, and Salton Trough area, southern California, showing faults with 

Quaternary activity and bwim dewrite in thesouthern sectionof theSan Andreas fault system. Fault patternsaregeneralized, and only major 

faults are ahown. 
100 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

quent strike slip along these faults has displaced the 

distal parts of the fans laterally, separating them from 

their source rocks and opposing them against unlike rock 

types in nearby parts of the range front. Recent mapping 

(Barrows and others, 1986, p. 196-197) along the fault 

segment first described by Wallace (1949) and Noble 

(1954) constrains the slip rate there to values of 1.5 to 3.0 

cmlyr (fig. 4.18). 

Most historical slip on the San Andreas fault has 

accompanied earthquakes of Ma6. A remarkable record 

of earlier earthquake-related faulting has been interpreted (Sieh, 1978,1984; Sieh and others, 1989) from natural 

and manmade exposures across the fault at Pallett 

Creek, about 55 km northeast of Los Angeles, within the 

eastern part of the segment mapped by Barrows and 

others (1986). At Pallett Creek, Sieh described sandblows and other liquefaction structures, buried ecarps, 

and truncated fault strands in dated peat and alluvium 

that record (fig. 4.19) a history of 12 earthquakes during 

the past 1,700 yr. The latest 10 episodes of faulting 

APPROXIMATE AGE RANGE OF OFFSET 

PLEISTOCENE GRAVELS. IN MILLIONS OF YEARS 

EXPLANATION 

Oh Harold Formationlundmerentiatedl 

Qhs Harold Formation (schist and sandstone clast member) 

On Nadeau Gravel 

OM Boulder gravel (Link RockCreek) 

Shoemaker Gravel 

Qbb Boulder gravel (BIg Rock Creek) 

Qbp Couliter gravel IPallcl Creek1 

FI~~~~4.18.-Estimatsdrangesmstrike-sliprateona90-hsegment 

of the main San Andreas fault north of Los Angeles. Stratigraphic 

nomenclature (from Barrows and others, 1985, fig. 4) identifies 

lithologicallydiictive Quaternary deposits that have been horizontally displaced by fault slip on the San Andreas fault system. 

establish an average recurrence interval of about 132 yr 

(Sieh, 1989) but show an irregular, clustered distribution 

over time for large earthquakes on this section of the 

fault. 

Slip rates along the San Andreas south of Pallett Creek 

are lower than those in the central section of the fault 

svstem: in nart. this difference ie accounted for bv slm on " ?. , ". 

the Sail Jacinto and other, more westerly faults. 

A site on the San Andreas fault in Cajon Canyon flat 

34'16.4' N., long 11727.9' W.) about 100 km east of Los 

Angeles, establishes a consistent slip rate (Weldon and 

Sieh, 1985) for the late Pleistocene and Holocene time. 

Along Cajon and Lone Pine Creeks, the fault displaces 

terrace risers, buried and active stream channels, and 

landslides. The alluvial and swamp deposits have yielded 

14 ^C ages. These dated fault displacements, in cornbination with a reconstructed fluvial history of the site, 

provide four independent measurements of fault slip that 

document an average slip rate for the past 14.4 ka of 

2.5k0.4 cdyr (fig. 4.20; Weldon and Sieh, 1985). 

Harden and Matti (1989) reported slip rates that are 

less well constrained and possibly more variable for the 

San Andm fault near Yucaipa, 45 km southeast of 

Cajon Canyon. There, displaced alluvial fans yielded 

average slip rates of 1.4 to 2.6 cm/yr for the past 14 ka, 

2.2 to 3.4 cmfyr for the past SO ka, and 1.2 to 1.6 cdyr for 

the past 65 or 90 ka. These rates imply an accelerating 

rate of latest Quaternary slip, but uncertainties in 

measurements of displacement and age permit rates that 

are constant or even diminishing over time. 

Paired stream offsets along the southern branch of the 

San Andreas fault (lat 34W.5' N., long 1lTlO.O' W.), 

between Yucaipa and Cajon Canyon and in San Bernardin0 flat 34W' N., long llTl7' W.), indicate a maximum 

slip rate of 2.5 cmlyr (Rasmussen, 1982, p. 112) based on 

estimated ages of 30 and 50 ka for faulted alluvial units. 

If, as Rasmussen suggested, these are minimum ages, 

the actual slip rate may be somewhat lower. 

On the San Andreas fault and about SO km north of the 

Salton Sea flat 33'46.9' N., long 116O14.4' W.), KeUer and 

others (1982) mapped a total otbet of 700 m where two 

strands of the fault cut an alluvial tan. Soils developed on 

the fan surface indicate an age of SO to 20 ka, bracketing 

the slip rate between 2.3 and 3.5 cdyr. At a nearby 

locality, a rate of only a few tenths of a centimeter per 

year can be derived from displaced (1 m) sedimentary 

deposits associated with the latest highstands of former 

Lake Cahuflla, 700-300 yr B.P. (Sieh, 1981). These 

deposits, however, may be too young to record slip 

accompanying large earthquakes with recurrence times 

longe; than 700 6. 

The San Andreas follows the east side of the Salton Sea 

southward for about SO km, its trace is marked by offset 
4 QUATERNARY DEFORMATION 

NE SW 

METERS I I LÃ‡Ã‘Ã‘ 1 METER Ã‘Ã‘Ã‘Ã I 

m -' 

and, with pebbles 

maw 

- 

Sill or ctey, dy 

in p1,ca - Fault zone, 

mixed lii~locy 

- Sedimeotar, 

contact 

- Fault ,'\l ilFIGURE 4.19.-San Andreas fault exposed in southeast waU of a trench visible here is a few percent of net displacement, which is chiefly 

at PaUett Creek, Calif., 55 h northeast of La Angeb. Black strata sirike dip, approximately normal to trench wall, with block on right 

are peat layers, datable by "C methods, that show increasing moving toward observer. Uppennoet, unfaulted deposits postdate 

amounts of vertical displacement with depth, owing to cumulative 1857 earthquake; lowermost peat bed on southwest side of fault was 

slip from repeated earthquakes. Vertical component of displacement deposited about A.D. 800. Modified from Sieh (1978). 
- - 

Trench wall 

Pink 

EXPLANATION 

[Units exposed in trench wall] 

Colluvlum (Holocene) Lakeday (Holocene&- 8.350-2.500 yr B.P. 

Alluvium lH01orniOÃ‘A.D 1725-1805 near tec Qpr Channel fill of Pink River IHoloccnc) 

Older alluvium (Holocene) B River gravels of Lone Pine Creek <Pleistocene+ - 13,000 yr B.P, 

Fault-scarp breccia IHolocene) Â Contact-Hachures indicate ground surface or buried ground 

surface Organic-rich sandy clay; swamp deposits (Holocene+ 

-2,510-500 yr B.P., in pan contemporaneous with unit Qsh -.... Fault-Dotted where concealed 
4. QUATERNARY DEFORMATION 

Holocene features and at least two historical slip events 

(Sharp, 1982, p. 9). Farther south, in the Imperial Valley, 

evidence of surface faulting disappears, and a 40-km gap 

separates the south end of the San Andreas from the 

north end of the Imperial fault. This gap, nearby geothermal and volcanic activity, and the broad structural 

trough occupied by the Salton Sea and the Imperial 

Valley indicate a major change in deformation processes 

and a transition from the San Andreas fault system to the 

ridge-transform system of the Gulf of California. 

OTHER FAULTS 

The seismically active San Jacinto fault diverges from 

the San Andreas northwest of Cajon Canyon and follows 

a nearly straight course for 215 km, ending along the 

southwest side of the Salton Trough. On small-scale 

maps, it appears as a linear southward extension of the 

San Andreas segment bounding the San Gabriel Mountains-a simpler, more direct break than the arcuate, 

branching pattern of the nearly parallel San Andreas, 45 

km to the northeast. Geomorphic and geologic evidence 

of Quaternary strike slip (Sharp, 1967, 1972, 1981) 

resembles that for the main San Andreas, but locally 

vertical motion, southwest side up, may account for as 

much as 10 percent of the net slip. 

Sharp (1981) determined minimum slip rates for two 

active segments of the San Jacinto fault zone: the Clark 

fault near Anza Valley and the Coyote Creek fault, 80 km 

farther south. At the Anza Valley locality (lat 3334' N., 

long 116'39' W.), the base of a distinctive gravel marker 

bed, displaced at least 5.7 km from its source terrane of 

schist, gneiss, and amphibolite, overlies 30 m of sediment 

which, in turn, overlie a rhyolite ash layer dated at 730 ka 

and correlated with the air-fall ash at the base of the 

Bishop Tuff. Geologic data require a slip rate of at least 

0.8 dyr (Sharp, 1981, p. 1755-1757), and a higher rate, 

as high as 2.5 cmlyr, is possible (R.V. Sharp, in Clark and 

others, 1984). North of Anza, geologic relations during 

4 

FIGURE 4.20.-Near Cajon Canyon, 75 km east-northeast of Los 

Angeles, a displaced terrace riser (TR) between two Quaternary 

terrace levels (Qt2 and Qt3 of Lone Pine Creek, out of view in 

foreground) records slip on the San Andreas fault. Sketch, baaed on 

a planetable map (Weldon and Sieh. 1985, fig. 6). exaggerates relief 

by about 1.4x; terrace risers actually slope no more than 25'. 

Relationsused to establish Holocene history offaultingand asliprate 

of about 2.6 cmlyr are derived from trench-wall exposures, here 

generalized from Weldon and Sieh (1985, fig. 12). and from wmorphic and stratigraphic evidence of stream piracy and other fluvial 

processes. 

the past 30 ka yield evidence of average slip rates of 1.2 

to 1.7 cmlyr (Merifield and others, 1987) for this fault. 

The minimum Holocene slip rate for the past 6 ka on the 

Coyote Creek fault-derived from a 14Cdated, displaced 

channel margin (lat 33'06.4' N., long 116'02.7' W.)-is 

0.3 dyr, but amuch higherrateis possible (R.V. Sharp, 

oral commun., 1986). Upper limits for the slip rate on the 

San Jacinto fault (Prentice and others, 1986; Rockwell 

and others, 1986; Meriiield and others, 1987) thus range 

from 1.0 to 1.7 cmlyr, in dose agreement with Sharp's 

(1981) estimate of 0.8 to 1.2 cm!yr for the late Quaternary. The Elsinore fault and its northern branches, the 

Whittier and Chino faults, lie about 35 km to the 

southwest of the San Jacinto fault. Seismic activity, 

occurring chiefly as small earthquakes, increases southward, but the strongest evidence for Quaternary faulting 

is along the north half of the fault, where Kennedy (1977) 

and Ziony and Yerkes (1985, table 5) reported examples 

of offset stratigraphy and landforms. South of lat 33'30' 

N., evidence of recent movement on the fault is sparse 

and widely spaced (Clark, 1975), although faulted Holocene fan deposits within 25 km of the United StatesMexican border indicate a strike-slip rate of about 

0.4k0.1 dyr (Pinault and Rockwell, 1984). 

South of its junction with the northeast-dipping Whittier reverse fault, the northern section of the Elsinore 

fault displays consistent evidence of oblique, dominantly 

horizontal slip. The vertical component is up to the 

southwest (Jahns, 1964, p. 45-46; Kennedy, 1977; Heath 

and others, 1982), as it is along the San Jacinto fault. For 

100 km, the northern section of the Elsinore fault follows 

the northeast front of the Santa Ana and Elsinore 

Mountains, juxtaposing a thick Quaternary sedimentary 

fill on the northeast against pre-Tertiary crystalline and 

sedimentary rocks in the mountain blocks. South of Lake 

Elsinore flat 33'39' N., long 11721' W.), a facies boundary between sandstone and conglomerate, dated at about 

700 ka by a layer of the Bishop Tuff, is offset 5 km 

(Kennedy, 1977, p. 5, 9), indicating a strike-slip rate of 

0.7 dyr. North of Lake Elsinore, Millman and Rockwell 

(1985) reported at least 1.6 km of strike-slip offset since 

the middle Quaternary, equivalent to a minimum slip rate 

of about 0.2 dyr. These observations and other geologic 

and geomorphic evidence along the fault suggest that the 

component of horizontal slip on the northern section of 

the Elsinore fault during the late Quaternary has been 

somewhat less than on the San Jacinto and probably does 

not exceed 0.7 dyr. 

The Newport-Inglewood zone of deformation (Hill, 

1971, p. 2958, 40 km southwest of the Elsinore fault, is 

a narrow (1-3 km wide) belt of echelon folds and 
104 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

left-stepping fault segments, characterized by oblique 

right slip, chiefly down to the west (Harding, 1973, fig. 

5). This zone contains six major oil fields, within which 

subsurface structural relations in upper Cenozoic strata 

are well defined by drilling. Quaternary folding and 

faulting is expressed as an aligned series of anticlmal hills 

and by deformed and faulted upper Pliocene and Pleistocene deposits (Poland and others, 1959; Barrows, 

1974). The overall structural trend is commonly mapped 

as the Newport-Inglewood fault (fig. 4.17). but individual 

segments are less continuous and less closely aligned 

than along other major strike-slip faults in the San 

Andreas system. Moreover, many anticlines along the 

zone retain their original ellipsoidal form even where 

they are cut by strike-slip faults (Hading, 1973, figs. 2, 

4, 7, 9-15). 

The onshore, northern part of the Newport-Inglewood 

fault zone extends 65 km from the Santa Monica Mountains Oat 34-03' N., long 118%' W.) at the southern 

margin of the Transverse Ranges southeastward to 

Newport Beach. Farther southeast, the offshore part of 

the zone continues for about 70 km, defined by folding 

and faulting in Tertiary and Quaternary deposits (Clarke 

and others, 1986, p. 359) beneath the sea floor. Whether 

this zone of deformation is continuous with aligned but 

more distant offshore faults, such as the Rose Canyon 

fault near San Diego, is unresolved; if the zone is 

continuous, its overall length is at least 240 km. 

Aligned swales and discontinuous eroded scarps, a few 

meters to as much as 80 m high, delineate individual fault 

segments; many of these scarps cut folded Pleistocene 

deposits. At the Inglewood and Long Beach oil fields on 

the northern, onshore part of the fault, folding and 

faulting are nearly as great in upper Pliocene strata as in 

older rocks (California Division of Oil and Gas, 1961, p. 

677, 595), indicating that most of the deformation is 

Quaternary in age. Historical faulting is evident from the 

1933 Long Beach earthquake (M=6.3) on an offshore 

segment of the zone, and from faulted artificial fill near 

Newport Beach (lat 3322' N., long 11733' W.) (Guptill 

and Heath, 1981). The horizontal slip in upper Pliocene 

strata, estimated from offset fold axes and from structure-contour maps of producing zones in faulted antidines (Poland and others, 1959, p. 75; California Division 

of Oil and Gas, 1961), is about 1.5 km, equivalent to a slip 

rate of less than 0.1 cmtyr. Vertical components of slip 

are about a fifth as large, and vertical separation of 

faulted basal Pleistocene deposits rarely exceeds 60 m 

(Yerkes and others, 1965, p. A48). 

West of the Newport-Inglewood fault zone, northwesttrending faults that exhibit evidence of Quaternary slip 

(Clarke and others, 1985, fig. 185) include the Palos 

Verdes Hills fault, and the offshore San Pedro Basin and 

Santa Cruz-Catalina Ridge faults. The direction and 

amount of Quaternary displacement on these faults are ill 

defined, but they probably have strike-slip components 

smaller than those on the San Jacinto and San Andreas 

faults. 

Many northwest-trending faults end near the 34th 

parallel against a major structural boundary that extends 

westward 250 km from the San Jaeinto fault to Santa 

Cruz Island. This boundary separates the east-westtrending structures of the Transverse Ranges, Santa 

Barbara Basin, and northern Channel Islands (lat 34' N., 

long 120Â W.) from northwest-trending structures to the 

south; it also separates higher land and sea floor to the 

north from lower to the south. A northdipping zone of 

aligned or echelon faults (Yerkes and Lee, 1987; Ziony 

and Jones, 1988), here named the "South Frontal fault 

zone" (of the Transverse Ranges block), follows the 

structural boundary and includes, from west to east, the 

Santa Cruz Island, Anacapa, Santa Monica, Hollywood, 

Raymond, Sierra Madre, and Cucamonga faults. All of 

these faults are down to the south, all exhibit evidence of 

Quaternary activity, and many have been active during 

the Holocene. The San Fernando fault, an element of the 

Sierra Madre trend, was the source of the damaging 

M-6.6 San Fernando earthquake in 1971. 

Slip rates on the South Frontal fault zone are probably 

lower than 0.5 dyr but are difficult to measure because 

most of the best exposed range-front faults consist of 

several branches. Typically, the youngest branches are 

farthest from the range front (Bull and others, 1979; 

McFadden and others, 1982) and cut tectonically derived 

fan deposits. 

The vertical component of slip on the Cucamonga fault, 

about 0.3 dyr (Matti and others, 1982), is derived from 

a progressive increase in scarp heights-all close multiples of 2 m-in successively older alluvial-fan deposits. 

These deposits range in age from middle Pleistocene 

(approx. 700 ka) to late Holocene. The ages are based on 

stratigraphic relations, the maturity of soils developed on 

the fan surface (McFadden and others, 19829, and correlation of these soils with others (Bull and others, 1979) on 

ancient stream terraces in the San Gabriel Mountains. 

FOLDING AND UPLIFT 

The southern section of the San Andreas fault system 

resembles an inclined block (fig. 4.21), with its raised 

northeast edge along the San Andreas fault and its 

downtilted southwest half submerged beneath the Pacific 

Ocean. Near Los Angeles, topographic relief across this 

block exceeds 4,300 m: from Mount San Antonio (3,067 m) 

in the San Gabriel Mountains to the floor of the Catalina 

Basin (-1,300 in). This sloping and apparently tilted 

surface actually incorporates the results of strike-slip and 

reverse or thrust faulting, and regional uplift. 
4. QUATERNARY DEFORMATION 106 

Marine Quaternary deposits of the Los Angeles Basin, 

Venture Basin, and adjoining offshore areas are locally as 

thick as 1,200 m (Yerkes, 1972). and lower Pleistocene 

beds, folded along northwestward trends, exhibit structural relief of more than 1 km in some parts of the Los 

Angeles Basin. Upper Pleistocene strata are deformed 

nearly as much, and evidence from drill holes and surface 

geologic mapping (Durham and Yerkes, 1964; Yerkes and 

others, 1966; Yerkes, 1972; Shoellhamer and others, 

1981) show that much of the folding is of late Quaternary 

age. Offshore Pliocene and Quaternary basins and folds 

also trend northwest. These basins, which border the 

coast as far south as San Diego (Howell and others, 1978), 

contain a thinner section of Upper Cenozoic strata than 

the Los Angeles Basin. 

Some of the Quaternary folding, like that along the 

Newpart-Inglewood fault zone, is probably directly related to strike slip on deeper wrench faults (Wilcox and 

others, 1973). Other, more westerly striking folds may 

reflect a compressional component of deformation similar 

18' 

1 

i 

FIGURE 4.21.-Generalized topographic map of southern California, 

showing major faults with Quaternary activity in the San Andreas 

firnit system. Faults dotted where concealed by water: hachures on contours indicate area of closed low. Compare with figures 4.9 and 

416. 
106 THE SAN ANDREAS FAL 

to that shown by folds in the central and southern Coast 

Ranges. Fold patterns in the Los Angeles Basin are 

complicated by unknown amounts of differential compaction over a faulted and folded basement surface. This 

surface, nearly 10,000 m deep at the center of the basin 

(Yerkes and others, 1965, fig. 2), underlies a nearly 

complete section of marine Neogene strata, of which the 

near-surface Quaternary deposits make up a small part. 

The basin fill thus records almost continuous subsidence 

and marine deposition from Miocene time to the present. 

Northeast of these late Cenozoic coastal basins. ~, m- 

~ " 

talline basement rocks either crop out or lie within a few 

hundred meters of the surface. Neogene subsidence of 

the coastal marine basins contrasts sharply with sustained and evidently rapid uplift, accompanied by thrust 

faulting, of the mountainous region to the east and 

northeast. The north-dipping San Fernando fault provided an example of such range-front tectonics in February 

1971. Failure on this fault caused an M=6.6 earthquake, 

opened a 15-km-long zone of surface thrusting, and 

elevated the land north of the fault by as much as 2 m. 

Similarly, major Quaternary upwarping and faulting 

have helped shape the San Gabriel and San Jacinto 

Mountains; the San Bernardiio Mountains (just beyond 

the boundary of the fault system as defined here) also 

display evidence of Quaternary uplift and thrust faulting 

and have apparently deformed with the other ranges. 

Geologically recent uplift and tilting of the mountain 

ranges is indicated by extensive summit areas above 

2,000 m, steep southwesterly-flowing streams in narrow 

V-shaped canyons, extensive alluvial fans and fan complexes at the northeast and southwest range fronts, and 

incompletely dissected Quaternary surfaces at elevations 

near 2,000 m. The dissected upland surfaces, most 

evident in the San Bernardino (Dibblee, 1975) and San 

Gabriel Mountains, appear as accordant ridge crests and 

as remnants of originally extensive uplands, some of 

which are underlain by Quaternary fan deposits and 

alluvium. 

Other geologic evidence also implies rapid late Pleistocene and Holocene uplift. The Blackhawk and Silver 

Reef debrisavalanche deposits at the north edge of the 

San Bernardino Mountains record high-velocity (31 ds), 

voluminous (0.28x1OS m") debris flows (Woodford and 

Harriss, 1928; Shreve, 1968), which surged northward 

from the range front as recently as 17.4 ka (Stout, 1975, 

1977), leaving relict lobes of slide debris on the desert 

floor 10 km from the slide source. Smaller and slower 

debris flows and rock slides still occur from time to time 

in both the San Gabriel and San Bedino Mountains. 

Rates of late Quaternary uplift can be estimated for 

parts of the southern and central San Gabriel Mountains: 

near the mouth of Little Tujunga Canyon flat 34O17' N., 

long 118-22' W.) (Menges and others, 1979), on the North 

C SYSTEM, CALIFORNIA 

Fork of the San Gabriel River flat 34WY N., long 11731' 

W.) (McFadden and others, la), and across the Cucamonga fault (lat 34-10' N., long 117-31' W.) (Matti and 

others, 1982). Strath and fill terraces at the first two 

localities have been dated by physiographic position, "C 

methods, and soil maturity; they range in age from 

middle Pleistocene to late Holocene (approx 1,000-700 

ka). Terrace elevations above modem stream channels 

reflect both climatic and tectonic influences, but Bull and 

others (1979) and Menges and others (1979) separated 

these effects to obtain uplift rates of about 0.3 cm/yr for 

the Holocene and about 0.03 cmlyr for the late Pleistocene. Interpretation (Matti and others, 1982) of the 

displacement history across strands of the Cucamonga 

fault, as discussed above, also results in an uplift rate of 

about 0.3 cm/vr for the latest Pleistocene and Holocene at 

the south front of the range. Together, these consistent 

estimates of uplift rate sample a 78-km segment of the 

San Gabriel front. Regional uplift rates to the north and 

northwest and within the main body of the Transverse 

Ranges are probably somewhat higher; rates for the 

western Transverse Ranges, as discussed in the section 

above entitled "Central Section of the San Andreas Fault 

System," are locally as high as 1 cm/yr near Ventura. 

Elevated wave-cut platforms and marine-terrace deposits document Quaternary uplift of the southern California coast from the Transverse Ranges southward to 

San Diego. Mappable, welldated (4BO-85 ka) marine 

terraces, which correlate with major worldwide highatands of the sea, show nearly uniform coastal uplift at 

ratesthat range fromabout 0.01 to0.05cm/yr(I.ajoieand 

others, 1979). 

Similar uplift rates (approx 0.02 cdyr) can be interpreted from the elevations and estimated ages of Quaternary erosional surfaces on the Pen% structural block, 

which ties between the Elsinore and San Jacinto faults. 

~a& Quaternary drainage systems on the Paloma erosional surface are presumed to be superimposed from the 

Gavilan-Lakeview surface (Woodford and others, 1971), 

which is nearly 200 m higher and an estimated 1 Ma older 

than the Paloma surface. Persistent uplift of the Perris 

block must be geologically recent because a late Miocene 

erosional surface, overlain by basalt dated at 8.3Â±0. Ma 

(J.W. Hawkins, in Woodford and others, 1971). is at an 

altitude intermediate between the two younger surfaces. 

The nearly horizontal remnants of the late Miocene 

erosional surface and the relatively undefonned detrital 

and eruptive strata of Neogene age attest to long-term 

vertical stability of the Perris structural block. 

Together, the Quaternary geomorphic and geologic 

evidence south of the Transverse Ranges implies systemwide Pleistocene and Holocene uplift, but at higher rates 

toward the northeast than on the Pems and coastal 

blocks. 
4. QUATERNARY DEFORMATION 107 

SUMMARY 

Right-lateral strike slip on fewer than 10 major faults 

dominates the Quaternary record of deformation within 

the San Andreas fault system. Of these faults, the San 

Andreas is the longest and most active, but the fault 

system includes other structural elements, such as shorter or discontinuous strike-slip faults, reverse and thrust 

faults, regional fold systems, minor folds genetically 

related to strike slip along wrench faults, and tilted, 

uplifted, or depressed crustal blocks. All of these structures result from relative, predominantly horizontal 

motion at the boundary separating the Pacific and North 

American plates. 

Though chiefly strike slip, the detailed pattern of 

Quaternary deformation differs from place to place within the fault system. Although the ultimate causes of these 

different structural patterns are ill defined and the 

subjects of continuing research, some geologic controls 

are evident: (1) the lithology and structure of basement 

rocks (see chap. 3), (2) the thickness of upper Mesozoic 

and Cenozoic strata overlying geologic basement, (3) the 

geometric orientation of major strike-slip faults, especially the San Andreas, and (4) the slip rates along 

structures controlling deformation. 

Plutonic basement includes the granitic and associated 

metamorphic rocks of the Salinian block west of the San 

Andreas fault in the central Coast Ranges, similar rocks 

of the southern California batholith in the Peninsular 

Ranges, and a more complex suite of metamorphic and 

crystalline rocks in the central and eastern Transverse 

Ranges. Elsewhere in the fault system, the dominant 

basement-rock unit is the intensely sheared and broken 

Franciscan complex, which locally contains ophiolite 

bodies and remobilized, cold intrusive masses of serpentinite. The simplest patterns of Quaternary deformation 

are within regions of exposed or near-surface plutonic 

basement, where block-bounding strike-slip faults and 

block uplift or tilting prevail. Deformation is greater and 

more complex in regions of exposed or near-surface 

Franciscan basement, where broad anticlimoria. local 

basins, and distributed shearing are more evident. It is 

most intense and complex where basement rocks are 

covered by a kilometer or more of stratified upper 

Mesozoic or Cenozoic rocks. 

Control of Quaternary deformation processes by fault 

geometry, or the interdependence of process and geometry, is best shown at the Big Bend in the San Andreas 

fault near the north boundary of the Transverse Ranges. 

There, west-northwest-trending folds and northwesttrending strike-slip faults of the northern and central 

sections of the fault system abut against a compressional 

domain characterized by east-west-trending folds, active 

thrust and reverse faults, and accelerated rates of 

vertical uplift. On much smaller scales, similar changes in 

structural trend and pattern appear at other changes in 

strike of the San Andreas or major branch faults. 

The effect of different slip rates is demonstrated by the 

contrast in structural style along the San Andreas fault in 

xntral California and along the Newport-Inglewood fault 

zone in southern California. At a slip rate of 3.5 cmlyr, 

the San Andreas has displaced middle Pleistocene (1 Ma) 

mtcrop belts and major structures by about 35 km, 

affectively creating-on opposite sides of the fault-two 

independent structural domains, each of which responds 

iifferently to its new structural setting. In the same 

period of time, the much lower slip rates on the Newport- 

[nglewood fault zone have separated Quaternary struc- 

;ures by less than 500 m, leaving the initial structural 

pattern at least partly intact and connected. 

Our current knowledge of Quaternary deformation 

within the San Andreas fault system can be expressed in 

arms of estimated rates for deformation processes (fig. 

$2). even though reliable auantitative measurements of 

slip,folding, and uplift arefew. Other lines of evidence 

from geologic mapping and geophysical investigations 

permit reasonable inferences regarding regional rates of 

teformation and the history of the fault system. Al- 

:hough some of these interpretations may change as new 

iata are acquired, several major characteristics of the 

hult system and its Quaternary history are documented 

~y the available evidence: 

1. Quaternary tectonism within the San Andreas fault 

system has deformed an area of 160,000 km2, extending from Punta Gloria to the Salton Sea, a system 

length of about 1,100 km, and from the Great Valley 

and the Mojave Desert to fault zones off shore in the 

Pacific Ocean, an average system width of about 145 

km. 

2. Throughout the fault system, the rate of right-lateral 

strike slip on northwesterly-trending faults typically 

exceeds the geologically determined rates for other 

deformation processes by an order of magnitude. 

Despite some important exceptions, horizontal slip 

approximately parallel to the plate boundary dominates the Quaternary history of the fault system. 

3. In thecentral section ofthe fault system, between lat 

35'46' and 3630' N.. most of this slip has followed the 

main San Andreas fault: slio rates on this section of 

the fault are the highest observed to date within the 

system, averaging about 3.5 cmlyr. 

4. Strike slip is distributed more broadly in both the 

northern and southern sections of the fault system, 

where the main San Andreas splits into several active 

branches; slip rates on some of these branches may 

equal or exceed the rate on nearby parts of the main 

fault. 
THE SAN ANDREA8 FAULT SYSTEM. CALIFORNIA 

5. Observed rates of Quaternary faulting, summed 

across the fault system, are insufficient to account for 

all of the relative movement (6.1 cm/yr) attributed to 

the Pacific-North American plate boundary, but 

neither the magnitude of the difference nor its cause 

is well constrained. 

6. East-westerly- to northwesterly-trending reverse 

and thrust faults occur near many strike-slip faults 

and elsewhere in the system; they are longer, more 

active, and best defined in and near the Transverse 

Ranges (between lat 33'45' and 35W N.), where 

they accompany rapid Quaternary uplift and intense 

folding on east-west-trending axes. 

7. Major west-northwest-trending fold belts deform 

Quaternary and older strata overlying deeply buried 

(1-5 km deep) basement rocks; most regional fold 

belts strike obliquely to strike-slip faults and are 

consistent with wrench-fault structural patterns, but 

the age and intensity of folding differ from place to 

place. 8. Several fault-bounded basement blocks are intemally little deformed: the Sebastopol block in northern 

California (Franciscan basement between the San 

Andreas and Rdgers Creek faults), the Salinian 

block in central California (granitic basement between the Rinconada and San Andreas faults), and 

the Perris block in southern California (granitic 

basement between the Elsinore and San Jacinto 

faults). 

9. Throughout the Quaternary, widespread uplift has 

characterized most of the fault system; locally downwarped, faulfrtontrolled basins that reverse this 

general pattern are the San Francisco Bay and Santa 

Clara Valley (northern California), the Santa Clara 

River valley-Ventura Basin (southern California), 

the Los Angeles Basin, and basins near the southern 

and southwestern margin of the San Joaquin Valley. 

10. Maximum observed uplift rates (0.5-1.0 on/yr) for 

the late Quaternary occur near Punta Go& and in 

the Transverse Ranges; elsewhere, measured uplift 

rates average about 0.05 cmtyr, but in large areas of 

the fault system no reliable measurements of uplift 

rates are available. 

11. Major strike-slip faults bound the west edge of 

several uplifted blocks in northern California: the 

SantaRosa block, bounded by the Rdgers Creek and 

Maacarna faults; the Diablo Range and Temblor 

Range blocks, bounded by the Calaveras and San 

Andreas faults; and the Santa Cruz Mountains and 

Santa Lucia Range blocks, bounded by the San 

Gregorio, San Simeon, and Hosgri faults. 

12. In the Transverse and Peninsular Ranges of southSacramento . 

0 100 200 300 KILOMETERS I , 1 

FIGURE 422.-Sketch m~ of western and southwestern California. showine selected mior faults exhibiting evidence of Quaternary activity in 

the~an~ndreaflfaultÃˆ~tem AverageQuaternary sliprate*~~bawd~measured&uesfrompublishedsourcesand,~here~& 

are incomplete or unavailable. on interpret^ plogic and geomorphic evidence. Slip components on northwest-trending faults are 

predonmiantJy horizontal, right lateral, and consistent with North American-Padfie phte motion. Group of e.^-wl~din~5ults between 
4. QUATERNARY DEFORMATION 109 

ern California, Quaternary uplift near 6trike-slip 

faults has prcduced two nearly orthogonal mountain 

chains: (1) the west-northwesterly-trending ranges 

that separate the Mojave Desert from mstal lowlands and that lie along the trend of Tmnsverse 

Ranges thrust faults and the southern branch of the 

San Andreas fault, and (2) the north-south-trending 

mountain system composed of diswte, northwesttrending ranges that are separated and possibly 

o&et by the Elsinore and San Jacinto faults west of 

the Salton &a. 

13. Rates of Quabq faulting, folding, and upm 

resemble those for the pre-Quaternary fault system; 

signiiicantly higher rates, evident near Punta Gorda 

and in the Transverse Ranges, signify late pulses of 

awelemted tecto~c activity that are still underway. 

Despite many still-unsolved problems and unanswered 

questions, the San Andreas fault system has proved a 

productive laboratory in which to pumue three task to 

quantitatively evaluate earthquake hazards in the den~ely ppdated, high-risk regions of Californkq to develop 

new investigative and analytical methods that can be 

applied in other regions of high earthquake risk and to 

better understand how geologic processes deform the 

Eds crust at plate margins. None of these tasks is 

fi~~hed, butthe progress so far has improved our 

howledge of crustal-defonnation processes and aided 

our efforts to reduce the hazards from damaging earthquakes. 

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4 QUATERNARY DEFORMATION 113 

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691. 
'I 

0 5 10 15 20 KILOMETERS @I 

I I 
T ens of thousands of small earthquakes occur in Califamia 

each year, reflecting brittle deformation of the margins of the 

Pacific and North American plates as they grind inexorably past 

one another along the San Andreas fault system. The deformational 

patterns revealed by this ongoing earthquake activity provide a 

wealth of information on the tectonic processes along this major 

transform boundary that, every few hundred years, culminate in 

rupture of the Sun Andreas fault in a great (?.-8) earthquake. 

5. SEISMICITY, 1980-86 

By DAVID P. HILL, JERRY P. EATON, and LUCILE M. JONES 

CONTENTS 

4 FIGURE 5.1. -Seiemicity from 1969 to 1980 superimposed on color Landsat image for the central Coast Rangen between the San Francisco Bay 

and Monterey Bay. Size of earthquake epicenters (yellow circles) increases with earthquake magnitude from 1.5 to6.0. Compiled by Jean Olsen. 

INTRODUCTION 

Earthquake activity along the San Andreas fault 

system (fig. 5.1) reflects brittle accommodation of the 

crust to the relative motion dong the dextral transform 

boundary between the Pacific and North American plates 

(see chap. 3). Great (M-8) earthquakes along the main 

115 

THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA; US. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 

branch of the San Andreas fault accommodate most of 

this relative plate motion. These plate-boundary earthquakes rupture the entire 15- to 20-km thickness of the 

brittle crust with right-lateral offsets of as much as 10 m 

extending for several hundred kilometers along the fault 

trace, and they recur along a given section of the San 

Andreas fault at intervals of several hundred years (Sieh, 

1981; Sieh and Jahns, 1984). The 1857 Fort Tejon 
116 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

earthquake in southern California and the 1906 San 

Francisco earthquake in northern California are only the 

two most recent such great events along the San Andreas 

fault (see chap. 6). 

In this chapter, we focus on the smaller, more frequent 

earthquakes that dominate the seismicity within the San 

Andreas fault system between recurrences of great, 

plate-boundary events. Although this interevent seismicity contributes only marginally to relative plate motion, 

it is symptomatic of processes underlying the earthquake 

cycle. In particular, the wealth of data generated by 

frequent smaller events provides important clues to the 

seismotectonic fabric, kinematics, and state of stress 

within the brittle crust and, ultimately, to the seismogenic processes common to earthquakes of all sizes within 

the San Andreas fault system. This persistent, interevent seismicity also captures widespread popular interest because it includes most of the felt earthquakes in 

California (earthquakes of Ma3 can be locally felt), and 

the larger of these interevent earthquakes (M=6-7) can 

cause extensive damage and loss of life when they strike 

near major population centers. 

We examine here the detailed patterns of earthquake 

occurrence along the San Andreas fault system recorded 

by the combined northern and southern California seismograph networks for the 7-yr interval 1980-86. These 

networks, which had evolved to much their present 

configuration (fig. 5.2) by mid-1979, enable uniform 

detection and location of all earthquakes of Mal.5 

throughout the San Andreas fault system and of Ma2 

throughout most of California. The telemetered seismic 

stations in the combined networks approach 550 in 

number. Signals from the 300 central and northern 

California stations are recorded and processed at the 

U.S. Geological Survey's (USGS) offices in Men10 Park; 

signals from the 260 southern California stations are 

recorded and processed under a joint USGS-California 

Institute of Technology (Caltech) effort on the Caltech 

campus in Pasadena. These dense telemetered networks 

overlie regional seismic networks operated by the University of California, Berkeley, and Caltech that provide 

records of Ma3 earthquakes in northern and southern 

California, respectively, from the early 1930's onward 

(table 5.1; Hileman and others, 1973; Bolt and Miller, 

1975). 

After a brief overview of the San Andreas fault system 

in the context of a broad transform boundary, we focus on 

the three-dimensional distribution of earthquakes along 

the San Andreas fault systemitself on the basis of aseries 

of detailed seismicity maps and cross sections for the 

years 1980-86. We then review selected focal mechanisms for the larger of these earthquakes as a guide to the 

kinematics of seismogenic deformation along the fault 

system. Finally, we discuss the implications of these 

seismicity patterns in terms of current tectonic processes 

dong the transform boundary. 

The seismicity maps and cross sections in this chapter, 

ffhich form the core of our presentation, are largely 

self-explanatory. The following points, however, deserve 

special emphasis: 

1. The reliability of hypocentral locations correlates 

closely with the local density of the seismograph 

network (fig. 5.2). Relative epicentral locations are 

better than k0.5 km for earthquakes within the 

densest sections of the network (corresponding focal 

depths are better than kl.O km). Relative locations 

may be uncertain by several kilometers or more, 

however, for earthquakes occurring beyond the margins of the network. 

2. All cross sections have a 2x vertical exaggeration as 

a means of illustrating patterns in the depth distribution of earthquakes beneath profiles that are many 

times longer than the limited range of focal depths 

(less than 15 km along most of the fault system). 

3. Hypocentral locations are plotted using small circles 

that scale only weakly with magnitude, to better 

emphasize detailed spatial structures within the seismicity patterns. 

4. The locations of the most commonly used place names 

and faults in this chapter are shown in figure 5.3 (see 

front of book for a more complete map). 

REGIONAL SEISMICITY AND 

THE SAN ANDREAS TRANSFORM BOUNDARY 

Dickinson (19811, among others, emphasized that the 

San Andreas fault system and the San Andreas trans- 

'om boundary are not strictly equivalent structures. The 

atter is more general, incorporating, for example, the 

;oncept of temporal evolution of Pacific-North American 

slate interaction and the recognition that the faults 

tccommodating most of the plate motion have changed 

)ver time. In this section, we emphasize that, although 

pat earthquakes along the San Andreas fault system 

:urrently accommodate most of the relative plate motion, 

slate interactions along the transform boundary influence 

iefonnation of the brittle crust over a much broader 

&on. 

The breadth of the seismicity pattern in California and 

western Nevada (fig. 5.4) suggests the lateral extent of 

ieformation associated with the San Andreas transform 

soundary. Indeed, it corresponds closely to the zone of 

distributed shear between those plates as interpreted by 

Ward (1988) from more than 5 years of very long baseline 

nterferometry (VLBI) observations at 20 Western Unitid States stations from 1982 through 1987. Figure 5.4 

ncludes all Mzl.5 events recorded by the telemetered 
5. SEISMICITY, 198(^86 117 

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

multicomponent station. 

seismic networks in figure 5.2 during the 7-yr interval 

1980-86, as well as events recorded by adjacent telemetered networks in Nevada (see Rogers and others, in 

press). Although details within this seismicity pattern 

fluctmk bm yew to yew, the broad =W of the 

pattern have remained stable for the entire historical 

TABLE 6.1.-Numbnr of ahart-psriod 8eÃ‘Ãˆ atations in CaS/mia 

Year 1940 1965 1971 1975 1981 

Ã£d.,or, <lo <20 -100 -200 -800 

%,,~- c&f,, . <lo <20 -80 -lw -m 
118 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

record of earthquake occurrence in California (see chap. In outline, the seismicity pattern for California and 

6; Ellsworth and others, 1981; Hill and others, in press; western Nevada forms a hollow ellipse with its long axis 

Hutton and others, in press). nearly coincident with the transform boundary. This 

FIGURE 5.3.-Place names and faults most commonly used in text (see 

front of book for more complete maps of place names and faults). 

Faults (dotted where concealed): B, Banning; BP, Big Pine; BS, 

Bartlett Springs; BZ, Brawley seismic zone; C, Calaverav; CN, 

Concord; CU, Cucamonga; DV, Death Valley; E, Elsinore; FC, 

Furnace Creek; G, Garlock; GV, Green Valley; H, Haywad; HG, 

Hoseri; HRC, Healdsburg-Rodger6 Creek, IM, Imperial- LVC, Long 

Valley caldera; M, Maacama; MC, Mission Creek; N1, NewportInglewood; OT, Otiiita; PM, Pinto Mountain; PV, Pake Verdes; 

PVA, Panamint Valley; R, Rinconada; RC, Rose Canyon; SA, San 

Andreas, SC, San Clemente Island; SG, San Greprio; SJ, San 

Jacinto; SN, Sierra Nevada, SNA, Sur-Nadmiento; W, Whittie7 

WM, White Mountains; WW, White Wolf. Arrows and numbera 

indicate direction and amount of motion, resoectively. . of Pacific and . 

Go& plates with respect to North ~m'eri&n plate to the east; red 

lines indicate l&S7 and 1906 ruptures of San Andreas fault. 
5. SEISMICITY, 1980-86 

pattern is widest across central California, where it 

approaches nearly a third of the 1,100-km length of the 

transform boundary, from the Mendocino triple junction 

in the north to the head of the Gulf of California at the 

south. Extended alignments of epicenters within this 

pattern suggest a coarse structural fabric linking the 

broad distribution of earthquakes to the transform 

boundary. Seismicity along the San Andreas fault system 

itself stands out as a series of subparallel, northwesttrending lineations extending the length of coastal California. The alignment of epicenter clusters along the east 

side of the Sierra Nevada in eastern California branches 

northward from the south end of the San Andreas fault 

system in the Salton Trough only to bend back toward the 

north terminus of the San Andreas fault system at the 

Mendocino triple junction in northern California. The 

FIGURE 5.4.-Locationsof64,000 Msl.5 earthquakesin Californiaand westernNevadaduring 1980-86andmapped Holocenefaults(dottedwhere 

concealed; major branches of the San Andreas fault system marked in red). 
120 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

Sierra Nevada-Great Valley and western Mojave Desert 

blocks form a broad quiescent region between the San 

Andreas and eastern California seismicity bands. In 

southern California, pronounced transverse seismicity 

belts coincident with the southern margin of the Sierra 

Nevada and Transverse Ranges, respectively, span this 

otherwisequiescent region. A weaker, more diffuse 

seismicity belt near lat 37Â N. spans the Sierra NevadaGreat Valley block in central California, forming a visual, 

if not structural, link between the San Andreas fault 

system and the dense cluster of epicenters in eastern 

California. This major cluster in the eastern Sierra 

Nevada represents an episode of intense earthquake 

activity in Long Valley caldera and vicinity that began in 

1978 and has persisted to the present (Van Wormer and 

Ryall, 198ft Hill and others, 1985b). 

The displacement pattern associated with earthquakes 

throughout California and western Nevada is broadly 

consistent with deformation under a coherent, regional 

stress field dominated by plate-boundary interaction 

along a northwest-striking, dextral transform boundary 

(Hill, 1982). Strike-slip focal mechanisms with rightlateral slip on northwest- to north-northwest-striking 

planes, for example, are common through most of the 

region. Regional variations within this pattern include a 

tendency toward normal slip on northerly-striking planes 

along the western margin of the extensional Basin and 

Range province in eastern California, and toward reverse 

slip on easterly-striking planes in the Transverse Ranges 

of southern California. Compressional deformation perpendicular to the San Andreas fault within the Coast 

Ranges, however, represents an important deviation 

from this regional pattern. 

SEISMICITY ALONG 

THE SAN ANDREAS FAULT SYSTEM 

Sections of the San Andreas fault system stand out on 

seismicity maps as a network of northward-branching 

alignments of epicenters through the central and northern Coast Ranges and as subparallel lineations of clusters 

of epicenters that branch northward from the south end 

of the Imperial fault toward the Transverse Ranges in 

southern California (fig. 5.4). One of the most remarkable 

aspects of the seismicity pattern associated with the fault 

system, however, is the nearly complete absence of 

earthquake activity down to even the smallest magnitudes (M-1.5) along those sections of the fault that have 

ruptured with the largest historical earthquakes, the 

great (Ma) 1857 and 1906 earthquakes (see figs. 5.3, 

5.4). The southernmost section of the San Andreas fault, 

from Indio to the Salton Sea, also lacks microseismicity, 

although no large earthquake has ruptured this section in 

the past 200 yr. These quiescent Clocked") segments of 

the fault stand in sharp contrast to the segments marked 

by persistent linear concentrations of small to moderate 

earthquakes. 

This dual expression of the fault system evident on 

current seismicity maps apparently reflects fundamental 

differences in the long-term behavior of the respective 

segments. In particular, seismic activity along the 

'locked" segments of the main trace of the San Andreas 

fault may be limited to the recurrence of major earthquakes at intervals of 100 to 300 yr accompanied by 

immediate foreshock and aftershock sequences, and 

these segments may remain quiescent for most of the 

interevent time associated with the cycle between great 

earthquakes (Ellsworth and others, 1981). In contrast, 

those segments with persistent microearthquake activity 

probably seldom, if ever, rupture with great earthquakes, although they may be capable of generating 

earthquakes as large as M4. 

Aseismic creep also characterizes and is largely confined to those fault segments along the San Andreas fault 

system that show persistent microearthquake activity 

(Wesson and others, 1973; Schuiz and others, 1982). 

Creep is most pronounced along the central California 

segments of the fault system, where average creep rates 

match the long-term displacement rates of 32 to 34 

mmlyr. Louie and others (1985) documented creep along 

sections of the seismically active fault segments in the 

Salton Trough, and Astiz and Alien (1983) documented 

creep along a section of the Garlock fault that is marked 

by microearthquake activity. The creep rates in these 

two areas, however, are more than an order of magnitude 

less than the long-term deformation rates. 

In the following subsections, we consider the 1980-86 

seismicity along and adjacent to the major sections of the 

San Andreas fault system in more detail. We begin with 

the Mendocino triple junction in the north and move 

southward, generalizing slightly Allen's (1968) subdivision of the fault system into four major sections of 

contrasting seismic behavior: (1) the quiescent 1906 

break and subparallel branches, (2) branches forming the 

central California active (creeping) section, (3) the quiescent 1857 break, and (4) branches forming the southern 

California active section south of the Transverse Ranges. 

MENDOCINO TRIPLE JUNCTION 

The three lithospheric plates that dominate the modern 

tectonics of coastal California meet at the Mendocino 

triple junction, which is marked by a dense cluster of 

epicenters just off Cape Mendocino (fig. 5.5A; see chap. 

3). North of this triple junction, oblique subduction 

dominates, with the eastern margin of the Gorda plate 
6. SEISMICITY, 1980-86 

(the southernmost section of the Juan de Fuca plate) 

slipping beneath the North American plate north-northeastward at arate of 30 to 50 rnmlyr (Wilson, 1986). South 

of the triple junction, the Pacific plate is moving past the 

North American plate along the San Andreas transform 

boundary on a heading of 35"'s W. of N. at a rate of 

approximately 60 mmlyr (DeMets and others, 1987). 

The dense lineation of epicenters that trends westnorthwest from Cape Mendodno corresponds closely to 

the Mendocino Fracture Zone (MFZ) and the PacificGorda plate boundary but follows a slightly more northerly trend. Details of bow the San Andreas fault ties into 

the triple junction, however, are unclear. The trace of the 

San Andreas fault lies off shore north of Point Arena, and 

the broad area of seismic quiescence south of the triple 

junction offers few clues to the kinematics of this 

northernmost segment of the fault. 

The conspicuous linear zone of epicenters marking a 

northeast-trending slice across the southwest corner of 

the Gorda plate is the aftershock zone of the M=7.2 

Eureka earthquake of November 8, 1980. This was the 

largest earthquake to occur in California during the 

interval 1980-86. The N. 53' E. trend of its aftershock 

zone agrees closely with the strike of the left-lateral slip 

plane inferred from the focal mechanism of the main 

shock, which was located about a fourth of the way 

downstrike from the northeastern end of the aftershock 

zone (Eaton, 1989). The aftershocks from this earthquake 

FIGURE 5.5.-Seismicity in the Mendodno triple junction area, north- ment of cross section. B, Depth sections perpendicular to N. 75" W, ern California. A, Earthquake locations, showing major branches of trendof the Mendorino Fracture Zone (A-A'andB-B' outlined in fig. 

the San Andreas fault system in red: faults dotted where concealed. 6.54. C. Depth sections parallel to N. 75" W. trend of the Mendodno 

Magnitude symbols shown in explanation are seated with enlarge- Fracture Zone (C-C' and D-D' outlined in fig. 5.5A). 
122 THE SAM ANDREAS FAULT SYSTEM, CALIFORNIA 

died off within a few months after the main shock-a 

notably short aftershock sequence for an earthquake of 

this size. 

A third, more diffuse group of earthquakes in the 

vicinity of the triple junction shows little tendency to 

concentrate in linear zones. These epicenters form an 

irregular zone with the greatest concentration in the 

vicinity of Cape Mendocino, gradually dying off with 

distance to the north, east, and west. To the south, the 

seismicity dies off abruptly across the MFZ and its 

landward extension. 

The three-dimensional aspects of this triple junction 

seismicity are illustrated by four cross sections (figs. 

5.5B, 5.50. The two cross sections perpendicular to the 

N. 75O W. trend of the MFZ (fig. 5.5B compare the 

distribution of focal depths within the submarine section 

of the Gorda plate adjacent to the triple junction (A-A') 

with that within the adjacent, subducted section of the 

Gorda plate and the overlying North American plate 

(B-B'). Earthquake hypocenters along the southern 

margin of the submarine Gloria plate define a dense, 

vertically elongate zone beneath the MFZ that dips 

70'-75Â N. and extends to depths of nearly 35 km. 

Earthquakes north of this zone cluster in a somewhat less 

dense, wedge-shaped core outlined by a northward 

shallowing of maximum focal depths accompanied by a 

northward deepening of minimum focal depths that 

converge at a depth of about 20 km (Eaton, 1989). 

Overlying this relatively dense wedge is a more diffuse 

distribution of hypocenters, the northernmost of which 

represent the 1980 M=7.2 event and adjacent aftershocks. The absence of seismicity south of the MFZ 

indicates that the northeast comer of the Pacific plate 

behaves as a rigid block in its interaction with the 

younger, thinner, and internally deforming Go* plate 

(Wilson, 1986; Eaton, 1989). 

In the onshore cross section (B-B', fig. 5.5ffl, the 

intense seismicity associated with the MFZ loses most of 

MENWCINO 

FRACTURE 

ZONE 

Eureka 

. . . . I A' 

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DISTANCE. IN KILOMETERS 

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5. SEISMICITY, 1980-86 

its expression. Here, the southern margin of the subducted Gorda plate is marked only by isolated clusters of 

hypocenters at depths of 10 and 25 km. To the north, 

however, a band of hypocenters concentrated at about a 

depth of 20 km corresponds closely to the base of the 

wedge-shaped distribution beneath the submarine section of the Gorda plate, including the downward flexing of 

this band as it approaches the southern margin of the 

Gorda plate. South of the landward extension of the 

MFZ, the seismicity shallows rather abruptly, reflecting 

the edge of the subducted Gorda plate beneath the North 

American plate and the rather thin seismogenic crust 

associated with the San Andreas fault system to the 

south. 

The cross section parallel to and including the MFZ and 

its landward extension (D-D', fig. 5.50 reveals that the 

dense seismicity cluster along the MFZ near the triple 

junction tapers westward along the MFZ with a wedgelike geometry to a 20-kmdeep band of hypocenters, 

much the same as it does to the north (cross see. A-A', 

fig. 5.5B). (The pronounced linear concentration of hypocenters at 15-km-depth beneath the submarine Gorda 

plate in cross sections A-A' and D-D' represents the 

default focal depth for the more poorly located earthquakes beyond the perimeter of the onshore seismic 

network.) The landward extension of the MFZ shows up 

only weakly as a diffuse scattering of hypocenters extending to a small, isolated cluster of 25-km-deep hypocenters some 50 km east of the triple junction (A= 190 km, 

cross sec. D-D') and, possibly, as far as a handful of 30- 

to 50-kmdeep hypocenters 100 km east of the triple 

junction (A=260 km, cross sec. D-D'). Focal depths of 

the shallow seismicity in the northern Coast Ranges are 

confined to the upper 10 km of the crust. Farther east, 

however, focal depths increase again to depths of 35 to 40 

km beneath the north end of the Great Valley. An 

east-dipping quiescent band, about 10 km thick, appears 

to separate the seismicity associated with the MFZ from 

that beneath the northern Coast Ranges and the Great 

Valley to the east. The geometry of this band suggests 

that it may somehow be related to the geometry of the 

subducted Gorda plate beneath the western margin of the 

North American plate. 

The parallel cross section immediately to the north 

(C-C', fig. 5.60 reinforces the impression that the 

distribution of hypocenters carries an image of subduct- 

< Klamath Mountains and > 

- Pacificocean > Coast Ranges < Cascade Range - 

C - Coast Lake Shasta LassenPwak 

LEVEL 

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D - Pacificocean - GreatValley -D' ? SEA 

50 - 

60 - 

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30 W 90 120 150 180 210 240 270 300 330 360 

DISTANCE, IN KILOMETERS 

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124 THE SAM ANDREAS FAULT SYSTEM, CALIFORNIA 

ed-Gorda-plate geometry. Maximum focal depths increase systematically from 25 to 30 km beneath the 

submarine Gorda plate to nearly 80 km beneath the 

southern Cascade volcanoes. Although the easternmost 

of these deep earthquakes are small and few, their 

locations are well constrained (Cockerham, 1984, Walter, 

1986). As in the section of the fault to the south, a 

seismically quiescent, east-dipping band appears to separate earthquakes within the Gorda plate from those in 

the overlying North American plate. 

THE 1906 BREAK AND THE NORTHERN COAST RANGES 

Aside from a light scattering of epicenters about the 

trace of the San Andreas fault through the San Francisco 

peninsula and the Santa Cruz Mountains to the south, the 

rupture zone of the 1906 earthquake is nearly aseismic. 

This pattern has persisted not only through the period 

1980-86 shown in figure 5.6 but also at least since the 

mid-1930's. when instrumental data became available for 

reliable earthquake locations in the area (see Bolt and 

Miller, 1975; Hill and others, in press). It was interrupted, however, by the M=7.1 Lorna Prieta earthquake on 

October 17, 1989, which ruptured the southernmost 45 

km of the 1906 break (cross sec. L-L', fig. 5.7; see chap. 

6). The cluster of epicenters along the fault just west of 

San Francisco coincides closely with Boore's (1977) estimate of the epicentral location for the 1906 main shock. 

We note that the greatest offsets along the 1906 rupture 

occurred north of the epicenter along the stretch of the 

fault that now shows the lowest seismicity (see Thatcher 

and Lisowski, 1987). 

The pair of subparallel epicenter lineations through the 

northern Coast Ranges east of the 1906 break closely 

follow the Rodgers Creek-Healdsburg-Maacama fault 

zone (west) and the Green Valley-Bartlett Springs fault 

zone (east). In both lineations, the epicenters tend to 

cluster along the eastern margins of these 2- to 3- 

km-wide fault zones, which are characterized by multiple 

strands of subparallel, curvilinear fault traces (see maps 

at front of book; Dehlinger and Bolt, 1984). These fault 

zones are essentially colinear with the Hayward and 

Calaveras faults to the south, although an aseismic 

interval coincident with the eastern arm of the San 

Francisco Bay (San Pablo and Suisun Bays) obscures the 

connection between these branches of the fault system. 

Most of the earthquakes defining the pair of subparallel 

lineations through the northern Coast Ranges are small 

(M^S), and, indeed, these fault zones were not recognized as seismically active branches of the San Andreas 

system until after the northern Coast Ranges section of 

the telemetered network (fig. 5.2) became operational in 

late 1979. 

Dense clusters of epicenters just south of Clear Lake 

define a northeast trending pattern transverse to and 

midway along these northern two branches of the San 

Andreas fault system. The southwesternmost of these 

clusters represents microearthquake activity associated 

with the Geysers geothermal field (Eberhart-Phillips and 

Oppenheimer, 1984; Oppenheimer, 1986). The cluster 

just to the northeast underlies the Clear Lake volcanic 

field, which last erupted about 10 ka (Donnelly-Nolan and 

others, 1981). Scattered clusters farther to the northeast 

suggest a tenuous link between this Geysers-Clear Lake 

trend and the north-south-trending lineation of epicenters along the axis of the north end of the Great Valley. 

In longitudinal cross section (H-H', fig. 5.6), earthquakes occurring along the two northern branches of the 

San Andreas fault system are moderate in number and 

rather evenly distributed except for the dense, shallow 

cluster beneath the Geysers-Clear Lake area. The depth 

to the base of the continuously seismogenic crust shows 

considerable relief, deepening to between 12 and 13 km 

south of the Geysers-Clear Lake area and shallowing to 

only 5 km beneath and just north of the geothermal field. 

Farther north, the base of the seismogenic crust deepens 

gradually to a maximum of about 10 km. Note the isolated 

cluster of small but well-located events at depths of 13 to 

18 km beneath Clear Lake, just north of the shallowest 

depths to the base of the continuously seismogenic crust. 

Another isolated cluster of deep earthquakes located 

beneath Suisun Bay have focal depths as great as 15 to 25 

tan. The more numerous earthquakes along branches of 

the fault system south of Suisun Bay paint in a dense 

distribution of hypocenters throughout the upper 10 to 15 

km of the crust. 

Although the two transverse cross sections across the 

northern Coast Ranges (E-E', F-F', fig. 5.8A) show the 

concentration of hypocenters around the Rodgers Creek 

and Green Valley branches of the San Andreas fault, they 

provide no hint of the location of the main branch of the 

San Andreas fault that ruptured in 1906. The southern of 

these two cross sections (F-F') includes the dense, 

shallow cluster of events associated with the GeysersClear Lake activity, as well as the cluster of deeper 

events beneath the eastern margin of the Coast Ranges. 

Focal depths in the latter cluster, which falls at the 

northeast end of the Geysers-Clear Lake lineation and at 

the south end of the north-south-trending lineation beneath the Great Valley, range from 10 to 25 km, a depth 

range common to earthquakes near the north end of this 

Great Valley seismicity lineation (cross sec. D-D', fig. 

5.50. Maximum focal depths increase by several kilometers from west to east in both cross sections. Farther 

east, they increase abruptly to depths of 25 km or so 

along the Great Valley lineation. 
5. SEISMICITY, 1980-86 

CENTRAL CREEPING SECTION 

Densely aligned epicenters mark the nearly straight, 

creeping section of the San Andreas fault that separates 

the south end of the 1906 break near San Juan Bautista 

from the north end of the 1867 break near Cholarne (fig. 

5.6B). Densely aligned epicenters follow the Calaveras 

fault northward to a point just east of the south tip of the 

San Francisco Bay, where the Hayward fault branches to 

the west and the Greenville fault zone branches to the 

east. Few epicenters fall along the northward extension 

of the Calaveras fault beyond this branching point, 

although a diffuse cluster of epicenters coincides with the 

right-stepping offset between the north end of the 

Calaveras fault and the Concord fault (fig. 5.6A). This 

dilatational offset was the site of pronounced earthquake 

swarms in June 1970 and August 1976 (Lee and others, 

1971;. Weaver and Hill, 1978179). 

Although these dense alignments of epicenters coincide 

closely with the mapped surface trace of the San Andreas 

fault system as first documented by Eaton and others 

(1970a), the coincidence is not everywhere perfect (fig. 

5.65). In the region where the Calaveras fault branches 

from the main trace of the San Andreas fault, for 

example, the densely aligned epicenters appear to be 

systematically displaced 3 to 4 km westward of the 

surface trace of the San Andreas fault and a somewhat 

smaller distance eastward of the surface trace of the 

Calaveras fault. Much of this apparent offset results from 

a strong contrast in rock type and P-wave velocity across 

the faults that is not adequately accounted for in routine 

hypocenter locations (Mayer-Rosa, 1973; Pavoni, 1973; 

Spieth, 1981). When the hypocenter locations are recalculated using a more appropriate, two-dimensional structural model, these offsets are much reduced but not 

completely eliminated. The remaining offsets reflect 

deviation of the faults from the vertical, with the San 

Andreas fault dipping 70Â W. (Pavoni, 1973, Spieth, 1981) 

and the Calaveras fault dipping 80' E. (Reasenberg and 

Ellsworth, 1982). 

Wesson and others (1973) recognized the dose correlation between active creep and persistent microseismic 

activity along the 200-km-long section of the San Andreas 

fault north of Parkfield and proposed that this correlation 

may hold for other branches of the fault in central 

California as well. Allen (1981) and Schultz and others 

(1982) pointed out that this correlation holds for the 

Calaveras-Hayward fault system, but creep measurements have yet to be made on the subparallel Rodgers 

Creek-Healdsburg-Maacama and Green Valley-Bartlett 

Springs faults. Creep is the dominant process for shallow 

slip along the central section of the San Andreas fault, 

and geodetic measurements spanning this section of the 

fault indicate that the long-term slip rate of 32 mmlyr 

along the fault accommodates nearly all of the local plate 

motion. Because it appears that little, if any, shear strain 

is accumulating in the blocks on either side of the fault, 

most seismologists believe that this section of the fault is 

unlikely to rupture in a great earthquake (see chap. 7). 

The creeping section of the fault system, however, has 

produced several moderate earthquakes during historical 

time (see chap. 6). The most recent of these events, which 

occurred along the right-branching segments northeast 

of the San Andmas fault, where creep rates average 

several millimeters per year, include (1) the M=5.9 

Coyote Lake earthquake of August 6, 1979, and the 

M=6.2 Morgan Hill earthquake of April 24,1984, which 

ruptured adjacent 20-km-long segments of the Calaveras 

fault south of its junction with the Hayward fault 

(Reasenberg and Ellsworth, 1982; Bakun and others, 

1984); and (2) the M-5.5 and 5.8 Livennore events of 

January 24 and 27, 1980, which ruptured a 20-km-long 

stretch of the Greenville fault north of Livennore (Bolt 

and others, 1981). 

The most noteworthy sequence of moderate earthquakes along the central section of the San Andreas fault 

involve the five virtually identical M=6 events that have 

ruptured the same 30-km-long stretch near Parkfield at 

nearly 22-year intervals since 1881 (Bakun and McEvilly, 

1984). This stretch of the fault is defined by a 1-km 

right-stepping offset on the south and a 6Â W. bend on the 

north; it coincides with the transition between the south 

end of the creeping section of the fault and the north end 

of the 1867 break (see Bakun and Lindh, 1985, fig. 1). The 

most recent of these characteristic Parkfield earthquakes 

occurred in 1966, and an intensive monitoring experiment 

is underway to capture a detailed instrumental record of 

the next Parkfield earthquake, which is predicted to 

occur sometime within a 10-year window centered on 

1987-88 (Bakun and Lindh, 1985). 

The scattered seismicity within the Coast Ranges 

surrounding the San Andreas fault system is distinctly 

more intense in the Franciscan terrane east of the fault 

than in the granitic Salinian block to the west. The large, 

dense cluster of epicenters along the eastern margin of 

the Coast Ranges adjacent to the south end of the 

creeping section represents the aftershocks of the M=6.7 

Coalinga earthquake of May 2, 1983, and the M=5.7 

Kettleman Hills earthquake of August 4, 1985, both of 

which involved reverse slip on northwest-striking planes 

subparallel to the adjacent section of the San Andreas 

fault (Stein and King, 1984; Eaton, 1989). Scattered 

clusters of epicenters within the Franciscan terrane show 

a crude northwest-trending alignment with the southwest edge of the Coalinga-Kettleman Hills aftershock 

zone and the Ortigalita fault, the north end of which 

passes beneath the San Luis Reservoir (LaForge and 

Lee, 1982). This weakly defined lineation is essentially 
126 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

colinear with the Greenville fault, east of the San 

Francisco Bay (fig. 5.4). 

The Salimian block west of the central section of the San 

Andreas fault and east of the Sur-Nacimiento fault forms 

a broad, nearly aseismic swath along the west flank of the 

Coast Ranges. These two seismically active fault zones 

separate the granitic Salinian block from the Franciscan 

terrane on either side (see chap. 3). The Rinconada fault, 

within the Salinian block, appears to be nearly aseismic 

except, possibly, toward the south where it approaches 

the Sur-Nacimiento fault zone. The small cluster of 

epicenters just east of the midpoint of the Rinconada fault 

represents a persistent spot of microearthquake activity 

at deoths of 8 to 10 km near the San Ardo oil field (Polev. . ", 

1988). 

In the cross section along the central section of the San 

Andreas fault (L-L', fig. 5.77, the actively creeping 

section of the fault shows up as a densely mottled 

distribution of hypocenters within the upper 12 to 15 km 

of the crust. The density of hypocenters within this 

creeping section tends to decrease with depth, and the 

denser clusters generally are concentrated at depths of 

less than 5 to 8 km. The base of the seismogenic zone 

undulates about an average depth of some 13 tan beneath 

most of the creeping section, but it deepens to 15 km 

beneath both the northern and southern transitions to 

the locked sections of the fault. In contrast to the 

creeping section of the fault, the sparse seismicity 

associated with the locked segments that ruptured in 

1906 (north) and 1867 (south) tends to be concentrated 

toward the deeper parts of the seismogenic crust. Note 

that the 1989 Loma Prieta M=7.1 earthquake ruptured 

the 45-km-long section of the San Andreas fault with a 

pronounced U-shaped gap in shallow earthquakes immediately north of the creeping section of the fault (cross 

sec. L-L', fig. 5.7; see chap. 6). 

The cross sections transverse to the central San 

Andreas fault system (G-G', fig. 5.8A; I-1', J-J', K-K', 

fig. 5.8B) reveal the seismically active, creeping branches 

of the fault as narrow, near-vertical hypocenter distributions coincident with the fault plane. The broadened 

distribution in cross sections I-/' and J-J' (fig. 5.85) 

reflects the oblique projection of earthquakes along the 

Calaveras fault zone and the clustering northeast of the 

fault in the Bear Valley region (Ellsworth, 1975), respectively. These transverse cross sections also emphasize 

the quiescence of the Salinian block relative to the 

Franciscan terrane on either side (note, however, the 

isolated cluster of deep events beneath San Ardo in the 

Salinian block in cross sec. K-K'), and the fairly uniform 

depth of 12 to 15 km to the base of the seismogenic zone 

that persists throughout the central Coast Ranges. As in 

the area farther north, however, maximum focal depths 

increase rather abruptly to 25 km beneath the eastern 

margin of the Coast Ranges and the Great Valley. This 

increase in focal depth is particularly pronounced beneath 

the dense cluster of hvnocenters associated with the 19B 

Coalinga earthquake and its aftershocks (cross sec. GG', 

fig. 5.8B). 

THE 1857 BREAK AND THE TRANSVERSE RANGES 

The 1867 rupture of the San Andreas fault began near 

Parkfield at the south end of the creeping section of the 

fault and propagated southeastward along the straight 

segment, through the Carrizo Plain and around the Big 

Bend near Tejon Pass, and thence along the relatively 

straight, east-southeastward trend of the Mojave segment to CajonPass, where the San Jacinto fault branches 

to the south (figs. 5.3, 5.9A). Right-lateral offsets assodated with this great earthquake generally decreased 

from 9 m along the Carrizo Plain segment, through 6 m 

around Fort Tejon, to 3-4 m along the Mojave segment 

(Sieh, 1978). 

The pronounced bends in the San Andreas fault at 

either end of the east-southeast-trending Mojave segment involve strong structural complexities and clusters 

of persistent seismic activity. Both bends, for example, 

spawn major left-lateral faults that form conjugate sets 

to the San Andreas system. Sykes and Seeber (1986) 

proposed that these two major bends in the San Andreas 

fault system represent large-scale asperities that exert a 

strong influence on the rupture patterns of great earthquakes along the San Andreas fault in southern Califorda. The San Andreas fault appears to maintain its 

integrity as a single, more or less continuous zone 

through the 30' bend at Fort Tejon. Convergence resulting from the pronounced counterclockwise cant of the 

Mojave segment of the San Andreas fault (N. 66' W.) 

with respect to the average N 37Â W. orientation of the 

transform boundary largely accounts for the crustal uplift 

and shortening expressed in the Transverse Ranges. 

The straight Carrizo Plain segment is, except for a 

small cluster of events near Simmler (lat 35126' N., long 

120%)' W.), almost completely aseismic, much like the 

Point Arenasegment of the 1906 break. The straight pan 

of the Moiave semnent also is seismicallv verv auiet. The " "A 

southernmost 80km of the 1857 rupture zone produces 

few earthquakes, forming a sinuous lineation around the 

fault. Again, an analogy can be drawn with the 1906 

rupture zone: The southernmost 80 to 100 km of the 1906 

rupture, which exhibited less slip than the rupture zone 

farther north, also has a slightly higher background 

seismicity rate than that to the north. The Mojave 

earthquakes are temporally clustered (Sauber and others, 1983) and, because of their reverse focal mechanisms 

and displacement from the surface trace of the San 
5. SEISMICITY, 1980-86 

EXPLANATION 

Mfl.3+ 

0 3.w 

0 4.S+ 

0 6.W 

A 

FIGURE 5.6.-Seismicity in the northern (A) and central (5) Coast Eangee. Major branches of the San Andreas fault marked in red; faults 

dotted where concealed. Earthquakes within rectangles are plotted in corresponding depth sectionsin figures 5.7 and5.8. C, Calaveras 

fault: CN, Concord fauft; G, The Geysers; SF Bay, San Francisco Bar. SIB, San Juan Bautista: SSB, Sirisun Bar. D, Dixon; V, 

Vacaville; W, Wmtera. 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

Andreas, are thought to be on secondary fault structures 

rather than on the San Andreas fault itself. 

Within the area of the Big Bend of the San Andreas 

fault near Tejon Pass, the level of seismicity is much 

higher than in the adjoiningregions. Much of this activity 

is associated with the Pleito-White Wolf fault system, 

which abuts the San Andreas fault in the northern bend, 

some 40 km west of the junction with the Garlock-Big 

Pine faults (fig. 5.9A; cross sec. N-N', fig. 5.9B; see fig. 

5.12). The White Wolf fault ruptured in 1952 with the 

M=7.7 Kern County earthquake, accompanied by leftoblique reverse slip on a southeast-clipping fault plane 

(Richter, 1968; Stein and Thatcher, 1981). This is the 

largest earthquake to occur in California since the M-8 

San Francisco earthquake. In contrast to the quiescent 

1906 rapture, however, the White Wolf fault continues to 
6. SEISMICITY, 1980-86 

be marked by persistent aftershocks of the 1952 Kern 

County earthquake. The Garlock and Big Pine faults are 

essentially quiescent within 40 to 50 km of the San 

Andreas, although two small clusters of epicenters form 

a nearly symmetrical pattern on either side of the 

junction of these sinistral faults with the San Andreas. 

Seismicity increases markedly near the south end of 

the 1867 rupture zone, where the San Jacinto fault and 

the east-west-striking Cucamonga fault, which forms the 

south front of the central Transverse Ranges, intersect 

the San Andreas fault. A bulbous lobe of epicenters 

extends westward along the Cucamonga fault from this 

junction. Scattered epicenters fill in the wedge of the 

Transverse Ranges between the Cucamonga and San 

Jacinto faults. In figure 5.9A, note the tight lineation of 

epicenters that appears to follow the trace of the Cucamonga fault westward to its intersection with the 

Elsinore fault (Chino branch). The significance of this 

lineation is unclear because the Cucamonga fault presumably dips north beneath the central Transverse Ranges 

(Morton, 1987). 

The cross section along the Mojave segment of the San 

Andreas fault (M-M', fig. 5.95) shows a wide variation in 

focal depths. The few events along the straight section of 

the Mojave segment are strongly clustered around 

depths of 10 km, with the maximum focal depth always 

above 15 km and almost no shallow earthquakes. This 

pattern is similar to the concentration of hypocenters in 

the lower half of the seismogenic crust beneath the locked 

1906 segment (compare with cross see. L-L', fig. 5.7). In 

contrast, the earthquakes at Tejon Pass cover the MI 

depth range from 0 to 25 km. At the southeast end of the 

1867 rupture zone at Cajon Pass, maximum focal depths 

increase with the seismicity level to a maximum of 20 km. 

In this section, there are no shallow (less than 5 km deep) 

events. 

Note, in particular, that focal depths of more than 20 

km are broadly associated with the convergent tectonics 

of the Transverse Ranges; they are not limited to the 

seismicity clusters in the vicinity of the bends in the San 

Andreas fault system as might be inferred from cross 

section M-M' (fig. 5.9B). Cross sections N-N' and 0-0' 

(fig. 5.95). for example, illustrate that the seismogenic 

crust reaches thicknesses of 20 to 25 km beneath the 

Santa Barbara Channel and the western Transverse 

Ranges, as well as beneath the Tehachapi Mountains to 

the east. 

SOUTHERN SECTION OF 

THE SAN ANDREAS FAULT SYSTEM 

Southeast of the 1857 rupture, the San Andreas fault 

splays into several branches associated with intense 

Northern Coast Ranges * - mabb Range 

Clear The 

H Lake Geysers - San Francisco Bay H' 

SEA 

LEVEL 

10 

20 

a 

2 - 1906 breek Creeping section z 1857 bra* - Son Francisco Ã‘Ã‘Ã of San Andreas faun 

San Juan Bautista 

30 I I I I I I I I I I I I 

0 30 60 90 120 150 180 210 240 270 300 330 360 

DISTANCE. IN KILOMETERS 

FIGURE 5.7.-Longitudinal depth sections along the San Andm fault system in the northern and central Coast Ranges. Bar above section L-L' 

indicates rupture extent of the October 17, 1989, Loma Prieta M=7.1 earthquake (see chap. 6). SSB. Suiaun Bay. See figure 5.6 for locations 

of sections and explanation of symbols, which are scaled with scale change of cross sections. 
130 THE SAN ANDREAS FA1 

seismicity in the Banning-San Gorgonio area. Like the 

Tejon Pass bend in the 1857 rupture zone, the San 

Gorgonio bend spawns a major left-lateral fault (the Pinto 

Mountain fault) conjugate to the San Andreas system. 

Unlike the situation at Tejon Pass, however, the San 

Andreas fault at San Go~pnio splays into a complex 

pattern of branching and intersecting fault segments. 

South of San Gorgonio, the San Andreas fault reconverges into a single strand and bends again to the more 

southeasterly trend that characterizes the southern section of the fault system. 

This section of the fault systemsouth of the Transverse 

Ranges is transitional from oblique spreading along the 

axis of the Gulf of California to the obliquely convergent 

strike-slip displacement that dominates deformation 

along the continental section of the San Andreas transform boundary to the north. Several major strike-slip 

faults run west of and subparallel to the main strand of 

the San Andreas fault south of the Transverse Ranges. 

These faults, which are considered part of the San 

Andreas system and include the Imperial, San Jacinto, 

and Elsinore faults, accommodate a significant proportion of the plate-boundary motion. The Imperial and San 

Jacinto faults, in particular, have produced more moderate earthquakes than any other segment within the fault 

system (see chap. 6; Hanks and others, 1975). 

SOUTHERN BRANCH OF THE SAN ANDREAS FAULT 

The most intense seismicity along the main trace of the 

southern section of the Ran Andreas fault is associated 

with the San Gorgonio bend and is concentrated between 

the two principal branches of the San Andreas fault: (1) 

the Mission Creek fault, or northern branch of the San 

Andreas; and (2) the Banning fault, which runs nearly 

due west from the south end of the Mission Creek fault 

toward an ambiguous junction with the San Jacinto fault 

just south of San Bernardino (see fig. 5.3A and maps at 

front of book). Neither strand forms a continuous structure through the bend. This San Gorgonio seismicity 

cluster produced numerous M=5.0-6.5 earthquakes in 

the 1930's and 1940's, and in 1986 it produced the ML=5.6 

North Palm Springs earthquake, which involved dextral 

strike-slip displacement on the north-dipping Banning 

fault (Jones and others, 1986). The background seismicity 

in this area is the highest in southern California, but it is 

distributed throughout a volume and cannot be clearly 

associated with any fault. To the west, seismicity associated with the Banning cluster abuts the dense lineation 

of epicenters coincident with the northernmost segment 

of the San Jacinto fault. Nicholson and others (1986) 

suggested that much of the seismicity within the upper 10 

km of the crust in this cluster involves left-lateral slip on 

a series of northeast-striking faults; however, Jones 

I SYSTEM, CALIFORNIA 

(1988) pointed out that the evidence for northeasttrending lineations of epicenters within the Banning 

cluster is less than clear. 

Diffuse seismicity extends northward from the Banning cluster into the San Bernardino Mountains and 

eastward into the Pinto Mountains, with no clear lineations along the sinistral Pinto Mountain fault. Indeed, a 

diffuse, north-south-trending limeation of epicenters 

seems to cut directly across the Pinto Mountain fault 

from the west-central Pinto Mountains. Two M=5.2 

earthquakes (see events 75, 76, fig. 5.11A) with rightlateral strike-slip planes parallel to this trend occurred at 

the north end of this zone in 1975 and 1979. Somewhat 

farther south, however, a broad, east-west-trending 

lineation appears to coincide with the Blue Cut fault. 

Even farther south, a second broad lineation extends 

eastward from near the junction of the Banning and 

Mission Creek branches, although this lineation does not 

coincide with a mapped fault. 

The southernmost section of the San Andreas fault, the 

Indio segment, which extends from the junction of the 

Banning and Mission Creek branches southward to the 

end of the San Andreas at the Salton Sea, has been almost 

completely aseismic in historical time. At the north end of 

this segment, periodic swarms of small (Ms4) earthquakes a few kilometers east of the San Andreas appear 

to occur on small northeast-trending structures (for 

example, Norris and others, 1986; Jones, 1988). The 

sparse background seismicity is also offset a few kilometers to the east from the surface trace of the San 

Andreas. Although the possibility of systematic offsets 

related to P-wave-velocity contrasts across the fault has 

not been investigated in detail, the observed offset seems 

too large to be explained entirely by lateral velocity 

contrasts. 

Although it has not ruptured with a major earthquake 

in historical time, the aseismic Indio segment of the San 

Andreas fault seems to have much in common with the 

1867 and 1906 rupture zones. Sieh (1986) presented 

geologic evidence for at least four major ruptures along 

the Indio segment since A.D. 1000; the last occurred 

approximately 300 yr ago. Unlike the two major locked 

sections, however, the south end of the Indio segment 

adjacent to the Salton Sea shows minor aseismic creep 

(Louie and others, 1985) and has shown episodes of 

sympathetic slip accompanying 44-6 earthquakes on the 

Imperial fault and the southern section of the San Jacinto 

fault (Sieh, 1982). Not only is the Indio segment aseismic, 

but also the entire Coachella block extending from the 

San Andreas fault on the northeast to the crest of the San 

Jacinto Mountains on the southwest. 

A cross section of hypocenters along the southern 

branch of the San Andreas fault (M-M', fig. 5.95) shows 

that the earthquakes associated with the bend at San 
6. SEISMICITY, 1980-86 131 

Northern Coast Ranges , Pacific Ocean P ____ Great ____ Valley E - SA - M BS E' 

SEA A 

LEVEL 

10 

SEA 

LEVEL 

& ," 

& 10 

3 Â 

a 

1- ia 

s 

Pacific Ocean - - Northern Coast Ranges - ,______ Great _____ me Valley F - SA M Geysers BS F' A - - 

Pacific Ocean 

< 

(San 

Coast 

FranciscoBay) 

Ranges - - Great ____ Valley G' RA H QV G' 

SEA 

LEVEL 

10 

20 

DISTANCE. IN KILOMETERS 

A 

FIGURE 6.8.-Tranaveree depth secliow amru the San Andreaa fault Springs; C, Calaveran; GV, Greenrille; H, Hayward: M. Maacama: 

system in the northern (A) and central (BJ Coasi Rangea. See fipme R. Rinconada, SA, San Andrew; SG. San Gregorio; SNA, Sur5.6 for locations of dona and explanation of symbols, which are Nacimienio. BV. Bear Valley on cross section J-J'. FRT, Franscaled with enlargement of cross sections. Faults; BS, Bartlett ciscan terrane. 
132 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

. ~oast Ranges . Pacific Ocean - <Ã‘Ã‘Ã rea at Valley - 

SEA 

LEVEL 

-FRT Salitian block 

-Pacific Ocean Ã‘Ã‘ 

SEA 

LEVEL 

s e 10 u 

5 

9 

1C 

z 

< Coast Ranges , -FRY SaKnian block FRT- 

-Pacific Ocean - . Great Valley 

SEA 

LEVEL 

I I I I I I I 

30 80 90 120 150 180 

DISTANCE, IN KILOMETERS 

B 
5. SEISMICITY, 1980-86 

Gorgonio are among the deepest in southern California, 

with maximum focal depths approaching 26 km. The 

maximum focal depths deepen southward along the San 

Andreas fault from about 12 km beneath the Mojave 

segment to 25 km beneath the San Gorgonio fault. At the 

south end of the San Gorgonio area, however, maximum 

focal depths abruptly decrease to 10 km. This shallowing 

of seismicity is associated with a shift in the most 

concentrated seismicity from between the two segments 

(Mission Creek and Banning) of the San Andreas to east 

of the Mission Creek fault. The sparse seismiaty of the 

Indio segment is limited to depths of 6 km or less. 

ASSOCIATED FAULTS 

Although the southernmost section of the San Andreas 

fault is almost completely aseismic, associated subpadlei faults are extremely active. These faults are marked 

by the three hold north-south- to northwest-trending 

alignments of epicenters that dominate the seismicity 

pattern within the San Andreas fault system south of the 

Transverse Ranges (fig. 5.10A), from east to west: (1) the 

Brawley seismic zone (Johnson, 1979), defined by a 

dense, spindle-shaped cluster of epicenters connecting 

the north end of the Imperial fault and the south end of 

the Indio segment of the San Andreas fault; (2) the 

northwestward alignment of densely clustered epicenters along the San Jacinto fault zone, which appears to 

branch from the northern section of the Imperial fault; 

and (3) the northwestward alignment of more diffusely 

clustered epicenters along the Elsinore fault, which 

appears to branch from somewhere near the south end of 

the Imperial fault. 

FIGURE5.9.-SeamIicitY inthe southernCoast RangesandTransverse Palmdale; SB, Santa Barbara; SMB, Santa Monica Bay; TP, Tejon 

Ranges. A, Earthquake locations, ahowingmajorbranchesof the San Pass. B, Depth sections outlined in figure 5.9.4. Faults: B, Banning; 

Andreas fault system in red; faidte dotted where concealed. Magni- G, Gal-loch; MC, UifiEuon Creek; N.Br.SA, northern branch of the San 

tude symbols shown in explanation are scaled with enlargement of Andreas; PM, Pinto Mountain; SA, San Andreas, SJ, San Jacinto; mom sections. C, Cholame; CP, Cajon Pax; P, &Meld; PI, WW, White Wolf. 

5. SEISMICITY, 1980-86 

The Brawley seismic zone and the cluster of epicenters 

at the south end of the Imperial fault (coincident with the 

Cerro Prieto volcanic-geothermal field in Mexico) represent the two northernmost in the series of small spreading centers offset by right-lateral transform faults that 

characterize oblique spreading in the Gulf of California 

(Lomnitz and others, 1970; Johnson and HiU, 1982). The 

Imperial fault itself, which is marked by a scattered 

alignment of epicenters, serves as the transform fault 

between these two small spreading centers. The M=7.1 

El Centre earthquake ruptured the entire length of the 

Imperial fault in 1940, and the M=6.6 Imperial Valley 

earthquake of 1979 ruptured the north two-thirds of the 

fault; intensity data suggest that moderate earthquakes 

(5.5<M<6.3) in 1906,1915,1917, and 1927 may also have 

been located on the Imperial fault (Johnson and Hill, 

1982). Most of the aftershocks associated with the 1979 

Imperial Valley earthquake were concentrated in the 

south half of the Brawley seismic zone, which was first 

recognized because of the many earthquake swarms it 

produced from 1973 through mid-1979 (Hill and others, 

1976; Johnson 1979; Johnson and Hutton, 1982). Many of 

the individual swarm sequences, as well as individual 

clusters of events in the aftershock sequence, defined 

lineations transverse to the strike of the Imperial fault 

and the long axis of the Brawley seismic zone. Most 

earthquakes within the Brawley seismic zone have 

strike-slip focal mechanisms; thus, kinematically, these 

transverse lineations represent conjugate structures to 

the dominant north-northwestward trend of the Imperial-Brawley fault system. 

Irregular clusters of epicenters mark the San Jacinto 

fault zone, which runs along the southwest base of the 

Santa Rosa and San Jacinto Mountains. These clusters 

tend to be concentrated near bends and junctions within 

the complex set of multiple fault strands that form the 

surface expression of this fault zone. In several places, 

particularly within the southern and northern sections of 

the fault zone, epicenters define linear concentrations 

that tend to be closely aligned with mapped fault traces. 

The San Jacinto fault zone has produced at least 10 

earthquakes of M=6.0-6.6 since 1890, the most recent of 

which were the M=6.2 earthquake of 1954, the M=6.6 

Borrego Mountain earthquake of 1968, and the M=6.6 

Superstition Hills earthquake of 1987. Thatcher and 

others (1975) pointed out that this series of historical 

M>6 earthquakes along the San Jacinto fault zone has 

left two seismic gaps: one along the northern 40 km of the 

fault, and the other along a 20-km-long stretch of the 

central section of the fault zone (the Anza gap). The Anza 

gap shows up in figure 5.10A as a relatively quiescent 

stretch of the fault zone between two dense clusters, with 

a third duster located off the fault zone some 20 km 

southwest of the gap (see Fletcher and others, 1987; 

Sanders and Kanamori, 1984). 

The Elsinore fault zone is defined not so much by a 

coincident alignment of epicenters as by the loci of 

western end points for clusters of epicenters elongate 

northeastward between the Elsinore and San Jacinto 

fault zones. This pattern is most pronounced along the 

southeast half of the fault; the northwest half, which 

defines the northeast scarp of the Elsinore Mountains, is 

marked by scattered clusters of epicenters. As the 

Elsinore fault enters the Los Angeles Basin to the north, 

it splays into the Whittier and Chino faults. Historical 

seismicity levels are considerably lower along the 

Elsinore fault than either the San Jacinto fault zone or 

the Imperial faultIBrawley seismic zone. The largest 

historical earthquake on the Elsinore fault was an M=6 

event in 1910 in the central section. The Whittier 

Narrows earthquake (Mr.=5.9) of 1987, which caused 

over $300 million in damage, was located at the north end 

of the Elsinore-Whittier fault. Because its mechanism 

was thrust faulting on an east-west-striking plane with a 

shallow dip, however, it does not appear to be simply 

related to the Elsinore system. 

Seismicity in the relatively quiescent southwestern 

corner of California between the Elsinore fault and the 

coast shows up in figure 5.10A as small, sparsely scattered clusters of epicenters. Activity picks up again, 

however, in the vicinity of the major northwest-striking 

faults along the coast (the Rose Canyon fault through San 

Diego and the Newport-Inglewood and Palos Verdes 

faults along the western margin of the Los Angeles 

Basin). Except for weak alignments along the NewportInglewood fault, which ruptured with an M=6.3 earthquake in 1933 (Richter, 1958), the seismicity patterns 

associated with these faults show little tendency to align 

along mapped fault traces. 

The cross sections in figure 5.105 emphasize that, 

except in the immediate vicinity of the Salton Sea, 

maximum focal depths associated with earthquakes 

aligned along the principal branches of the southern 

section of the San Andreas fault system are systematically deeper than those aligned along its central arid 

northern sections. Maximum focal depths, for example, 

decrease from 15 to 18 km beneath the central section of 

the Imperial fault near the United States-Mexican border 

to less than 10 km beneath the north end of the Brawley 

seismic zone at the southeast tip of the Salton Sea (cross 

sec. Q-Q', fig. 5.10B). The focal depths associated with 

earthquakes along the Coyote Creek, Superstition Hills, 

and Superstition Mountain faults forming the southwestern section of the San Jacinto fault zone adjacent to the 

Salton Sea are concentrated in the upper 10 km of the 

crust (cross sec. P-P', fig. 5.10B). 
136 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

Those segments of the San Andreas fault system in 

southern California with maximum focal depths shallower than 12 to 15 km are also those that show evidence of 

aseismk creep (see Louie and others, 1986). Indeed, 

actively creeping segments of the San Andreas fault 

system throughout California seem to be confined to 

those along which microearthquakes are concentrated in 

the shallow crust (focal depths of less than 12-15 km). 

Moving northwestward along the San Jacinto fault 

zone, the base of the seismogenic crust deepens systematically to a maximum of 20 km beneath the stretch 

adjacent to San Jadnto Mountain (which at 3,293 m, is 

the second highest point in southern California) midway 

along the fault zone (cross sec. P-P', fig. 5.10B). The 

base of the seismogenic crust maintains this 20-km depth 

farther northwestward along the fault zone to its junction 

with the Banning fault just south of San Bernardino (fig. 

5.10A), beyond which it begins to shallow again. Note, in 

particular, that earthquakes tend to be concentrated 

between 10- and 20-km depth beneath the San Jacinto 

fault zone, leaving the upper 10 km of the crust relatively 

quiescent along the middle stretch of the fault zone. The 

dense knot of hypocenters in the upper 5 km of the crust 

midway along cross section P-P' corresponds to the 

cluster of epicenters 15 km southwest of the fault zone 

near the Anza gap (fig. 5.10A). The Anza gap itself shows 

up between A=120 and 140 km in cross section P-P' as a 

quiescent zone below and southeast of the shallow cluster 

of hypocenters (Fletcher and others, 1987; Sanders, 

1987). The distribution of hypocenters beneath the 

Elsinore fault zone (cross sec. R-R', fig. 5.10B) is in 

many ways similar to that beneath the San Jacinto fault 

zone. Maximum focal depths increase northwestward 

from 12-15 km at the southeast end of the fault near the 

United States-Mexican border to about 20 km midway 

along the fault zone (generally coincident with the 

highest topography in this section of the Peninsular 

Ranges) and then gradually decrease farther northwestward toward the Los Angeles Basin. Maximum focal 

depths show evidence of increasing again at the northwest end of the fault as it approaches the Transverse 

Ranges and branches into the Whittier and Chino faults. 

The hypocenters along the south half of the Elsinore fault 

also tend to concentrate in the lower 10 km of the 

seismogenic crust, although this pattern is not as well 

defined in the diffuse seismicity of the Elsinore fault zone 

as in the dense clustering along the San Jacinto fault 

zone. 

FOCAL MECHANISMS AND 

TRANSFORM-BOUNDARY KINEMATICS 

Focal mechanisms of selected earthquakes recorded in 

California from 1933 through 1988 are shown in figure 

5.11A. and the corresponding source parameters are 

listed in tables 5.2 and 5.3. Primary considerations in the 

selection of these events were (1) size-larger events 

were chosen where available, because they represent 

large-scale processes along major boundaries; (2) date of 

occurrence-the quality of data for focal-mechanism 

determinations improved significantly during the mid1970's; and (3) location-some larger events were omitted 

because they were redundant in terms of mechanism and 

location, and some smaller events were included because 

they occurred in regions of significant seismicity where 

no larger events were available. Most focal mechanisms 

were determined from first arrivals at stations in the 

northern and southern California seismic networks. The 

evolving capability of these networks for such studies is 

reflected in the number of stations in the networks, 

summarized in table 5.1. Fault-plane solutions for the few 

large earthquakes on the list before the mid-1970's were 

supported by observations from stations outside the 

California networks. 

Since the mid-197Vs, focal mechanisms have been 

determined for only a fraction of the events for which 

adequate local first-motion data were available. Therefore, in addition to the three considerations listed above, 

there was a fourth, the interests of the investigators who 

analyzed the data. These interests included topical studies of large earthquakes and aftershock sequences, analyses of regional traveltimes on the basis of Ms4 

earthquakes, and a special study of the focal mechanisms 

of earthquakes on or near the San Andreas fault in 

southern California (Jones, 1988). 

Focal mechanisms discussed in the first two subsections below are for earthquakes in the contiguous Coast 

Ranges-Transverse/Peninsular Ranges-Mojave Desert 

region associated with the principal seismic expression of 

the San Andreas fault system, where the seismic networks are best developed. Outside that region, except for 

the Cape Mendocino area and the vicinity of Long Valley 

caldera, the few welldetermined focal mechanisms that 

are available provide only limited information on tectonic 

processes. 

STRIKE-SUP KINEMATICS OF 

THE SAN ANDREAS FAULT SYSTEM 

Most moderate and large (Ma) earthquakes along the 

San Andreas fault and its major branches produce nearly 

pure right-lateral displacements along near-vertical 

planes that closely follow the surface traces of the 

respective fault segments. This relatively simple kinematic pattern holds for the great earthquakes that 

rupture 'locked" sections of the fault every few hundred 

years (Sieh, 1981), as well as for nearly all the moderate 

earthquakes that rupture limited patches along persistently active segments of the fault system (Ellsworth and 

others, 1982; Jones, 1988). Displacements associated with 
5. SEISMICITY, 1980-86 137 

these earthquakes dominate the kinematic pattern along 

the transform boundary in California. DeMets and others 

(1987) and Minster and Jordan (1987). for example, 

argued that the cumulative displacement from earthquakes along the faults in the San Andreas system, 

together with the contribution from aseismic slip along 

its creeping segments, accounts for 60 to 70 percent of the 

total displacement between the Pacific and North American plates. 

The fault-parallel strike-slip displacements typical of 

San Andreas earthquakes are illustrated in figure 6.11.4 

by focal mechanisms along the San Andreas fault and its 

major branches from the United States-Mexican border 

to north of Clear Lake. In central California, such 

mechanisms mark the San Andreas fault itself from San 

Francisco to Cholame (events 26, 36, 38, 46, 46), the 

Calaveras-Greenville fault (events 23, 28-34) and the 

Hayward fault (event 27). Farther north, such mechanisms occur along the Green Valley-Bartlett Springs 

fault (event 16) and the Rodgers Creek-HealdsburgMaacama faults (events 16, 17, 19, 20). In southern 

California, such mechanisms mark the San Jacinto fault 

(events 78, 82-86) and the Imperial fault (event a). 

Along the coast west of the San Andreas fault, similar 

focal mechanisms occur along the San Gregorio-Palo 

Colorado fault (events 39, 40) in northern California and 

along the Newport-Inglewood fault zone (events 62, 71), 

the Rose Canyon fault (event 781, and the San Clemente 

fault (event 70) in southern California. 

Exceptions to this simple pattern for moderate (M24) 

events along the San Andreas fault and its major 

branches appear to be limited to regions of unusual 

complexity, such as the major bends in the San Andreas 

near Cajon Pass (event 69) and San Gorgonio Pass (event 

80). Jones and others (1986) attributed the July 8, 1986, 

earthquake (event 80) to right-lateral slip on the Banning 

FIGURE 5.10.-Seismicity along the southern section of the San zone; LB, Long Beach; MSJ, Mount San Jacinto, SB, Santa 

Andreas fault system. A, Earthquake locations, showing major Barbara, SBd, San Bernardino. 9, Depth sections outlined in 

branches of the San Andreas fault system in red; faults dotted 5.10A. Faults: CU, Cueamong% N1, Newwrt-Inglewood, W, 

where concealed. Magnitude symbols shown in explanation are Whittier. 

scaled with enlargement of cross sections. BZ, Brawley seismic 
Â£ 

INDIO SEGMENT m OF SAN ANDREAS 4 SAN JACINTO FAULT - - FAULT- BZ - -IMPERIALÃ‘ 

p' Q S07ea FAULT Q' 

SEA 

LEVEL ec 

s 10 G 

s 

z 30 ? 

UJ O30 11 I I I I I I I I I I I I I L 

0 30 60 90 120 150 180 210 240 0 30 60 90 120 i-3 DISTANCE. IN KILOMETERS 3 

Transverse - Ã ELSINORE FAULT * 

Ranges I-" 

R -&sin Angeb+ w R' ?i 

0 

E Â 

5 

s 

CO 3 a 

DISTANCE. IN KILOMETERS 

Transverse - E 

Ranges - Los Angeles Basin 2 

SEA 

n LEVEL 

I 

1 10 

G 

s 

z 

20 

30 

0 30 60 90 120 150 180 210 240 

DISTANCE, IN KILOMETERS 

B 
5. SEISMICITY. 1980-86 

segment of the San Andreas fault where it dips 46' N. 

beneath the San Bemardino Mountains. The October 17, 

1989, M=7.1 Loma Prieta earthquake involved nearly 

equal amounts of right and reverse slip along a section of 

the San Andreas fault that takes a slight westerly bend 

through the Santa Cruz Mountains and dips 70Â°SW (see 

chap. 6). Smaller (M<4) events near, but probably not 

on, the fault show a great variety of focal mechanisms 

that reflect varying conditions along the fault; these 

mechanisms range from reverse or reverse-oblique slip 

FIGURE 5.11.-Focal mechanisms for lareer earthquaken. A. California. SMB, Santa Monica Bay SS, San Simeon. Numbers refer to table 5.2. 

B, Coalinga-Kettleman Hills region (events 47-60, fig. 5.11A). Letters and numbers refer to table 5.8. Circle size increases with magnitude 

from 3.5 to 6.7. 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

TABLE 6.2.-Locations and famimmhanism parameters for selected eadkqmkea in 

California 
5. SEISMICITY, 1980-86 141 

TABLE 5.2.-Localimt and focal-mechanism pammeters for selected earthquakes in 

Califmia-Continued 

on easterly-striking planes (events 37, 67), through 

right-lateral strike slip on planes parallel to the San 

Andreas fault (events 65, 66, 68, 87), to normal or 

normal-oblique slip on northerly-striking planes (events 

77, 81). 

Moderate earthquakes with strike-slip focal-mechanisms that are not located on major faults of the San 

Andreas system but yet are broadly associated with it 

commonly have right-slip planes, with strikes ranging 

from northwestward (event 42) to north-southward 

(events 35, 44, 74, 75, 76, 86). In most cases, these 

right-slip planes agree in strike with local mapped faults 

or with alignments of epicenters that strongly suggest 

active faults (events 75, 76). 

CRUSTAL CONVERGENCE ADJACENT TO 

THE SAN ANDREAS FAULT SYSTEM 

One of the more important results to emerge from 

high-resolution focal-mechanism studies in recent years 

is that earthquakes occurring even a short distance off 

faults of the San Andreas system can involve displacements that diverge sharply from local San Andreas 

strike-slip displacements. This pattern is particularly 

pronounced in the strong component of reverse slip at 

large angles (more than 607 to the local strike of the San 

Andreas fault on both sides of the San Andreas fault 

system in both the Transverse and Coast Ranges. 

North-south convergence within the Transverse Ranges is dominated by reverse slip on easterly-striking 

planes. The M=7.7 Kern County earthquake of 1952 

(event 64, which occurred on the south-dipping White 

Wolf fault along the north flank of the Transverse Ranges 

about 25 krn north of the junction of the San Andreas and 

Garlock faults, and the M=6.6 San Fernando earthquake 

of 1971 (event 71), which ruptured a 20-km-long stretch 

of the northeast-dipping San Gabriel-San Fernando 

thrust faults (Whitcomb, 1971; Heaton, 1982), are two 

striking examples of this deformation. So, also, is the 

alignment of M=5-6 reverse-slip earthquakes (events 59, 

60, 63) along the southern margin of the Transverse 

Ranges. The reverse slip on east-west-striking planes 

associated with these earthquakes suggests that the 

north-dipping Santa Monica-Cucamonga fault serves as 

an important convergent boundary between the Peninsular and Transverse Ranges. 

Figure 5.11A also shows that the east-west-trending 

zone of convergence associated with these earthqutikes 
142 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

TABLE 5.3.-Locations and focal-meehanism pammetem for earthgtdes in the CoalingaKettleman Hilts region 

[Same symbols .a in table 5.21 

curves northward near Santa Monica Bay and continues 

northwestward along the coast at least as far as Point Sal 

and probably as far as San Simeon. Focal mechanisms of 

earthquakes along this zone from Point Sal to Whittier 

(events 55-60, 63) are predominantly reverse slip, with 

slip directions nearly normal to the local trend of the 

zone. The focal mechanism of event 43 near San Simeon, 

which indicates right-oblique reverse slip on a northeastdipping plane parallel to the coast, is intermediate 

between those of event 40 at Point Sur and event 55 at 

Point Sal. 

Reverse-slip focal mechanisms for offshore events 18 

and 41 in central California and for event 72 in southern 

California suggest that the offshore crust is undergoing 

compression normal to the coastline throughout the 

length of the San Andreas fault system. The CoalingaKettleman Hills earthquake sequence of 1982-85 (events 

47-60, fig. 5.11A) emphasizes the important role of 

crustal convergence along the southern Coast RangesGreat Valley boundary in central California. The principal events in this sequence (event 48, Coalinga, and event 

49, Kettleman Hills, fig. 5.11A) involved reverse slip on 

subparallel planes at depths of 10 to 12 km that dip gently 

(approx 20') southwest. Much of the aftershock activity, 

however, occurred at shallower depths and involved 

high-angle reverse slip on planes dipping steeply (45'- 

70') northeast (events f, g, i, o, q, fig. 5.11B). Displacements associated with these earthquakes, which are 

nearly perpendicular to the San Andreas fault, represent 

168 

90 

90 

73 

Ill 

so 

168 

102 

85 

161 

90 

90 

85 

90 

105 

82 

82 

113 

112 

90 

- 

98 

a convergent process in which Franciscan melange on the 

west is being wedged between crystalline basement and 

the overlying Great Valley sedimentary sequence on the 

east (Wentworth and others, 1983, Baton, 1990). 

The boundary between the Coast Ranges and Great 

Valley is marked by reverse-slip earthquakes throughout 

much of its length: event 54 southeast of the Kettleman 

Hills, the Coalinga-Kettleman Hills sequence, event 21 

near Winters east of Lake Berryessa, and event 14 west 

of Oroville. The similarity in focal mechanism of event 21 

near Winters to the Coalinga and Kettleman Hills main 

shocks suggests that the convergent process acting in the 

southern Coast Ranges is common to the entire eastern 

margin of the Coast Ranges. Indeed, the strong earthquakes that shook the Winters-Vacaville-Dixon area in 

1892, just south of event 21, resemble the CoalingaKettleman Hills sequence in both setting and intensity 

distribution. Focal mechanisms of smaller earthquakes 

along the Coast Ranges-Great Valley boundary in central 

California studied by Wong and others (1988) also suggest convergence across that boundary. 

Convergence normal to the strike of the San Andreas 

fault is not limited to the coast and the Coast RangesGreat Valley boundary described above. In a detailed 

examination of the focal mechanisms of aftershocks of the 

1984 Morgan Hill earthquake, Oppenheimer and others 

(1988) concluded that the direction of maximum compression immediately adjacent to the Calaveras branch of the 

San Andreas fault is at an angle of about 80Â to the N. 10' 
5. SEISMICITY, 1980-86 143 

W. strike of the fault. Along the entire stretch of the San 

Andreas fault from Parkfield to the Salton Sea, Jones 

(1988) found a constant angle of 66' between the strike of 

the fault and the maximum-principal-stress direction for 

earthquakes occurring off the fault. 

This evidence from earthquake focal mechanisms and 

other stress indicators (such as horehole breakouts and 

fold axes) that the maximum principal compressive stress 

may be oriented at a high angle to the local strike of the 

San Andreas fault seems to contradict long-accepted 

ideas for brittle failure in the crust based on laboratory 

experiments in rock mechanics. Zoback and others (1987) 

and Oppenheimer and others (1988) suggested that these 

relations can be explained by an exceptionally low 

average shear strength for the San Andreas fault system. 

As pointed out by Lachenbruch and McGarr in chapter 

10, however, the strength and state of stress along the 

San Andreas fault are still matters for discussion. 

EAST-WEST EXTENSION IN THE SIERRA NEVADA 

The three moderate-earthquake focal mechanisms for 

the Sierra Nevada and its western foothills shown in 

figure 5.11A all indicate normal faulting on northerlystriking planes and suggest pervasive east-west extension throughout the Sierra Nevada. Event 52 is in a dense 

north-south-trending band of epicenters about 15 km east 

of the Kern River canyon, and event 53 is in a northsouth-trending hand of earthquakes about 10 km west of 

the Sierra frontal fault near Walker Pass. These relations 

suggest that the east-west spreading and associated 

normal faulting on northerly-striking faults of the Great 

Basin are encroaching into the southeast comer of the 

Sierra Nevada block (Jones and Dollar, 1986). 

Event 13 is the main shock (M=5.7) of an earthquake 

sequence on a north-south-striking, westdipping normal 

fault near the Oroville Dam that occurred in 1975. The 

uplift of the Sierra Nevadarelative to the Great Valley to 

the west indicated by the focal mechanism of this event is 

also visible in the Chico monoclie, which marks the 

Sierra Nevada-Great Valley boundary northwest of 

Oroville. 

CONJUGATE FAULTING IN 

THE SIERRA NEVADA-GREAT BASIN BOUNDARY ZONE 

The Sierra Nevada-Great Basin boundary zone is 

represented in figure 5.11-4 by three focal mechanisms. 

Events 11 and 12 lie northwest of Lake Tahoe along the 

edge of a minor gap in the band of seismicity along the 

east edge of the Sierra Nevada. Both events appear to 

have resulted from left-lateral slip along steeply dipping, 

northeast-striking faults; both events had aftershock 

regions that were elongate northeast-southwest. About 

250 km southeast, the 44-6.4 July 21, 1986, Chalfant 

Valley earthquake (event 51) resulted from right-lateral 

strike-slip displacement on a north-northwest-striking 

surface dipping 60' SW. An M=5.7 foreshock on July 20 

resulted from left-lateral strike slip on a northeaststriking, northwest-dipping surface. These two conjugate slip surfaces merge at their north ends (Cockerham 

and Corbett, 1987; Smith and Priestly, 1988). 

The zone of intense seismicity in the vicinity of Long 

Valley caldera and the Sierra Nevada block to the south 

produced 11 M~5.5 earthquakes from 1978 through 1984 

(Savage and Cockerham, 19871, as well as many thousands of smaller events and numerous earthquake 

swarms. Most of the larger events occurred in the Sierra 

Nevada block south of Long Valley caldera, involving 

left-lateral slip along near-vertical, north-south- to north 

northeast-strikine - faults. One ~ of ~ four ~~~ M-6 -- ~ events that - -~~~~ ~~~~~. 

occurred on May 25-27, 1980, however, was located 

within the south moat of the caldera along the westnorthwest-striking fault zone that produced most of the 

earthquake swarms (see Hill and others, 1985a, b). 

FRAGMENTATION OF 

THE SOUTHEAST CORNER OF THE GORDA PLATE 

The 1980 Eureka M=7.2 earthquake occurred along a 

fault break that extended from the continental slope 40 

km west of the coastline at lat 41" N. for a distance of 140 

km southwestward to the MFZ, virtually cutting off the 

southeast comer of the Gorda plate. Focal mechanisms of 

the main shock and largest aftershock (events 1, 3, fig. 

5.11A) both indicate left-lateral strike-slip displacement 

along a vertical fault that coincides with the line of 

aftershocks. Some early aftershocks, including event 4 

and other moderate events farther east along the MFZ, 

have focal mechanisms that indicate right-lateral slip 

along the MFZ. Although the main shock occurred 

beneath the Continental Shelf, it seems clear that the 

1980 earthquake primarily involved the Gorda plate 

because the fault broke well beyond the base of the 

continental slope and the edge of the North American 

plate. Moreover, left-lateral slip along the 1980 break 

stimulated right-lateral slip along the adjacent part of the 

MFZ. Ongoing right-lateral displacement along the MFZ 

is also indicated by event 5 (Dec. 1983). 

Two moderate earthquakes near Cape Mendocino in 

1981 and 1987 (events 6 and 7, respectively) had focal 

mechanisms similar to that of the 1980 Eureka earthquake, indicating left-lateral strike-slip displacement on 

steeply dipping, northeast-striking planes. Aftershocks 

of the 1987 M=5.8 event outlined a narrow, steeply 

dipping, northeast-trending, 20-km-long zone between 

about 15- and 25-km depth that extended southwestward 

from the shoreline just north of Cape Mendocino to the 
144 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

MFZ. This aftershock zone appears to cut off the 

southeasternmost comer of the Gorda plate just north of 

the abrupt eastward termination of intense seismicity 

along the MFZ, at a point that might be taken as the 

Mendocino triple junction from the viewpoint of seismicity. 

Relative horizontal extension at seismogenic depths is 

suggested by events 2 and 8. Event 2 (Nov. 10, 1980; 7 

km deep) was the largest in a detached cluster of shallow 

aftershocks 20 km east of the 1980 main shock, and event 

8 (Apr. 9,1987; 26 km deep) occurred about 100 km east 

of the 1980 main shock in the zone of seismicity associated 

with the subducting Gorda plate. 

DISCUSSION 

The Pacific plate moved northwestward with respect 

to the North American plate by 300 to 400 mm during the 

7-yr interval 1980-86. Earthquakes occurring along the 

San Andreas fault system during the same interval, 

however, accommodated only a small fraction of this 

relative plate motion. Only four earthquakes of M>5 

occurred along branches of the San Andreas fault system 

during 1980-86. the pair of M=5.9-6.3 Livermore earthquakes (events 29, 30, fig. 5.10A) on the Greenville fault 

(Jan. 24-27, 1980), the M=6.2 Morgan Hill earthquake 

(event 33) on the Calaveras fault (Apr. 24,19841, and the 

M=5.6 North Palm Springs earthquake (event 80) on the 

Banning segment of the San Andreas fault (July 8,1986). 

Each of these moderate San Andreas earthquakes ruptured fault segments limited to 20 to 30 km in length, with 

average displacements over the respective rupture surfaces of 100 to 200 mm (see Hartzell and Heaton, 1986). 

As is typical of earthquakes along the San Andreas fault 

system, each of these events involved nearly pure 

right-lateral strike-slip displacement coincident with the 

local strike of the fault. As is also typical of San Andreas 

earthquakes, slip on the first three events occurred on 

near-vertical fault planes with a northwestward to northnorthwestward strike. The North Palm Springs earthquake, which ruptured a section of the east-west-striking 

Banning fault in the structurally complex San Gorgonio 

bend in the fault system at the southern margin of the 

Transverse Ranges, represents an important deviation 

from typical San Andreas earthquakes. Although its 

displacement was dominantly right-lateral strike slip, it 

occurred along a plane that dips 46Â N. (Jones and others, 

1986) and included a small but significant component of 

reverse slip (Mendoza and Hartzell, 1988). With the 

arguable exception of the North Palm Springs earthquake (arguable because of the complex section of the 

fault systemin which it occurred), however, none of these 

M>5 earthquakes ruptured the main trace of the San 

Andreas fault. Indeed, the two most recent M>5 earthwakes to clearly do so were the M=6 Parkfield earthluake of 1966 (Bakun and McEvilly, 1984) and the M=7.1 

Loma Prieta earthquake of 1989 (see chap. 6). 

Thus, aside from the displacement accommodated by 

iteady aseismic slip at a rate of 32 to 37 mmlyr along the 

creeping section of the fault in central California, most 

relative plate motion across the San Andreas transform 

boundary during this 7-yr interval accumulated as elastic 

shear strain. Accordingly, the earthquakes plotted in 

figures 5.3 through 5.9 are symptomatic of accumulating 

strain along the San Andreas fault system rather than of 

affective strain release. The latter requires rupture with 

a major earthquake along one of the locked stretches of 

the San Andreas fault. 

SEISMICITY PATTERNS AND THE EARTHQUAKE CYCLE 

What changes in spatial-temporal patterns of earthquake occurrence might we expect to see as the next 

great earthquake on the San Andreas fault approaches? 

Both historical and instrumental seismicity records indicate that the spatial distribution of earthquakes in 

California changes only slowly over periods of decades to 

centuries, although the intensity of activity within this 

distribution fluctuates year to year (Ellsworth and others, 1981; Hill and others, in press; Hutton and others, in 

press). Temporal fluctuations in activity during the 

intervi-1 1980-86, for example, were dominated by a 

short-lived aftershock sequence following the 1980 Eureka M=7.2 earthquake and by the long-lived aftershock 

sequence following the 1983 Coalinga M=6.7 earthquake. 

The overall spatial distribution of earthquakes in California, however, remained nearly stationary throughout 

this 7-yr interval. Furthermore, the spatial pattern 

defined by 1980-86 seismicity is much the same as that 

outlined by the record of Ma5 earthquakes that extends 

back nearly 200 yr (see chap. 6). 

variations in the historical rate of moderate to large 

(M>5\ earthouakea in central California before and after 

the 1906 San Francisco earthquake appear to mimic those 

described by Fedotov (1965) and Mogi (1968) for the 

earthquake cycle associated with great, subduction-zone 

earthquakes in Japan, Kamchatka, and the Kurile Islands 

(see chap. 6; Ellsworth and others, 1981). The history of 

instrumentally recorded M<5 earthquakes in California 

is too short, however, to indicate whether we might 

expect to see distinctive changes in the seismicity pattern 

a short time (months to years) before the next great 

earthquake on the San Andreas fault. We have yet to see, 

for example, whether the quiescent (locked) segments of 

the San Andreas fault remain aseismic except for the 

rupture of a great earthquake, or whether these segments become active with small to moderate earthquakes 

as foreshock activity to great earthquakes. 
6. SEISMICITY. 1980-86 145 

DISTRIBUTED SEISMICITY AND 

DEFORMATION OF THE PLATE MARGINS 

The two largest earthquakes in California during the 

interval 19M occurred off the faults of the San 

Andreas system, and their occurrence emphasizes the 

importance of deformation within the plate margins along 

the San Andreas transform boundary. The M=7.2 Eureka event (Nov. 8, 1980). for example, involved deformation internal to the Gorda plate; and the M=6.7 Coalinga 

event (May 2, 1983) involved crustal shortening with 

reverse slip perpendicular to the San Andreas fault. 

These two earthquakes and the many smaller, "off fault" 

events (fig. 5.44 reflect local deviations from the simple 

rigid-plate approximation of plate tectonics. 

DEFORMATION OF THE GORDA PLATE 

As the small, youthful Gorda plate is subducted obliquely northeastward beneath the North American 

plate, it is being subjected to north-south compression in 

response to a component of convergence between the 

larger, older Juan de Fuca plate to the north and the 

Pacific plate to the south (Jachens and Griscom, 1983, 

Wilson, 1986). Distorted marine magnetic anomalies 

within the Gorda plate indicate that it has undergone 

progressive internal deformation over the past 5 Ma in 

response to this compression (Silver, 1971), and current 

seismicity within the plate (fig. 5.4) indicates that this 

deformation continues to the present. 

The 1980 Eureka M=7.2 earthquake emphasizes that 

part of this deformation occurs with left-lateral slip on 

northeast-striking faults within the plate. The seismicity 

map and cross sections (fig. 5.4) demonstrate that deformation associated with the Gorda plate terminates 

abruptly against the Pacific plate in a steeply northdipping zone of interaction along the MFZ, which can be 

followed on shore beneath the North American plate as a 

gently east-dipping, subhorizontal zone of widely scattered small to moderate earthquakes. Thus, convergence 

between the Gorda and Pacific plates across the MFZ 

apparently occurs by crushing and thickening of the 

southern margin of the Gorda plate as it is jammed 

against the anvil-like mass formed by the thicker and 

colder Pacific plate. Diminished east-west stress in the 

Gorda plate resulting from the subducting limb of the 

plate farther east serves to increase the difference 

between the maximum (north-south) and minimum (eastwest) compressive stresses within the plate, leading to 

left-lateral strike-slip displacements along northeaststriking faults, as in the M=7.2 Eureka earthquake. This 

process accommodates the convergent component of 

Gorda-Pacific plate motion along the east end of the MFZ 

at the expense of fragmentation and eastward expansion 

of the Gorda plate north of the MFZ. 

THE SAN ANDREAS DISCREPANCY 

Much of the seismicity adjacent to the San Andreas 

fault system is attributable to differences between the 

long-term slip rate and direction (slip vector) along the 

Sail Andreas fault system and that predicted for relative 

motion between the Pacific and North American plates 

along the San Andreas transform boundary on the basis 

of global models of plate motion. Minster and Jordan 

(1978, 1987) predicted that the direction of dextral slip 

between the Pacific and North American plates along the 

Sail Andreas transform boundary in central California is 

N. 35O W. The main trace of the San Andreas system, 

however, strikes N. 41'W. through central and northern 

California and N. 65'-70Â W. through the Transverse 

Ranges in southern California. DeMets and others (1987) 

concluded that the marine magnetic anomalies at the 

mouth of the Gulf of California constrain the slip rate to 

an average of 49 mmlyr over the past 3 to 4 Ma. Both 

long-term geologic offset data and geodetic data measured over the past several decades, however, indicate 

that the average slip rate along the San Andreas fault 

system is only about 35 mm/yr. The contribution to 

deformation of the western margin of the North American plate from spreading across the Basin and Range 

province is about 10 mmlyr in a N. 56' W. direction 

(Minster and Jordan,1987). Ellsworth (see chap. 6) suggests that most of the San Andreas discrepancy can be 

explained if the component of dextral slip associated with 

historical Basin and Range earthquakes reflects a longterm trend superimposed on the N. 56' W. spreading 

direction. If so, then the residual component of Basin and 

Range extension perpendicular to the San Andreas fault 

system is approximately balanced by convergence across 

the Coast Ranges and continental margin. 

CONVERGENCE NORMAL TO 

THE SAN ANDREAS FAULT SYSTEM 

Focal mechanisms of earthquakes occurring off the San 

Andreas fault system suggest that the component of the 

Sail Andreas discrepancy normal to the fault system 

may, indeed, be accommodated by distributed brittle 

deformation on either side of the fault system. These 

mechanisms range from dextral strike slip on planes 

subparallel to the San Andreas fault, through ohliquereverse slip, to nearly pure reverse slip with a slip 

direction perpendicular to the San Andreas fault. 

The Coalinga-North Kettleman Hills earthquake sequence provides clear evidence for crustal convergence 

perpendicular to the San Andreas fault system in the 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

ga earthquake may beactive the length of the Coast 1984). Namson and Davis (1988) proposed that the entire 

Ranges (Wong and others, 1988, Eaton and Rymer, systemofcoast Rangefoldsmay begenetically relatedto 

Coast Ranges. The several smaller events with similar 

mechanisms to the north along both the eastern and 

western (coastal) margins of the Coast Ranges (fig. 5.11) 

suggest that the convergence responsible for the CoalinFIGURE 5.12.-Seismiaty flwn 1980 to 1986 superimposed on digital shaded-reHefima@ of central California, showing faults @hie) and fold axes 

(red). Size of symbol for epicenters (yellow) increases with magnitude from 1 to 6. Shaded reliefby Raymond Batson, U.S. Geological Survey 

(illumination from north at 30Â¡) overlays from ROBS Stein (unpub. data, 1989). SAF, San Andreaa fault. 

1990). The subparallelism of fold axes within the Coast 

Ranges with the San Andreas fault indicates that faultnormal convergence has been important for the past 3 Ma 

in central California (fig. 5.12; Page and Engebretson, 
5. SEISMICITY, 1980-86 

Coalinga-like earthquake sequences and low-angle (blind) 

thrust faults that are rooted in a decollement near the 

base of the seismogenic crust. The reverse focal mechanisms for earthquakes associated with offshore faults 

along the western margin of the Coast Ranges suggest 

that, here, convergence involves westward thrusting of 

the Coast Ranges over oceanic crust of the Pacific plate. 

The pronounced discrepancy in the strike of the San 

Andreas fault through the Transverse Ranges with 

respect to the Pacific-North American plate slip direction 

provides an obvious source of local crustal convergence 

(Hill and Dibbles, 1953; Atwater, 1970), and the associated structural complexities serve to distribute brittle 

deformation (seismicity) much more widely about the San 

Andreas fault svstem in southern California than about 

the relatively straight sections of the fault system in 

central and northern California. The largest earthquake 

in California since the great 1906 San Francisco earthquake occurred near the northern margin of this convergent regime; this M=7.7 Kern County earthquake 

ruptured some 35 krn of the southeast-dipping White 

Wolf fault with lefroblique reverse slip on July 21, 1952. 

The focal mechanisms of larger Transverse Range 

earthquakes, together with the mapped attitudes of 

major faults with Holocene offsets, show that much of 

this convergence occurs with slip on north-dipping thrust 

faults within and along the southern margin of the central 

Transverse Ranges (fig. 5.11A). For earthquakes in the 

western Transverse Ranges, the direction of reverse slip 

is more southwestward, consistent with thrusting of the 

western Transverse Ranges over the Pacific plate similar 

to that in the Coast Ranges to the north. 

EXTENSIONAL DEFORMATION AND THE SOUTHERN 

SECTION OF THE SAN ANDREAS FAULT SYSTEM 

The fault-normal convergence that dominates deformation adjacent to the San Andreas fault system through 

both the Coast Ranges and Transverse Ranges gives way 

rather abruptly to the extensional regime of the Salton 

Trough near the southern margin of the intensely active 

San Gorgonio bend in the fault. Focal mechanisms of 

earthquakes occurring on secondary structures adjacent 

to the seismically quiescent Indio segment of the San 

Andreas fault, for example, show a mix of strike- and 

dip-slip mechanisms. As is the case farther north, however, P-axes for these earthquakes tend to be oriented at 

a high angle (W-659 to the fault, suggesting that the 

Indio segment of the fault may also be relatively weak 

(Jones, 1988). 

One particularly noteworthy aspect of seismicity south 

of the Transverse Ranges is the. tendency for earthquakes to occur along conjugate strike-slip structures. 

Recall that the Sierra Nevada-Great Basin boundary 

zone also shows this tendency and that both regions are 

subject to extensional deformation, earthquake swarms, 

and late Quaternary volcanism. Earthquake sequences 

within the southern section of the San Andreas fault 

system commonly produce epicenter lineations that intersect at nearly a 90' angle with the northwest-striking 

right-slip plane and the northeast-striking left-slip plane. 

Earthquake-swarm sequences in the Brawley seismic 

zone, for example, typically occur along northeast-striking lineations normal to the trace of the adjacent Imperial 

fault (Johnson, 1979), and the M-6.7 Westmorland 

earthquake of 1981 involved left-lateral slip along several 

subparallel, northeast-striking planes (Johnson and Hutton, 1982). The diffuse lineations of epicenters spanning 

the area of the Peninsular Ranges between the San 

Jacinto and Elsinore faults also tend to be orthogonal to 

these two branches of the San Andreas fault system (fig. 

5.10A). An impressive recent example of this orthogonal 

conjugate pattern is the M=6.2 and6.6 superstition- ills 

earthquakes of November 24, 1987 (Magistrate and 

others, 1988). 

The kinematics of these conjugate structures remains 

a matter of conjecture. Dextral slip along throughgoing 

faults of the San Andreas system must certainly dominate deformation, and the shorter, northeast-striking 

structures must play only a secondary role. Nicholson 

and others (1986) proposed that the northeast-striking 

lineations represent the boundaries between blocks rotating clockwise much like roller bearings, between 

subparallel pairs of dextral strike-slip faults. Hill (1977) 

and Weaver and Hill (1978179) suggested that within local 

spreading centers, such as the Brawley seismic zone, 

conjugate strike-slip structures form miniature triple 

junctions with a dike or normal fault that subtends the 

acute angle between the conjugate strike-slip faults. 

MAXIMUM FOCAL DEPTHS AND THICKNESS OF 

THE SEISMOGENIC CRUST 

Maximum focal depths of earthquakes beneath the San 

Andreas transform boundary range from less than 5 km 

beneath the Geysers geothermal field in the northern 

Coast Ranges to more than 20 km beneath the Transverse Ranges, the eastern margin of the Coast Ranges, 

and the San Jacinto and Elsinore faults in southernmost 

California. Beneath relatively straight segments of the 

San Andreas fault system through central California, 

maximum focal depths range from 12 to 15 km (figs. 5.7, 

5.8). Sibson (1983) pointed out that these variations in 

maximum focal depth along the San Andreas fault system 

are inversely correlated with surficial heat flow, and he 

armed that the maximum death of earthouakes coincides 

with the temperature-dependent transition from brittle 

failure in the upper crust to aseismic, quasi-plastic flow in 
148 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

the lower crust and upper mantle. For quartz-bearing 

rocks typical of the upper crust and deformation rates 

typical of the San Andreas fault system (l~lO'~~ to 

lx lo-" s-I), this brittlelductile transition occurs at about 

300 'C (Sibson, 1983). By this interpretation, the thin 

seismogenic crust beneath both the Geysers and Brawley 

geothermal fields in northern and southern California, 

respectively, reflects elevated temperatures in the shallow crust, whereas the relatively thick seismogenic crust 

beneath the Transverse Ranges and the eastern margin 

of the Coast Ranges reflects depressed temperatures in 

the midcrust associated with crustal convergence. Although temperature may dominantly influence the thickness of the seismogenic crust, local variations in rock 

composition (particularly the presence or absence of 

modal quartz and structural water) and in strain rate can 

also be important. These variations, for example, may 

help explain isolated clusters of deep earthquakes, such 

as the 20- to 24-km-deep events north of San Pablo Bay 

in central California (see cross sees. F-F', GG', fig. 

5.85). 

In any case, the thickness of the seismogenic crust 

beneath the San Andreas transform boundary seems to 

be much more strongly related to temperatures in the 

crust than to the structural thickness of crust defined by 

the depth to the Moho (see chap. 8). This relation is 

strikingly illustrated by the twofold increase in thickness 

of the seismogenic crust beneath the rootless Transverse 

Ranges. 

DECOLLEMENT AT THE BASE OF 

THE SEISMOGENIC CRUST? 

A theme common to models of crustal convergence 

along the San Andreas fault system involves low-angle 

reverse slip on decollement surfaces near the base of the 

seismogenic crust (Wentworth and others, 1983; Webb 

and Kanamori, 1985; Namson and Davis, 1988; Baton and 

Rymer, 1990). A natural extension of this theme leads to 

a view of the seismogenic crust as a conglomeration of 

relatively rigid blocks interacting by frictional slip along 

weak preexisting faults (block boundaries) in response to 

regional stresses transmitted through both the brittle 

crust and quasi-plastic deformation in the underlying 

lithosphere (Hill, 1982). However, the nature of a decollement surface at the base of the brittle crust and the 

relation of the seismogenic San Andreas fault system to 

the aseismic transform boundary in the underlying litbosphere remain speculative. It is not yet clear, for 

example, whether the San Andreas fault continues below 

the seismogenic crust as a narrow, near-vertical boundary (possibly offset a substantial distance from the 

seismogenic fault by slip on the horizontal decollement 

surface) that slips by quasi-plastic, mylonitic deformation 

or whether it broadens rapidly with depth into a wide 

shear zone spanning, say, the entire width of the Coast 

Ranges (see chap. 7; Sibson, 1983). 

CONCLUSIONS 

The spatial-temporal pattern of earthquake occurrence 

within the seismogenic crust along the San Andreas fault 

system is the brittle manifestation of distributed deformation of the lithosphere between the Pacific and North 

American plates along the San Andreas transform boundary. As we develop a more complete model of the 

long-term behavior of the seismogenic crust, including 

relations between great, plate-boundary earthquakes 

that periodically rupture the principal strand of the San 

Andreas fault system and the persistent background of 

small to moderate earthquakes on adjacent structures, 

our image of the deeper deformation will improve. 

Together will come a more complete understanding of the 

processes controlling deformation along the transform 

boundary and of the earthquake cycle. 

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150 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

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5. SEISMICITY. 1980-86 

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Healy, J.H., Oppenheimer, D.H., Reasenbe%, P.A.. Jones, L.M., 

Raleigh, C.B.. Wong, I.G., Scotti. Oona, and Wentworth, C.M., 

1987, New evidenceon the state of stress of the San Andmas fault 

system: Science, v. 238, no. 4830, p. 1106-1111. 

M otion between the North American and Pacific plates at the 

latitude of the Son Andreas fault produces a broad zone of 

large-magnitude earthquake activity extending more than 500 km 

into the continental interior. The San Andreas fault system defines 

the western, limits of plate interaction and dominates the overall 

pattern of seismic strain release. Few of the Ma6 earthquakes that 

have occurred in the past 2 centuries were located on the San 

Andreas fault proper, an observation emphasizing the importance 

of seconSary faults for both seismic-hazard assessment and tectonic 

processes. 

6. EARTHQUAKE HISTORY, 1769-1989 

CONTENTS 

4 FIGURE 6.1.-Seu>mi&y highiighte broad and complex zones of active tectonism in the western United States. Dots, earthquakes acaled by 

magnitude in unit-magnitude steps, emallst for M<4; circles, earthquakes of M>7; red dots, recent events, generally since 1975; pink 

symbols, historical events. From Engdahl and Rinehart (1988). 

IBS 

THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA; US. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

INTRODUCTION 

Between Punta Gorda on the northern California coast 

and the head of the Gulf of California, 1,360 km to the 

southeast, lies the active transform boundary that forms 

the modem San Andreas fault system (fig. 6.1). Dextral 

motion between the North American and Pacific plates 

along this system is accommodated within an elongate 

zone, broadening from about 100 km at its north end to 

about 300 km in southern California. The San Andreas 

fault proper hugs the east side of this zone at its south 

terminus and gradually migrates across the zone, lying 

on the west edge of the zone at its north terminus. The 

San Andreas fault system transmits about three-fourths 

of the relative motion across the plate boundary, a8 

shown by various geologic and geodetic evidence. Much 

of this motion is stored elastidy in the upper crust along 

the major faults in the system, ultimately to be released 

in large plate-boundary earthquakes. These large earthquakes and their implications for the mechanics of North 

American-Pacific plate interactions are the subject of this 

chapter. 

Earthquake activity in California and Nevada at the 

latitude of the San Andreas fault extends well beyond the 

confines of the San Andreas system (fig. 6.2). In the past 

century alone, only about half of the MaÂ activity has 

0 - 200 KILOMETERS 

FIGURE 6.2.-Seismicity of California, Nevada, and northern Baja California, 1769-1989. Earthquakes are listed in table 6.1 and plotted by 

magnitude class. 
6. EARTHQUAKE HISTORY, 1769-1989 

fallen within the San Andreas system; of the rest, half is 

associated with the western Basin and Range province, 

and the other half with the Mendocino triple junction and 

the Gorda plate. Although activity in the latter region 

reflects the tectonics of the triple junction and the 

collision of the Gorda plate with the North American 

plate, seismicity east of the San Andreas system along 

the east flank of the Sierra Nevada and in the Basin and 

Range province reflects the incomplete accommodation of 

plate motion along the San Andreas fault system. A 

significant proportion of this "missing" motion occurs in 

the Basin and Range, the seismicity of which plays an 

integral role in the tectonics of the plate boundary. 

EARTHQUAKE HISTORY OF 

THE SAN ANDREAS FAULT SYSTEM 

The historical record of major earthquakes affecting 

California, western Nevada, and northernmost Baja 

California (table 6.1) includes basic seismologic data on 

206 of the largest earthquakes occurring between 1769 

and 1989. This catalog lists all known events of M26 and 

includes new and updated information on their locations. 

The record of seismicity within the San Andreas fault 

system and surrounding regions is both geographically 

and temporally uneven and incomplete before the introduction of practical seismographic instrumentation 

around the turn of the 20th century. In general, the 

density and distribution of people who left written 

accounts of their experiences determines the reliability of 

the catalog during the preinstrumental period. From the 

establishment of the Franciscan missions beginning in 

1769 until their secularization in the 1830's, detailed 

accounts of events that damaged the missions are available, and these accounts form the primary source material for earthquakes occurring during this period. Life in 

California was a constant struggle for survival at that 

time; posting to a mission evidently was considered a 

hardship assignment, and so essentially nothing was 

recorded about events that were only felt. even when 

they were destructive at nearby missions. After secularization and before the gold rush, the quality of the record 

degrades with the cessation of the annual reports of the 

missions. Other sources of records also are notably weak 

during the Mexican period, from the early 1830's until 

1846. 

The discovery of gold in 1848 transformed the written 

record of earthquakes with the advent of newspapers 

throughout the gold fields in the Sierran foothills and in 

the San Francisco Bay region. Printed accounts of 

earthquakes have been extensively used, notably by 

Toppozada and others (1988), to quantify the seismicity of 

California from 1850 onward. They estimated that their 

historical catalog is probably complete for the San 

Francisco Bay region and central Sierra Nevada from 

1850 on for earthquakes of M-6. The same level of 

completeness is not achieved, however, for the San 

Andreas fault system in southern California until the 

1890's. Statewide, the catalog of earthquakes is substantially complete for earthquakes of M-7 after about 1850 

(see Agnew, 1985). The quality of the catalog for central 

Nevada, where much significant 20th century seismicity 

has occurred, is less complete. Questions remain today 

about purported events as late as 1903 in this region 

(Slemmons and others, 1959). 

Reports of the local effects of earthquakes continue to 

play a major role in determining the locations and sizes of 

earthquakes well into the 20th century. The earliest 

seismographs capable of systematically detecting California and Nevada earthquakes were installed throughout the world by John Milne beginning in 1896. 

Seismograms from these instruments and their successors provide useful instrumental magnitudes from 1898 

onward. However, not until the development of the 

Wood-Anderson seismograph and its deployment 

throughout California beginning in 1926 do instrumental 

measurements fully supplant noninstrumental magnitudes and epicentral locations. 

The objective in assembling a single catalog from these 

many sources, spanning many different types and qualities of information, has been to achieve uniform spatial 

coverage without sacrificing any events of historical 

significance. M=6 was chosen as the threshold magnitude because probably all events of this magnitude are 

known from the instrumental period beginning in 1898, 

and the preinstrumental record is reasonably complete at 

this level in some areas for an additional half-century. All 

earthquakes with at least one reported magnitude of at 

least 6.0 have been included in the catalog. Because 

magnitude is an estimated quantity and has some inherent uncertainty, events with reported magnitudes within 

a few tenths of a unit of 6.0 are also included. In addition 

to those earthquakes with cataloged magnitudes, original 

documents for others with reported high intensities or of 

particular historical significance have been reexamined in 

an attempt to refine their locations and magnitudes. 

A word of introduction should be added about earthquake locations and magnitude scales and their use in this 

chapter. Earthquakes are complex physical processes 

generated by sudden slip on faults, and as such they can 

only be grossly characterized by simple concepts. Two 

seismologic conventions are in common use for assigning 

a single geographic coordinate to an earthquake: One 

measures the center of energy release, frequently as 

estimated from the intensity distribution for preinstrumental events; the other measures the location of the 

initial point of rupture, or hypocenter, as determined 

from seismic traveltime measurements. Either point on 
156 THE SAM ANDREAS FAULT SYSTEM. CALIFORNIA 

the Earth's surface above the hypocenter or the center of 

the intensity distribution is sometimes referred to as the 

epicenter, and each type of location appears in table 6.1, 

with preference given to instrumental epicenters. Fortunately, the geographic differences between these distinct physical measures become significant only for the 

largest events, M=-7, when viewed at the scale of the 

entire San Andreas fault system. 

Magnitude, as commonly used to compare the sizes of 

different earthquakes, also represents an extreme simplification of the earthquake process and by itself cannot 

fully characterize the size of any event. Traditionally, 

seismologists have developed a suite of magnitude scales, 

each with its own purpose and range of validity to 

measure an earthquake. Because no single magnitude 

scale can be systematically applied to the entire historical 

record, a summary magnitude, M, is introduced here to 

facilitate comparisons between events. As described 

below in the subsection entitled "Quantification of Earthquakes and Magnitude Scales," M is taken as the 

surface-wave magnitude (Mg), when available, and as a 

modified intensity magnitude (M,) during the preinstmmental era. Generally speaking, M provides a better 

relative measure of the static, geologic increment of fault 

slip in the earthquake than it does of the severity of 

shaking. 

The earthquake history of California, western Nevada, 

and northern Baja California presented here has apparent limitations and can doubtlessly he improved through 

further research. Nevertheless, it providesa firm observational basis for assessing the tectonic implications of 

the 2-century-long seismic history, as well as of the 

prospects for future earthquake activity. 

PRINCIPAL EARTHQUAKES 

In this section, we briefly discuss some events of 

particular historical, social, or scientific significance. 

Although each of the 117 San Andreas fault system 

events in table 6.1 merits discussion, this task is far 

beyond the scope of this review, and so the reader is 

referred to the reports by Richter (1968), Cofiinan and 

others (1982). and Townley and Allen (1939) for an 

introduction to many of these events. Table 6.1 also omits 

several historically significant events with magnitudes 

well below the nominal threshold of M-6 adopted here, 

and so it something less than a complete reference on San 

Andreas seismicity. 

My major effort in constructing this catalog has gone 

into identifying and validating all reported events ol 

M^G. Two conspicuous omissions from table 6.1, events 

that are commonly mentioned in the literature but that 

could not be substantiated upon further inspection, 

should he noted. The first is the 1862 earthquake alleged 

Ã have ruptured the Big Pine fault (for example, 

lenninga, 1975). Toppozada and others (1981) failed to 

End any evidence supporting the occurrence of a major 

sarthquake at that time in the region. Geologic inspection 

if the surface trace of the fault by M.M. Clark (oral 

;ommun., 1988) similarly failed to provide evidence of 

my historical activity. The other deleted event appears 

m the seismicity map by Goter (1988) at lat 35O N., long 

126' W., with an epicenter from the catalog of Abe and 

Noguchi (1983). Although a large (MÃ§=6.8 earthquake 

sertainly took place on March 22, 1902, no evidence has 

been uncovered to support a location anywhere on shore 

in California or, for that matter, in the Western United 

States. The original location determined by Milne in 1903 

placed the event well off the California-Oregon coast at 

lat 42' N., long 130' W. 

JULY 28, 1769 (M-6) 

The earthquake history of California serendipitously 

begins with the first overland expedition through the 

State in 1769. In response to the perceived threat posed 

by Russian expansion into the northern Pacific and 

growing British presence in the northwestern Pacific, 

Spain embarked on the colonization of present-day Califomia through the establishment of a series of Franciscan 

missions, supported by military garrisons at San Diego 

and Monterey. In the summer of 1769, Gaspar de Portola 

led the first expedition from San Diego to establish a land 

mute to Monterey. 

On July 28, while camped along the Santa Ana River, 

about 60 km southeast of Los Angeles, a sharp earthquake was felt that "* * * lasted about half as long as an 

Ave Maria" (fig. 6.3). From the diaries of three members 

of the expedition, we know that earthquakes were felt on 

nearly a daily basis through August 3, as the party 

traveled northwestward to near San Gabriel and then 

westward across Los Angeles to the Pacific. The diary of 

Fray Juan Crespi (Bolton, 1927) mentions no fewer than 

a dozen aftershocks, some described as violent. After 

August 4, no further earthquakes were mentioned as the 

expedition traveled into the San Fernando Valley and 

exited to the north. 

These sketchy reports suggest that the explorers 

traveled near or through the epicentral area of a moderate earthquake (Richter, 1973; Toppozada and others, 

1981). Comparisons between the accounts of the aftershocks and more recent events suggest an event of 

similar size and location to the 1933 Long Beach, 1971 San 

Fernando, or 1987 Whittier Narrows earthquake. If 

significance is placed on the absence of aftershocks while 

crossing the source region of the 1971 San Fernando 

earthquake, the evidence would seem to favor a source in 

the Los Angeles Basin. An event on either the San 
6. EAETHQUAKE HISTORY, 1769-1989 157 

Andreas or San Jacinto faults, some 50 kin to the 

northeast, could conceivably have been the source of the 

1769 earthquake. The description of the duration of 

strong shaking, however, suggests a magnitude more of 

5-6 than of 7-8. 

A more distant source would make the long, felt 

aftershock sequence even more remarkable because it 

would he well removed from the expedition route. 

DECEMBER 8. 1812 (M-7) 

The first of two significant earthquakes to occur in 

southern California in 1812 occurred on December 8 and 

destroyed the church at Mission San Juan Capistrano, 

killing40 neophytes (fig. 6.4); damage was also sustained 

FIGURE 6.3.-Early accounts of significant earthquakes reflect the 

sparse settlement of California in anarrow coastal corridor before the 

population explosion accompanying the gold rush in 1849. Accounts of 

the few welldocumented events (dates shown) principally derive 

from mission records at San Diego (SD), San Luis Rey (SLR), San 

Juan Capistrano (SJC), San Gabriel (SG), San Fernando Rey (SFR), 

San Buenaventura (SBV), Santa Barbara (SB), Santa Inez @I), and 

La Purisima Concepcibn (LPC), and from the towns of Los Angeles 

(LA) and Fort Tejon (FT) in southern California. Accounts from the 

Spanish capital Monterey (M), SanFrancisco (SF), and San Jose (SJ), as well as mission sources, detail events in north half of the State. 

Uncertainties in the interpretation of every event before the peat 

earthquake of 1857 (rupture shown: arrows indicate direction of 

relative movement) are welt illustrated by newly uncovered evidence 

suggesting a San Andreas origin for the December 8, 1812, shock 

near Wrightwood (head of connecting arrow), well inland of traditional location along the coastal Newport-lnglewod fault (tail of 

arrow). Earthquake of December 21, 1812, locates in the Santa 

Barbara Channel (SBC). Foreshodni of the great earthquake of 1857 

locate near Parkfield (P), suggesting unilateral rupture propagation 

to the southeast. 

at San Gabriel. The accounts of this earthquake and the 

later one on December 21 cannot be readily disentangled 

at San Fernando Rey and at San Buenaventura, considerably complicating the interpretation of this event. 

Analyses of these scanty data by Toppozadaand others 

(1981) and Evernden and Thompson (1985) place the 

epicenter along the south half of the Newport-Inglewood 

fault zone (fig. 6.3). This location is somewhat constrained by the interpretation of no damage at Buenaventura during the event. The Los Angeles Star of January 

10, 1857, however, stated that the December 8 event 

severely damaged the church tower (Agnew and Sieh, 

1978). The same story attributed the collapse of the stone 

arch roof of the church at San Juan Capistrano to poor 

construction, a possibility made credible by the death of 

the master mason before completion of the church (fig. 

6.3; Duncan Agnew, oral commun., 1988). 

Recently, Jacoby and others (1988) proposed that this 

event ruptured the San Andreas fault at Wrightwood 

(fig. 6.3), on the basis of dendrochronologic dating of 

distress to trees growing on the fault trace. Sieh and 

others (1989) argued that this rupture extended at least 

25 km northwestward into the peat bog at Pallet Creek. 

The fault rupture in this event preserved at Pallet Creek 

is comparable in size to the rupture formed in the 1857 

earthquake. 

The preferred location of the December 8, 1812, 

earthquake on the San Andreas fault as proposed by 

Jacoby and others appears in table 6.1. A magnitude of 

about 7 is consistent with the inferred extent of damage. 

The lateral extent of rupture is unconstrained to the 

southeast and may well have extended into the San 

Bernardino Valley. However, the accounts of the earthquake from Indians living in the San Bernardino Valley 

that were thought to place some constraint on the 

rupture are now believed to be fictitious (Harley, 1988), 

leaving Mission San Gabriel, some 40 km from the 

rupture, as the nearest point of observation. 

DECEMBER 21, 1812 (A(=?) 

The second major episode of earthquake activity in 

1812 damaged the missions along the Santa Barbara 

Channel and western Transverse Ranges just 13 days 

later, on December 21 (fig. 6.3). All investigators place 

this event in the Santa Barbara Channel and assign a 

magnitude of about 7 (see Toppozada and others, 1981, 

and Evernden and Thompson, 1985, for two recent 

analyses). This sequence appears to have involved two 

events of comparable magnitude separated in time by 

about 15 minutes. A vigorous aftershock sequence accompanied the earthquakes and lasted until the end of the 

year at Mission Santa Barbara and Mission La Purisima 

Concepci6n. Reports of a tsunami appear to be exagger- 
158 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

ated, although some kind of wave activity probably 

accompanied the earthquake (Toppozada and others, 

1981; McCulloch, 1985). 

JUNE 10, 1836 (M=f^i) 

Little is known about the strong earthquake of June 10, 

1836, that struck the then lightly populated San Francisco Bay region. An account of the event, published in the 

aftermath of the 1868 earthquake, provides the principal 

rationale for associating this event with the Hayward 

fault. Louderback (1947) systematically compared the 

two events and concluded that the 1836 earthquake 

probably ruptured the Hayward fault. Lindh (1983) 

proposed that the 1836 event ruptured the north half of 

the fault, whereas the 1868 event is known to have 

ruptured the south half, thereby avoiding the paradox of 

two large events on the same segment separated by a 

scant 32 years. 

JUNE 1838 (M=7) 

The pioneering historical work of Louderback (1947) 

reveals that a major earthquake with probable rupture of 

the San Andreas fault occurred in June lEB8. Documentation of the event is so poor that its date cannot be fixed 

more precisely than "late June." Louderback concluded 

that the shock was comparable in magnitude to the 1906 

earthquake. Current opinion suggests a smaller event 

involving only the 60+-km-long segment of the fault on 

the San Francisco peninsula (Working Group on Earthquake Probabilities, 1988). 

JANUARY 9, 1857 (M=8'1<) 

The great Fort Tejon earthquake of January 9, 1857, 

ruptured 300 km of the San Andreas fault from near 

Parkfield to Wrightwood and offset the fault by as much 

as 9% m on the Carrizo Plain. The fault rupture and the 

effects of the earthquake have been extensively studied, 

notably by Agnew and Sieh (1978) and Sieh (1978b). The 

epicenter of this event appears to have been at the 

extreme northwest end of the fault rupture, as determined by the intensity patterns of two M=6 foreshocks 

centered near Parkfield (Sieh, 1978a). Strong shaking 

lasted from 1 to 3 minutes, consistent with unilateral 

rupture propagation to the southeast (fig. 6.3). 

The earthquake caused only two deaths in the sparsely 

settled southern California region. Damage was most 

severe along the fault zone; nearly every building sustained damage at Fort Tejon. In Los Angeles, then a city 

of about 4,000 people located approximately 60 km from 

the fault, some houses were cracked, hut none were 

severely damaged (Agnew and Sieh, 1978). Modified 

Mercalli intensities (MMI's) of VII or more occurred in 

the San Fernando Valley, San Gabriel Valley, and 

Ventura region. 

It is natural to compare the 1857 and 1906 earthquakes, 

the two greatest earthquakes of the San Andreas fault in 

historical time. The 1906 fault break was longer, whereas 

maximum and average surface offsets were larger in 

1857. These differences approximately balance each other, and so the seismic moments of the two events are 

approximately equal. Moment magnitudes computed using comparable data are M=7.8 for the 1857 earthquake 

FIGURE 6.4.-Mission San Juan Capistrano as drawn by Henry Miller in 1856.43 years after the December 8, 1812, earthquake. Vaulted stone 

church at right collapsed in that earthquake, killing 40 worshipers. Photograph courtesy of the Bancroft Library. 
6. EARTHQUAKE HISTORY, 1769-1989 159 

and M=7.7 for the 1906 event. A summary magnitude ol 

M=8V4 was assigned by analogy with the 1906 earthquake. 

OCTOBER 21, 1868 (M=i) 

Known as the "great San Francisco earthquake" until 

1906, one of California's most destructive earthquake? 

occurred on October 12, 1868, resulting from slip on the 

Haywad fault. Heavy damage occurred in communities 

situated along the fault and in San Francisco and San Jose 

(fig. 6.5). Sadly, many of the engineering lessons learned 

from this earthquake and openly discussed at the time, 

such as the hazards of building on "made ground" 

reclaimed from the San Francisco Bay or the admonition 

to "build no more cornices,"were long forgotten by the 

time of the 1906 earthquake. 

FEBRUARY 24, 1892 (M=7) 

The strong earthquake of February 24, 1892, located 

near the United States-Mexican border was assigned to 

the Agua Caliente fault north of the border by Toppozada 

and others (1981) and to the Laguna Salada fault in Baja 

California by Strand (1980). The literature on earthquakes in Baja California contains numerous references 

to this earthquake as having originated near the Agua 

Blanca fault, about 100 km southwest of Strand's epicenter (for example, Richter, 1958). The two recent intensity 

maps clearly rule out this epicenter and place it on the 

southern section of the Elsinore fault system. 

APRIL 19 AND 21. 1892 (M=6'/2 AND 6'14) 

A pair of strong earthquakes rocked the west side of 

the Sacramento Valley on April 19 and 21, 1892, heavily 

damaging the towns of Vacaville, Dixon, and Winters. 

The frst shock was stronger and caused heavy damage at 

Vacaville; the aftershock was more severe at Winters. 

The earthquakes are reminiscent of the 1983 Coalinga, 

Calif., earthquake, in that both sequences were positioned along the western margin of the Central Valley. 

Focal mechanisms of small earthquakes located along this 

boundary zone show numerous examples of low-anglethrust focal-mechanism solutions of similar orientation to 

the Coalinga earthquake, in addition to strike-slip mechanisms (see chap. 5; Wong and others, 1988), suggesting 

the possibility of a similar mechanism for these 1892 

earthquakes. 

DECEMBER 25, 1899 (M=6A) 

Heavy damage occurred in the towns of San Jacinto 

and Hemet, located along the San Jacinto fault, from an 

earthquake on Christmas Day 1899. Six fatalities were 

attributed to the earthquake. 

APRIL 18, 1906 (M=81/4) 

The California earthquake of April 18, 1906, ranks as 

one of the most significant earthquakes of all time. 

Today, its importance comes more from the wealth of 

scientific knowledge derived from it than from its sheer 

size. Rupturing the northernmost 430 km of the San 

Andreas fault from northwest of San Juan Bautista to the 

triple junction at Cape Mendocino (fig. 6.6), the earthquake confounded contemporary geologists with its 

large, horizontal displacements and great rupture length. 

Indeed, the significance of the fault and recognition of its 

large cumulative offset would not be fully appreciated 

until the advent of plate tectonics more than half a 

century later. Analysis of the 1906 displacements and 

strain in the surrounding crust led Reid (1910) to 

formulate his elastic-rebound theory of the earthquake 

source, which remains today the principal model of the 

earthquake cycle. 

As a basic reference about the earthquake and the 

damage it caused, geologic observations of the fault 

rupture and shaking effects, and other consequences of 

the earthquake, Lawson's (1908) report remains the 

authoritative work, as well as arguably the most important study of a single earthquake. In the public's mind, 

this earthquake is perhaps remembered most for the fire 

it spawned in San Francisco, giving it the somewhat 

misleading appellation of the "San Francisco earthquake" 

(fig. 6.7). Shaking damage, however, was equally severe 

in many other places along the fault rupture. The 

frequently quoted value of 700 deaths caused by the 

earthquake and fire is now believed to underestimate the 

total loss of life by a factor of 3 or 4. Most of the fatalities 

occurred in San Francisco, and 189 were reported elsewhere. 

At almost precisely 5:12 a.m. local time, a foreshock 

occurred with sufficient force to be felt widely throughout the San Francisco Bay area. The great earthquake 

broke loose some 20 to 25 s later, with an epicenter near 

San Francisco (Bolt, 1968; Boore, 1977). Violent shocks 

punctuated the strong shaking, which lasted some 45 to 

60 s. The earthquake was felt from southern Oregon to 

south of Los Angeles and inland as far as central Nevada 

(fig. 6.6). The highest MMI's of VII to IX paralleled the 

length of the rupture, extending as far as 80 km inland 

from the fault trace. One important characteristic of the 

shaking intensity noted in Lawson's (1908) report was the 

clear correlation of intensity with underlying geologic 

conditions. Areas situated in sediment-filled valleys sustained stronger shaking than nearby bedrock sites, and 

the strongest shaking occurred in areas where ground 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

FIGURE 6.5.-San Francisco Mmiw Chmicle of October 28, 1868, richly illustrates severe damage sustained by 

buildings of noor desim or located on filled land durino-earthauake on the Haward fault. Reduction of this fi@re 
6. EARTHQUAKE H: 

reclaimed from San Francisco Bay failed in the earthquake. Modern seismic-zonation practice accounts for the 

differences in seismic hazard posed by varying geologic 

conditions (see Borcherdt, 1975, and Ziony, 1985, for 

analyses of the San Francisco Bay and Los Angeles 

regions, respectively). 

The characteristics and amount of surface fault slip in 

this earthquake varied to a remarkable degree along the 

length of the rupture. Peak displacements of 6 m were 

measured near Olema on the Point Reyes peninsula, 

where the surface trace of the rupture formed a sharp, 

well-defined break (fig. 6.8). In contrast, the fault break 

was extremely difficult to recognize along its southemmost 90 km, where the surface offset averaged only about 

1'12 m or less (see chap. 7). 

The magnitude of 8.3 commonly quoted for the 1906 

earthquake comes from Richter (1958) and, within the 

precision of reporting, is identical to the 8% listed by 

Gutenberg and Richter (1954). Table 6.1 also lists other 

magnitudes for this earthquake, derived from recent 

analyses of both the same data used by Gutenberg and 

Richter and new data. Strictly speaking, a "Richter 

magnitude" (ML) for the earthquake cannot be determined because no appropriate seismographs were in 

FIGURE 6.6.-California earthquake of 1906 showing extent of fault 

rupture along the San Andreas fault, location of epicenter near San 

Francisco, maximum extent of structural damage, and limits of 

perceptionof shock. Modified from Lawson (1908) and Toppozadaand 

Purke (1982). 

[STORY. 1769-1989 161 

operation at the time. Jennings and Kanamori (1979) used 

related measurements extracted from simple pendulums 

at Yountville, Calif., and Carson City, Nev., to derive 

ML=6.9, substantially smaller than the traditionally 

quoted value. ML, which is based on the single largest 

peak on a seismogram at approximately 1-s period and 

takes into account neither the duration of the event nor 

longer period motions, is saturated for this event. 

Geller and Kanamori (1977) used the unpublished 

worksheets of Gutenberg and Richter to compute a 

body-wave magnitude of mb=7.4, using the procedure of 

Gutenberg and Richter (1956). Because long-period (14 s) 

P-waves were used in this calculation, it cannot be 

directly compared to the short-period mb values routinely 

reported today. 

Other workers since Gutenberg and Richter have 

studied the long-period surface waves of the 1906 earthquake and computed Ms values. Bolt (1968) confirmed an 

Ms of about 8'14, whereas Lienkaemper (1984) found Ms= 

8.3 from an analysis of all the records collected by Reid 

(1910). Lienkaemper's magnitude combined data from 

both damped and undamped instruments, correcting each 

for magnification at the appropriate period of motion. 

Abe (1988), who analyzed only the undamped Milne 

seismograms, obtained Ms= 7.8, using slightly different 

procedures and a systematic set of station-magnitude 

corrections. Also, the four damped seismometers (all in 

Europe) give Ms=8.1. Longer period (50-100 s) surface 

waves analyzed by Thatcher (1975) indicate a seismic 

moment of 4x10m dyne-cm, equivalent to M=7.7, in 

agreement with the seismic moment of 5x1027 dyne-cm 

obtained from geodetic data, thus giving M~7.8 (Thatcher and Lisowski, 1987). Finally, Toppozada and Parke 

(1982) assigned an intensity magnitude (MI) of 7.8 on the 

basis of the total area (48,000 km2) undergoing shaking of 

MMI VII or higher. 

The "traditional" magnitude of 8% is retained here, 

except where seismic moment is used for quantitative 

purposes. 

NOVEMBER 21, 1915 (M=7.1) 

The major earthquake of November 21, 1915, triggered a spectacular steam eruption of a mud volcano, 

creating a 100+-m crater in Volcano Lake, Baja California, near the north terminus of the Cerro Prieto fault. 

Extensive cracking of the levee around the lake was 

noted at the time of the shock, but no tectonic ground 

displacements were found (Seismological Society of 

America Bulletin, 1916). This event may well be related 

to the November 29, 1852, earthquake (M=6'/2Â±) which 

also triggered a mud-volcano eruption at Volcano Lake 

that was observed at Fort Yuma, Ariz. Each of these 

events was probably associated with the Cerro Prieto 

fault. 
162 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

APRIL 21, 1918 (Af-6.9) 

The communities of Hemet and San Jacinto were 

severely damaged for the second time in 19 years by the 

large earthquake of April 21, 1918, on the San Jacinto 

fault. Both the 1899 and 1918 earthquakes produced 

similar intensity patterns throughout the southern California region, and these two events have been compared 

to each other. However, surface waves on Milne seismograms at common stations (Victoria, British Colombia, 

and Toronto, Ontario, Canada; San Fernando, Spain) 

average 3 times larger for the 1918 earthquake, corresponding to a difference in Mr of 'I2 unit. As with the 1836 

and 1868 earthquakes on the Hayward fault, the relation 

between the rupture zones in these two events is unclear. 

Surprisingly, no surface rupture was found for an event 

of this size, despite a specific search for it. 

NOVEMBER 4, 1927 (Af=7.3) 

The Lompoc earthquake of November 4, 1927, is the 

largest known event in the San Andreas system west of 

the San Andreas fault proper. This event produced a 

tsunami with local runup heights of 1.5 to 1.8 m (McCulloch, 1985). The exact location of the earthquake and its 

association with any causative structure remain the 

subject of a spirited debate (Gawthrop, 1978, 1981; 

Hanks 1979, 1981). 

MARCH 11, 1-3 (At=6.3) 

Rupture of the Newport-Inglewood fault on March 11, 

1933, caused major damage and a loss of 115 lives in Long 

Beach and surrounding parts of the Los Angeles Basin. 

Structural damage to public schools was particularly 

serious, and had the event occurred when schools were in 

session, the calamity would have been far worse. The 

Field Act, mandating construction standards for schools 

in California, was enacted as a consequence of the 

earthquake. 

DECEMBER 30 AND 31, 1934 (A4=6.5 AND 7.0) 

The major sequence that occurred along the Cerro 

Prieto fault on December 30 and 31, 1934, appears to 

FIGURE 6.7.-San Francism. Calif.. on the moho- of Andl 18. 1x16. nhotrtmnh-t; taken of the citv before fire sweet through show no 

sustained heavy damage in the earthquake, this and many other Graphic Arts. 
6. EARTHQUAKE HISTORY, 1769-1989 163 

have ruptured the surface trace of the fault near where it 

enters the Gulf of California. Aerial photographs of the 

fault crossing a tidal flat taken in 1935 show very fresh 

appearing fault morphology; subsequent photographs 

display a substantially subdued morphology (Kovach and 

others, 1962). 

MAY 19. 1940 (M=7.1) 

The Imperial fault was discovered from its 60+- 

km-long rupture in the Imperial Valley earthquake of 

May 19, 1940. Faulting was predominantly right-lateral 

strike slip and attained a peak offset of more than 6 m at 

the United States-Mexican border (fig. 6.9). The first 

instrumental measurement of strong ground motion 

adjacent to a fault rupture was obtained from an accelerograph located about 7 km from the surface trace. This 

record, which provides clear evidence of irregular seismic-energy release during the course of the event (Trifunac and Brune, 19701, has played a major role in 

shaping building codes for earthquake-resistant design. 

FIGURE 6.8.-Trace of lii06 earthquake rupture near point of maximum offeet (6 m) near Olema on the Point Reyes peninsula north of 

San Francisco. Photograph by G.K. Gilbert. View northwestward. 

JULY 21, 1952 (M=7.7) 

The Kern County or Arvin-Tehachapi earthquake of 

July 21, 1952, ruptured the White Wolf fault in the 

largest event to strike California since 1906. The earthquake led to 12 fatalities, and 2 more occurred during a 

large aftershock on August 22. Field studies of the 

earthquake (Oakeshott, 1955) describe the geologic, 

seismologic, and engineering aspects of the earthquake. 

From a tectonic standpoint, this event is notable for its 

conjugate relation to the San Andreas fault. Left-lateral 

slip with a significant reverse-slip component occurred on 

the northeast-striking, southdipping fault plane. 

FIGL'RE 6.9.-Surface faultingin lii40 Imperial Valley earthquake offset (6 m) rcguliir rows of orange trees. Fault dkpliicement alongthis section 

of the Imperial fault was confined to a narrow zone. 
164 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

FEBRUARY 9. 1956 (hf-6.8) 

More than 19 krn of the hitherto-unknown San Miguel 

fault in Baja California ruptured in the earthquake 

sequence of February 9, 1956. The fault offset was 

consistently right lateral and up to the northwest, and 

attained maximum horizontal and vertical separations of 

78 and 91 cm, respectively (Shor and Roberts, 1958). The 

sequence contained numerous aftershocks, including 

three of M26. About 2 years earlier, a pair of M=6 

events that occurred to the south and west of the San 

Miguel fault may have been associated with the Agua 

Blanca fault. 

APRIL 9, 1968 (At=6.5) 

The Borrego Mountain earthquake of April 9, 1968, 

produced the first documented rupture of the San Jacinto 

fault system when right-lateral displacements of nearly 

0.4 m &curred along a 30-km-lone &gnient of the ~oyot-e 

Creek fault. The Y.S. Geological Survey (1972) published 

a detailed description of the event. 

FEBRUARY 9, 1971 (hf-6.5) 

The San Fernando earthquake of February 9, 1971, 

ranks as one of the most serious California earthquakes in 

historical time. The event claimed 58 lives and caused 

more than half a billion dollars in property damage, 

including the destruction of two hospitals, a freeway 

interchange, and the Van Norman Dam. The earthquake 

ruptured north-dipping, high-angle reverse faults beneath the southern margin of the San Gabriel Mountains 

and broke the surface along a discontinuous, 15-km-long 

zone. Surface displacements averaged about 1 m. Seismograms of the earthquake reveal a steeply dipping deep 

fault and a more shallowly dipping near-surface fault 

(Langston, 1978; Heaton, 1982). Numerous publications 

report on detailed investigations of this event, including 

the summary report published by the U.S. Geological 

Survey (1971). 

OCTOBER 15, 1979 (hf=6.5) 

The Imperial fault ruptured for the second time in less 

than 40 years in a major surface-faulting earthquake on 

October 16, 1979 (U.S. Geological Survey, 1982). The 

event broke the north 30 km of the fault, or approximately half the length of the 1940 fault break. However, it was 

clearly much smaller than the earlier event; maximum 

surface offsets were well under 1 m, in contrast to 6 m 

observed in 1940, and the seismic moment was smaller by 

nearly an order of magnitude. Within the zone of overlapping surface rupture, the two events display nearly 

identical displacement profiles (Sharp, 1982), suggesting 

that the 1979 earthquake represents a characteristic 

rupture of this segment of the fault. Strong-groundmotion records for the 1979 earthquake form an unparalleled suite of near-field recordings and have stimulated 

numerous investigations into the dynamics of the source. 

MAY 2. 1983 (At-6.5) 

Our understanding of the nature of the earthquake 

hazard posed by active faults in the San Andreas fault 

system was fundamentally altered by the occurrence of 

the Coalinga earthquake of May 2, 1983, on a low-angle 

thrust fault deep beneath the western margin of the San 

Joaquin Valley (Rymer and Ellsworth, 1990). Before this 

event, it had been thought that the major, seismically 

active faults in California could be recognized on the basis 

of their surface exposures and record of late Quaternary 

activity. However, no surface expression exists for the 

fault system responsible for either this event or the 

M=5.9 North Kettleman Hills earthquake of August 4, 

1985, that adjoins it to the southeast. Instead of a surface 

fault, the buried deformation is expressed at the surface 

by active folds (the Coalinga anticline and the Kettleman 

Hills) that grew during the earthquakes (Stein and King, 

1984). 

The orientation of the fault and the style of movement 

on it present another major challenge to prevailing 

models of the San Andreas system, because this earthquake resulted from a release of compressive forces 

oriented nearly perpendicular to the trace of the San 

hdreas fault. Accumulating evidence on the orientation 

of -- the ~~~- stress ~~-~ ~ field ~ astride the San Androas fault sueersts -- 

that only a small component of the total stress acts to 

accommodate the plate motion along the San Andreas 

fault itself (Mount and Suppe, 1987; Zoback and others, 

1987). 

NOVEMBER 24, 1987 (Ak6.6) 

The Superstition Hills fault ruptured in its entirety on 

November 24, 1987. The total amount of separation 

substantially increased by persistent afterslip in the 

months after the main shock: in fact. the rate of afterslip 

was so great on the south half of the surface break as to 

leave open the possibility that all of its displacement 

occurred as afterslip. The earthquake was preceded by a 

major foreshock (M=6.2), on a conjugate, northeasttrending, left-lateral strike-slip fault that intersected the 

Superstition Hills fault at the main-shock epicenter. The 

surface-faulting pattern of the entire sequence was 

particularly remarkable for the occurrence of numerous 

breaks on other conjugate faults in the north quadrant 

around the main break (see Hanks and Allen, 1989). 
6. EARTHQUAKE HISTORY, 1769-1989 165 

OCTOBER 18, 1989 (M-7.1) 

In the late afternoon of October 17, 1989, the San 

Andreas fault ruptured in its first major earthquake since 

1906 at 5:04 p.m. P.d.t. (0004 G.m.t. on Oct. 18). 

Centered along a remote segment of the fault in the 

southern Santa Cruz Mountains, the Loma Prieta earthquake reruptured the southernmost 40 km of the 1906 

fault break, producing the Nation's most costly natural 

disaster. The earthquake claimed 62 lives and injured an 

additional 3,757 people. It destroyed 963 homes and 

damaged more than 18,000 others, displacing 12,000 

people from their residences. The combined dollar loss to 

the private and public sectors exceeded $6 billion (Plafker 

and Galloway, 1989). 

Damage in the epicentral region was most severe 

where the earthquake shaking was compounded by local 

ground failures, commonly involving landslide movement 

but also including some fractures of probable tectonic 

origin; the shaking clearly reactivated some fissures 

observed in 1906. Primary fault displacement, however, 

did not reach the surface. In the hard-hit communities of 

Santa Cruz, Watsonville, and Los Gatos, unreinforcedmasonry buildings bore the brunt to the damage, and 

ground conditions played a significant role in the damage 

patterns. 

The earthquake also caused grave damage and claimed 

the greatest number of lives far to the north, in San 

Francisco and Oakland, about 100 km from the epicenter. 

There, the earthquake selectively destroyed structures 

known to be at risk or located on poor ground (Plafker 

and Galloway, 1989). The root cause of the devastation in 

the Marina District of San Francisco (fig. 6.101, as well as 

at most other sites along the margin of the San Francisco 

Bay, was liquefaction-induced ground failure. All of these 

localities sit on land reclaimed from the bay and are 

underlain by young, water-saturated sedimentary deposits. As we know from the clear lessons of history, 

provided by the earthquakes of 1865, 1868, and 1906 

(Lawson, 1908), such materials perform poorly even 

under modest levels of earthquake shaking. The collapse 

of the double-decked section of California Interstate 

Highway 880 in Oakland (fig. 6.10), where 41 people died, 

resulted principally from design defects. The section of 

the viaduct that collapsed was founded on soft estuarine 

sedimentary deposits that amplified the strong ground 

motion; the adjoining section, founded on alluvium, rode 

through the earthquake. 

The earthquake broke the San Andreas fault where it 

makes a conspicuous leftward bend, connecting straighter subparallel segments to the north and south. The fault 

plane dips 70' SW., and movement in the earthquake 

involved comparable amounts of right-lateral strike slip 

and reverse slip, a kinematic response driven by the need 

to remove material from this compressional fault bend as 

the Pacific plate moves to the northwest around it. The 

rupture nucleated at the base of the seismic zone, at 

18-km depth, and spread unilaterally upward and bilaterally along strike, filling a conspicuous void in the 

preevent seismicity. Geodetic data collected immediately 

after the event suggest an average strike-slip displacement of 1.6 m and an average reverse-slip displacement 

of 1.2 m, rising from the hypocenter at 18 km to within 6 

krn of the surface. 

The Loma Prieta earthquake fulfilled a long-term 

Forecast for the rupture of this specific segment of the 

San Audreas fault (Lindh, 1983; Sykes and Nishenko, 

FICUHL 6.10.-Damage in October 18, 1989, Loma Prieta earthquake 

red at distances as far as 100 km from the epicenter in areas 

underlain by water-saturated, unconsolidated material. A, Liquefaction-induced ground failure in the Marina district of San Francisco 

(top) was restricted to land reclaimed from the San Francisco Bay. B, 

In Oakland. the second deckof Interstate Hichwav880 colla~sed onto 

the shaking. 
166 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

1984; Working Group on California Earthquake Probabilities, 1988). The high earthquake potential assigned to 

this segment stemmed from its behavior in the 1906 

earthquake, when the fault displacement, as measured at 

the surface, averaged about 1.5 m, far less than the 

average for the entire rupture. Estimates of the longterm slip rate along this segment of the San Andreas fault 

suggested that the strain released in the 1906 earthquake 

would be renewed in 75 to 136 years, implying that 

another earthquake was possible in the coming decades. 

With its occurrence, the Loma Prieta earthquake became 

the second event in 2 years to fill arecognized seismic gap 

along the San Andreas; the first was the 1987 Superstition Hills earthquake. The Loma F'rieta earthquake also 

represents the third historical rupture of this segment of 

the San Andreas fault; the first was the October 8, 1865, 

earthquake, nominally assigned M=6'/z, which also 

caused liquefaction-induced ground failure in San Francisco. 

SEISMICITY OF 

THE WESTERN BASIN AND RANGE PROVINCE 

The advent of plate tectonics and its application to 

western North America by Atwater (1970) provided a 

unifying framework for the contemporary tectonics of the 

western Basin and Range and its interaction with the San 

Andreas fault system to the west. Deformation withm 

the province reflects soft coupling of the San Andreas 

fault system to the North American craton and distribution of the relative plate motion-by mechanisms yet 

unknown-well over 500 km into the continental interior. 

The ubiquity of normal-fault-bounded ranges throughout 

the province tends to belie the underlying nature of 

present-day deformation within the region. Within historical time, this region has undergone nearly equal 

proportions of extension on normal faults and dextral 

shear on strike-slip faults. 

The earthquake history of the western Basin and 

Range province is poorly known before the instrumental 

period, owing to sparse settlement of this highdesert 

region. The deficiencies of this record are illustrated by 

the uncertainties associated with fresh-appearing fault 

scarps discovered in 1911 near the north end of what 

would become the rupture zone of the 1954 Fail-view 

Peak earthquake. Upon reviewing the scant historical 

evidence, Slemmons and others (1959) concluded that 

these scarps formed about 1903. The absence of an event 

of sufficient size in the instrumental record, however, 

suggests that the scarp forming event is older (or 

substantially smaller than M=6). Current understanding 

of 19th-century seismicity includes an episode of activity 

along the ~alifomia-~rv&ka State line, including a 

able rupture of the Olinghouse fault on December 27, 

1869 (Sanders and Slemmons, 1979), although this conclusion was questioned by Toppozada and others (1981). 

Surface faulting has accompanied numerous earthquakes in the region, the most significant of which are 

discussed below. Notable additional surface-faulting 

events include the M=6.3 Excelsior Mountain, Nev., 

earthquake of 1934 and the Mr=5.6 Fort Sage Mountain 

earthquake of December 14, 1950, located in northeastern California (Gianella, 1957). Ground rupture may have 

also accompanied the M=6 earthquake of January 24, 

1875 (see Gianella, 1957). If so, this observation would 

move the epicenter listed in table 6.1 to lat 39-14' N., long 

120'/2Â w. 

PRINCIPAL EARTHQUAKES 

MARCH 26, 1872 (M-7.6) 

The town of Lone Pine, Calif., was virtually leveled 

when the entire 100 to 110-km length of the Owens Valley 

fault ruptured on March 26, 1872, in one of the largest 

earthquakes in U.S. history. This fault, which lies in the 

middle of Owens Valley, is distinct from the normal faults 

bounding the front of the Sierra Nevada to the west. 

Considerable confusion has existed in the literature 

regarding the style of faulting in the 1872 earthquake, 

including interpretations of right-lateral, left-lateral, and 

normal-fault movement. A recent study of the earthquake offsets by Beanland and Clark (in press) unambiguously demonstrates that fault movement was 

predominately right-lateral strike slip, with an average 

horizontal displacement of 6 m (fig. 6.10). The vertical 

offsets were clearly smaller and averaged about 1 m down 

to the east. Beanland and Clark estimate a moment 

magnitude of M=7.5-7.7. Faulting in 1872 largely reactivated earlier Holocene scarps, as recognized by G.K. 

Gilbert when he visited the area in 1883. 

The event was felt throughout most of California and 

Nevada, and as far east as Salt Lake City, Utah. Adobe 

and brick buildings in Owens Valley sustained the brunt 

of the damage. Minor damage also occurred in the San 

Joaquin and Sacramento Valleys on the opposite side of 

the Sierra Nevada, at distances of as far as 400 km. In 

Yosemite Valley, John Muir witnessed a spectacular 

rockfall triggered by the earthquake. As severe as the 

ground shakingmust have been, it was noted in the Inyo, 

Calif., Indqw&nt of April 6, 1982, "* * * that not a 

person would have been killed or hurt had their houses all 

been made of wood." It is of some historical interest that 

the first long-term earthquake forecast, made by G.K. 

Gilbert in 1883 to the citizens of Salt Lake City, was 

based in part on his observations of the 1872 earthquake. 

In it, he noted that the rebuilding of Independence with 

wood-frame buildings was an extravagance, because this 
6. EARTHQUAKE HISTORY. 1769-1989 167 

great shock had relieved the accumulated strain, and so 

many generations would pass before conditions would 

permit another similar shock to occur (Gilbert, 1884): 

The old maxim, "Lightning never strikes in the same spot twice" is 

unsound in theory and false in fact; but something similar might truly 

be said about earthquakes. The spot which is the focus of an earthquake 

(of the type here discussed 11872 Owens Valley!) is thereby exempted 

for a long time. 

Many comparisons have been drawn between the 

Owens Valley earthquake and the great San Andreas 

earthquakes of 1857 and 1906. The size of the regions 

shaken in all three events are comparable, as are the 

maximum fault displacements. The two San Andreas 

events have significantly longer rupture lengths, and 

their seismic moments are larger by a factor of 2 to 3. 

Whether or not any or all of these earthquakes can be 

classified as "great" earthquakes becomes a question of 

semantics. All of them can be classified as great on the 

basis of their rupture lengths of 100 km or more (Kanamori, 1977), but they all have seismic moments more 

than 100 times smaller than the largest known earthquakes, such as the M=9.2 Alaska earthquake of 1964. 

Practically speaking, these events are among the largest 

known strike-slip events, and they must be close to the 

size of the largest possible strike-slip events along the 

San Andreas fault system. 

OCTOBER 3. 1915 (M-7.3) 

The 1915 Pleasant VaUey, Nev., earthquake of October 

15, 1916, created a series of spectacular normal-fault 

scarps in the central Nevada seismic zone of the Basin 

and Range province (figs. 6.11.6.12). Four major scarps 

formed during the earthquake, with an aggregate length 

of 59 km, and reruptured Holocene scarps at the base of 

the mountain blocks (Wallace, 1984). Fault movement in 

the earthquake appears to have been purely dip slip and 

averaged about 2 m on the steeply dipping fault plane. 

The earthquake was felt from western Utah to the Pacific 

coast and from northeastern Oregon to the United 

States-Mexican border. Instrumental measures of the 

magnitude range from 7.3 to 7% and exceed the moment 

magnitude of 7.2 derived from field measurements 

(Mo=6.1x102' dyne-an). 

The Pleasant VaUey earthquake lies at the north end of 

a 500-km-long belt of historical surface-faulting earthquakes within the central Nevada seismic zone and 

Owens Valley fault system. The four major earthquake 

sequences in this zone since 1872 leave two conspicuous 

seismic gaps that have been discussed as the potential loci 

of future major earthquake activity (fig. 6.11; Wallace, 

1984). 

DECEMBER 21, 1932 (Mz7.2) 

The second historical surface-faulting event in the 

central Nevada seismic zone on December 21, 1932, 

produced a discontimuous zone of surface faulting and 

fissures in the valleys west and north of Cedar Mountain 

(Gianella and Callaghan, 1984). Within the 60-km-long, 

north-northwest-trending zone where faulting was observed, most breaks struck east of north and showed 

clear evidence of right-lateral displacements (fig. 6.11). 

JULY 6, 1954 (M-6.6). AND AUGUST 24. 1954 (M-6.8) 

The Rainbow Mountain earthquakes of July 6 and 

August 24, 1964, produced a zone of east-facing normalfault scarps along the base of Rainbow Mountain, extendEXPLANATION 1915 

Pleasant Magnitude 

Valley > 5.5 

- * ' 01954 '/ Dixie Valley 

A 

1954 .\ \ Fairview Peak 

1932 

Cedar Mountain 

'. - 

KILOMETERS 

FIGURE 6.11.-California-Nevada region, showing locations of major 

historical earthquakes in the western Basin and Range province, 

1857-1989. Focal mechanisms of five largest events in lowei'-hemisphere projection show compressional quadrant shaded and indicate 

significant shear as well as extensional strain in province. Seismic 

gaps (labeled) are potential loci of future major earthquake activity 

(Wallace. 1984). 
168 THE SAN ANDREAS FACLT SYSTEM, CALIFORNIA 

ing northward into the Carson Sink. The July 6 event 

produced an 18-km-long surface rupture at the south end 

of this zone striking N. 15' E., with maximum displacements of about 30 cm. The August 24 shock extended the 

zone by 22 km in a N. 20Â E. direction, with as much as 

75 cm of normal-fault slip. Tocher (1956) noted that the 

displacement on the northern part of the July 6 break 

approximately doubled in amplitude between July 16 and 

September 9; the timing of the additional slip could not be 

determined. 

DECEMBER 16, 1954 (M-7.1 AND 6.8) 

The Dixie Valley-Fairview Peak earthquakes of December 16, 1954, produced a 90-km-long zone of rightlateral oblique and normal faulting in the central Nevada 

seismic zone (fig. 6.11; Slemmons, 1957). The first shock, 

which occurred east of Fairview Peak, produced lateral 

displacements of more than 4 m and vertical displacements of as much as 3 m. Faulting along this 50-km-long 

zone was predominantly down to the east opposite 

Fairview Peak and changed polarity to the north. The 

second shock, which occurred 4 minutes later, had an 

epicenter on the east side of Dixie Valley in a leftstepping echelon arrangement with the earlier event. 

Normal-fault scarps formed along a 40-km-long zone at 

the base of the Stillwater Range some 20 km west of the 

Rainbow Mountain faulting. Vertical displacements exceeded 2 m, and consistent strike-slip displacements 

were absent. 

The Chalfant Valley earthquake of July 21,1986, is the 

largest event to date in a series of 33 earthquakes of 

M,z6 to occur since 1978 in the White Mountain seismic 

gap (Savage and Cockerham, 1987). Other principal 

events in this series include the May 25-27, 1980, 

Mammoth Lakes earthquakes (M=6.l, 5.9, 5.8, 6.0) and 

the November 23, 1984, Round Valley earthquake 

(M=5.7). The series of shocks is of interest not only 

because of its wide geographic distribution in the White 

Mountain seismic gap but also because of the contemporaneous uplift of Long Valley caldera (Hill and others, 

1985). The Chalfant Valley earthquake created a lo+- 

km-long zone of fractures with as much as a few centimeters of dextral slip on the frontal-fault zone of the 

White Mountains (Lienkaemper and others, 1987). The 

earthquake focal mechanism and aftershock distribution 

show that the predominately dextral strike-slip displacement associated with this event occurred on a westdipping fault plane that projects upward to meet the 

surface break. 

FIGURE 6.l'.-Fault trace of IN5 Pleasant Valley, Nov., carthq~akc remains clearly visible in this photograph by R.E. Wallace more than 6 

years after event (Wallace, 1984). 
ti. EARTHQUAKE HISTORY. 1769-1989 

SEISMICITY OF 

THE MENDOCINO TRIPLE JUNCTION 

AND THE GORDA PLATE 

The San Andreas fault terminates at its north end in a 

transfonn/transfonn/trench triple junction just seaward 

of Punta Gorda. Major earthquake activity lies along the 

Mendocino Fracture Zone. where it is an active transform fault, and to the north within the Gorda plate, which 

is undergoing intense internal deformation. The WadatiBenioff zone is well defined to a depth of 30 km and can 

be traced eastward to adepth of mote than80 km (see fig. 

5.6: Walter. 1986). Strong earthquakes within the Gorda 

plate locate off shore and span the position of the 

megathrust; these events appear to lie entirely within the 

oceanic lithosphere. The 1980 Eureka earthquake, for 

example, ruptured the Gorda plate from the landward to 

the seaward side of the megathrust. Despite the high 

level of seismicity, underthrusting events are rare. 

PRINCIPAL EARTHQUAKES 

NOVEMBER 23. 1873 (M=W+) 

The severe earthquake of November 23,1873, was felt 

from San Francisco to Portland, Oreg.; it inflicted the 

heaviest damage to Crescent City, Calif., and surrounding communities in the Klamath Mountains. The macroseismic epicenter near the California-Oregon State line 

and probably inland of the coastline is unique within both 

the historical and instrumental records. 

APRIL 16, 1899 (M=7) 

Little is known about the large earthquake of April 16, 

1899, with an epicenter seaward of Eureka, where it was 

described as "one of the severest shocks of earthquake 

ever experienced." Toppomda and others (1981) corrected the origin time of this event and assigned a nearshore 

epicenter and an M, of 5.7. The earthquake was assigned 

an epicenter in the Gulf of Alaska by Milne (1901) on the 

basis of the traveltime of the maximum amplitude from 

the five reporting stations; however, a California location 

satisfies his data equally well. The absence of significant 

damage along the coast suggests an epicenter well out to 

sea. An instrumental magnitude (Mg) of 7.0 is derived 

from the surface-wave amplitudes reported by Miie (see 

Abe and Noguchi, 1983). 

JANUARY 31,1922 (M-7.3) 

The intensity pattern of the large earthquake of 

January 31,1922, is generally similar to that of the 1899 

event. This event was well recorded throughout the 

world. 

JANUARY 22, 1923 (M-7.2) 

The earthquake of January 22, 1923, strongly shook 

the Cape Mendocino region and toppled many chimneys 

in the area. This earthquake was probably associated 

with the Mendocino Fracture Zone. 

DECEMBER 21, 1954 (M=6.6) 

The strong earthquake of December 21, 1964, apparently was located in the crust of the North American 

plate above the descending Gorda plate. The relocation of 

this event by Smith and Knapp (1980) suggests a possible 

association with the active Mad River fault zone. One 

fatality is attributed to the earthquake. 

NOVEMBER 8, 1980 (M~7.2) 

The Eureka earthquake of November 8,1980, resulted 

from 100-km-long, left-lateral strike-slip rupture of the 

Goda plate along a northeast-striking fault (see fig. 5.5). 

Aftershocks of the earthquake extended from within 30 

km of the coastline southwestward to the Mendocino 

hture Zone. TTie focal mechanism of ~ the ~~- earthouake -~~-~~is >~~~--- -- 

thus conjugate to the San Andreas, with its tension axis 

aligned in the downdip direction. This event argues for 

high rates of internal deformation within the subducting 

oceanic lithosphere and against the extension of San 

Andreas-style faulting northward of the triple junction. 

DISCUSSION 

The spatial distribution of large earthquakes during 

the past 2 centuries defines the San Andreas fault system 

as a 100- to 300-km-wide zone containing numerous active 

faults in addition to the San Andreas fault proper (fig. 

6.1). Except for the two largest events, the great 1867 

and 1906 earthquakes that together ruptured two-thirds 

of the total length of the San Andreas fault, large 

earthquakes are conspicuously absent along the master 

fault itself. Although these two great earthquakes account for half of the seismic-strain release since 1769, 

most of the rest occurs on other, smaller elements of the 

fault system. Major historical events on these secondary 

faults, such as the 1927 Lompoc and 1952 Kern County 

earthquakes, serve to define the boundaries of the San 

Andreas system. Their mechanisms differ significantly 

from right-lateral strike slip parallel to the plate-motion 

vector and illustrate the diversity and complexity of 

seismic-strain release within the plate-boundary zone. 

Over the timespan of the historical catalog, the most 

enduring characteristic of the earthquake distribution 

may be the spatial clustering of activity at specific 

localities along the plate boundary. Notable hotspots 
170 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

include the Cerro Prieto, Imperial, San Jacinto, and 

Calaveras faults, all of which are major branches of the 

San Andreas fault, and the Parkfield segment of the San 

Andreas fault in the transition zone between the 1857 

rupture and the 150-kin-long central, creeping segment 

of the fault. In each of these areas, the seismic activity 

coincides with these high-sliprate faults (1-3.5 cmlyr), 

and in some places it dearly represents recurrent rup 

tore of the same segment of the fault. At greater 

distances from the San Andreas fault, the historical 

events (or sequences) tend to represent isolated occurrences on slower moving faults. Thus, although the 

overall seismicity spans the broad plate-boundary zone, 

seismic-strain release over the past 2 centuries correlates 

with the local rate of fault movement. 

In general, the locations of historical earthquakes 

resemble the overall distribution of microearthquake 

activity, despite more than six orders of magnitude 

difference in average seismic moment (fig. 6.13; see chap. 

5). One important difference between the distribution of 

large and small earthquakes is the virtual absence of 

smaller events along the San Andreas fault segments 

that ruptured in 1857 and 1906. Similarly, seismic activity is distinctly absent on the potentially dangerous 

segment between the 1857 break and the Imperial 

Valley. Except for the central, creeping segment, where 

numerous small events occur, the San Andreas fault is 

almost completely aseismic during the long intervals 

between its ruvture in maim earthquakes (see figs. - 5.6. 

5.9). 

This inverse correlation between the source regions of 

large earthquakes and the distribution of smaller events 

can also be observed for smaller main shocks. Recent 

studies of the rupture dynamics of M=6 events occurring 

within seismically active regions indicate that the rupture zones of these events are similarly aseismic, with 

smaller events occurring predominantly outside the slip 

surface, even during the aftershock sequence (Reasenberg and Ellsworth, 1982; HartzeU and Heaton, 1986; 

Mendoza and Hartzell, 1988). Thus, the rites of future 

large earthquakes cannot be identified on the basis of 

minor seismicity alone. 

FIGURE 6.13.-Distribution of large and small earthquakes along the 

San Andreas fault system. Ingeneral, modem instrumental data (C; 

see chap. 5) portray same pattern of activity seen in large earthquakes from preinstnunental (A) and instrumental (5) eras. Some 

areas characterized by high levels of microearthquakes, such as 

well-defined faults east of northern section of the San Andreas fault 

(red line), have not produced significant earthqudies in historical 

time and BO are considered probable sites of future activity. 
6. EARTHQUAKE HISTORY. 1769-1989 

RATE OF SEISMICITY 

The average rate of earthquake activity within the San 

Andreas system can be estimated from the GutenhergRichter frequency-magnitude relation log N=a-bM, 

where N is the cumulative number of events of magnitude equal to or greater than M during a given time 

period. For the 77 events along the fault system with 

summary magnitudes Mz6 since 1852, this relation well 

describes the population with a=5% and b=l (fig. 6.14). 

Comparable results are obtained for subsets of M=6 

events, such as the instrumental period (1898-1989). 

It is useful to compare these results from the historical 

record with the frequency-magnitude relation determined from systematic microearthquake observations. If 

the historical rate of activity applies today and the 

frequency-magnitude relation for microearthquakes 

(M23) is described by the same relation, then about 5,600 

MÃˆ events should be observed each year. This prediction exceeds the number of events observed during the 

interval 1980-87 by about a factor of 2 (see chap. 5) and 

suggests that a somewhat smaller value of b=0.93 may be 

more appropriate for the extended magnitude range. 

For the catalog as a whole, the rate of earthquake 

occurrence is well described by a Poisson process, in 

which the probability of finding one or more events in any 

$ A Son Andreas system, 1852-1989 

Z 5 0.1 1 7 

. 

Sam 

North 

(Af 

boiindary, 

Andreas 5 

Amerlcan.Paclfic 

4.5, 

1852-1989 

system, 198046 

plate 

\ ' North American-Paclnc plate 

boundary, 1980-86 (M 5 4.5) 

10.~ I 

1 2 3 4 5 6 7 e 

MAGNITUDE 

FIGURE 6.14.-Annual frequency of earthquakes of magnitude ZM as 

derived from historical and modern instrumental catalogs. Gutenberg-Richter frequency-magnitude relation, log N=a-bM with 

o=VJyr and b=l, describes observed distribution of earthquakes of 

M26 within the San Andreas fault system during 138-yr interval 

from IS52 to 1989. Also shown are annual fpeouenpv of Ms-6 events 

from the broader Pacific-Nonh American bo&, including 

the San Andreas fault system and the western Basin and Range 

province, and of M=2-4.5 evenu in both regions during 1980443. 

interval oft years is P=l-e-'t, where X is the average 

rate of earthquake occurrence. It follows from the 

observed frequency-magnitude relation that the odds of 

having at least one M26 event per year are 0.43. There 

is also an even chance of at least one Mz6 event within 

any 15-month interval, one M^l within any 12V~year 

interval, or one Ms8 within any 125-year interval. 

The rate of earthquake activity along the plate boundary can also he usefully compared with plate-motion 

estimates derived from plate-tectonic theory. Current 

estimates of the relative velocity across the North 

American-Pacific plate boundary, determined from the 

spreading rate in the Gulf of California of 5 cmlyr 

(DeMets and others, 1987), imply an annual seismicmoment rate of 2x1026 dyne-cdyr for a 10-km-thick 

brittle crust, equivalent to a single M=6.8 earthquake. 

Earthquakes of this size occur far less often, and the 

principal seismic contribution to the plate motion comes 

from infrequent large events. The erroneous notion that 

the smaller events substantially contribute to the total is 

demonstrably false, as shown by summing the contrihutions of all the earthquakes below some magnitude. The 

innumerable events of Ms6 occurring each year contribute less than 10 percent to the total seismic-strain 

release. 

Within the San Andreas fault system, the total seismicmoment release since 1852 corresponds to 70 percent of 

the total North American-Pacific plate motion predicted 

by plate-tectonic models. This proportion is somewhat 

inflated because not all of the earthquakes act to transmit 

slio alone the date boundary: for examole. the 1952 Kern 

county earthquake, the third largest historic time, 

directly accommodated little plate-parallel motion. Although aseismic displacements account for some of the 

deficit, notably along the central, creeping section of the 

San Andreas fault, deformation occurring elsewhere, 

notably within the Basin and Range, contributes substantially to the relative motion between the North American 

and Pacific plates. 

PARADOX OF THE MISSING PLATE MOTION 

The discrepancy between plate-tectonic estimates of 

relative motion across the North American-Pacific date 

boundary and seismic estimates also holds for geologic 

and geodetic estimates of motion along the San Andreas 

fault system. The explanation of this apparent paradox is 

thought to include deformation within the Basin and 

Range province in western Nevada and eastern California (Atwater, 1970), which has been the locus of major 

seismic activity in historical time, including the third 
172 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

largest event, in 1872, and 3 of the 11 M--7 events in the 

20th century. 

It has long been recognized that the Basin and Range 

province has undergone substantial extension during the 

Cenozoic and is presently opening in a N. 60Â°W direction 

(Zoback and Zoback, 1980). Historical seismicity partly 

agrees with this geologically derived pattern; however, it 

also indicates a significant component of dextral shear in 

nearly every well-studied historical event (Shawe, 1965, 

Doser, 1986). Because the geologic expression of strikeslip displacement is much more difficult to recognize and 

quantify than vertical slip, a major question is raised 

about the significance of the historic seismicity for the 

total strain within the western Basin and Range. 

Since the 1872 earthquake, the net seismic strain 

within the Basin and Range province can be estimated by 

summing the contributions of individual events. The net 

shear strain thus determined indicates nearly equal 

components of extensional strain in a N. 60- W. direction 

and dextral shear trending N. 10' W. The resulting 

average-motion vector nearly coincides with the orientation of the San Andreas fault, and the lateral slip largely 

balances the coastward expansion of the province that 

results from extension alone. If both the rate and style of 

historical faulting accurately portray the long-term deformation within the region, they diminish the discrepancy between the predicted and observed rates of motion 

across the North American-Pacific plate boundary. 

EARTHQUAKE RECURRENCE AND 

CHARACTERISTIC EARTHQUAKES 

Over geologic time, the net displacement across a fault 

accumulates through the action of countless individual 

slip events. Measured over many displacement cycles, 

the average interval between events must equal the 

average event displacement divided by the remote slip 

rate. First principles, however, provide little guidance as 

to the properties of the recurrence, which might range 

from a totally random distribution of events in both space 

and time to identical earthquakes repeating at fixed 

intervals. If recurrence is essentially random, then 

long-term seismic hazard is described by the Poisson rate 

of activity, as discussed above. Greater regularities and 

systematics in recurrence imply that useful time-dependent forecasts of future activity can be derived from 

knowledge of the past behavior of the fault system. 

Results for San Andreas earthquakes have played a 

central role in establishing the existence of broad regularities in the recurrence process. At Parkfield, the San 

Andreas fault has ruptured six times since 1857 in M-6 

events with highly repeatable characteristics every 22Â± 

years. The latest three events, in 1922, 1934, and 1966, 

for which instrumental records exist, are virtually identical (fig. 6.15; Bakun and McEvilly, 1984). Amplitude 

iata from Milne seismographs uncovered in the preparation of table 6.1 show that the 1901 and 1922 events 

produced the same surface-wave amplitudes on common 

stations, strengthening earlier speculations that all the 

20th-century events are similar. Intensity data for the 

1881 event (Toppozada and others, 1981) and for foreshocks to the great 1857 earthquake (Sieh, 1978b) place 

them along the Parkfield segment as well. These regularities in the size, location, and timing of all known 

events at Parkfield led Bakun and Lindh (1985) to 

propose a specific recurrence model for Parkfield earthquakes. On the basis of this model, the next in the series 

af characteristic events is anticipated before 1993, and its 

forecast represents the first formally endorsed earthanake prediction in the United States. 

Geodetic analysis of the strain released in the 1966 

earthquake and its subsequent reaccumulation led Segall 

and Harris (1987; see Harris and Segall, 1987) to identify 

the zone where strain accumulates and is released, the 

DBN-EW 

- I Rayleioh wave 

Love wave 

FIGURE 6.15.-surface waves of lgZZ,l$34, and 1966 Parkfield, Calif., 

farthniiakpa aa recorded on same seimoirranh -. in DeBiIt. The 

Netherlands tDBN; EW, east-west; NS. north-south). These nearly 

identical waveforms and amplitudes led Baku" and McEvilly (1984) to 

propose recmnt rupture of same segment of the San Andreas fault 

&mechanism of Parkfield earthquakes, 
6. EARTHQUAKE HISTORY. 1769-1989 

'Parkfield asperity," as the center of the 1966 rupture 

zone. This zone of strain accumulation appears to be 

effectively locked during the interseismic period and 

corresponds to the center of the 1966 aftershock zone 

(Eaton and others, 1970) between about 4- and 10-km 

depth. The significantly fewer events in this part of the 

aftershock zone than in its periphery suggests that 

Parkfield earthquakes occur when this locked zone suddenly releases. Aftershocks appear to result from transfer of stress to the perimeter of the asperity. 

This same pattern of concentrated coseismic slip occupying a quiet region within the overall aftershock distribution characterizes several recent, well-studied events 

(Mendoza and Hartzell, 1988), three of which, the Coyote 

Lake earthquake (Aug. 6, 1979). the Imperial Valley 

earthquake (Oct. 15, 1979), and the Morgan Hill earthquake (Apr. 24, 1984), all have probable antecedents 

within the historical record. Reasenberg and Ellsworth 

(1982) identified the June 20, 1897, earthquake as a 

predecessor to the 1979 event and noted that the 82-year 

interval between events equaled the 1.2 m of coseismic 

slip determined by Liu and Helmberger (1983) divided by 

the long-term slip rate of 1.5 cmlyr for the Calaveras 

fault. Similarly, the 73-year interval between the July 11, 

1911, event and the 1984 Morgan Hill earthquake (Bakun 

and others, 1984) well predicts the 0.8 to 1.0 m of 

maximum coseismic slip determined by Hartzell and 

Heaton (1986). The 1979 Imperial Valley earthquake is 

more complex because it reruptured only the northern 30 

km of the May 19, 1940, fault break. Again, both the time 

interval between events and the fault-slip rate compare 

favorably with the fault slip at depth, as determined from 

seismograms (Hartzell and Heaton, 1983; Archuleta, 

1984). Earlier ruptures of this or other segments of the 

Imperial fault may well be in the historical record, 

possibly including the April 19, 1906, event, which 

occurred the afternoon of the great 1906 earthquake in 

northern California. 

Similar observations of recurrent faulting in events 

with characteristic magnitudes and locations from around 

the world (Nishenko and Buland, 1987) suggest a simple, 

first-order model for seismic potential. In this model, the 

future behavior of a specific segment of a fault can be 

forecast from knowledge of the size of past earthquakes, 

the timing and amount of slip in the latest event, and the 

long-term rate of fault movement (Lindh, 1983; Sykes 

and Nishenko, 1984). Accordingly, the probability of an 

event on a recently ruptured fault segment is low until 

the elastic strain rebuilds, which may he estimated from 

the geologic slip rate. As the strain rebuilds, the probability of another earthquake increases. Empirically, the 

time intervals between successive ruptures of a specific 

fault segment define a bell-shaped distribution that may 

be used to estimate the odds of the next event within 

some future time interval, given that it has not yet 

occurred. 

Probabilities for large earthquakes along the major 

branches of the San Andreas fault derived from this 

methodology differ markedly from Poisson estimates 

(Working Group on California Earthquake Probabilities, 

1988). For example, the chance of a repetition of the 

great 1906 earthquake within the next 30 years 

(1988-2018) is less than 0.1. In contrast, the chance of an 

M=7V2-8 earthquake on the southern section of the San 

Andreas fault is 0.6. When the Working Group's report 

was written, the southernmost part of the 1906 fault 

break was assigned the highest chance of failure of any 

segment of the north half of the San Andreas fault. Now 

that it has ruptured in the October 18,1989, Loma Prieta 

earthquake, the probability of another rupture will be 

small for several decades. A clearer understanding of 

past seismicity can only help to improve and refine 

estimates of future seismicity. 

THE SEISMIC CYCLE 

An important implication of the characteristic-earthquake model is the existence of a repetitive cycle of strain 

accumulation and release (Fedotov, 1968). Mogi (1981) 

suggested the existence of definite stages in the cycle, 

including a low level of seismicity in the first part of the 

cycle once aftershock of the latest event subside, a rise in 

regional activity as strain accumulates, and ultimately 

the occurrence of another earthquake with its attendant 

foreshocks and aftershocks, which initiates the next 

cycle. 

The long-term seismicity within the San Andreas fault 

system displays these characteristics along the rupture 

zone of the great 1906 earthquake (figs. 6.16, 6.17; 

Ellsworth and others, 1981). Activity was relatively high 

during the 19th century, as becomes particularly apparent after 1850, when the record is virtually complete. 

After the great 1906 earthquake, the level of seismicity 

changed drastically, and moderate events essentially 

ceased for 50 years. Since the mid-1950's, the activity 

level has increased and begun to approach the 19thcentury level (Tocher, 1959). This change in activity 

associated with the 1906 earthquake has been noted 

many times (for example, Gutenberg and Richter, 1954). 

and it is an open question whether it represents a 

premonitory increase (Toppozada and others, 1988) or 

whether the long quiescent period since 1906 is the 

essential feature (Ellsworth and others, 1981). 

Comparable variations in seismicity appear to be 

present in southern California, although the historical 

record there is less reliable until about 1890. Along the 

rupture zone of the great 1857 earthquake, the available 

data suggest a similar period of low activity for several 
174 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

decades after the event (fig. 6.17). Farther south, along 

the section of the fault that has not ruptured in 3 

centuries, the activity level since at least the 1880's is 

reminiscent of the activity in the San Francisco Bay 

region before the 1906 earthquake (fig. 6.18). As a 

potential long-term indicator of high seismic potential, 

the seismicity surrounding the dormant southern section 

of the San Andreas fault agrees with independent estimates of long-term potential derived from paleoseismolom. 

FUTURE USES OF EARTHQUAKE HISTORY 

At this stage in our understanding of the San Andreas 

fault system, seismicity is still best described as arandom 

EXPLANATION 

Munitaide 

FIGURE 6.16.-Seismicity of the San Francisco Bay region in quarter- bepm to approach levels last seen in the 19th century. However, 

century epochs. Activity was high in the region during at least a both geologic and geodetic evidence suggest that the next great 

half-century before 1906 earthquake and drastically declined after- earthquake will not occur for a century or more. 

ward for the next half-century. Since the mid-1950'8, activity has 
6. EARTHQUAKE HISTORY, 1769-1989 175 

process over time, with a highly clustered spatial distribution. There are, however, tantalizing hints of underlying regularities, such as those in the characteristic 

earthquakes at Parkfield, or in the striking changes in 

seismicity associated with the 1906 earthquake. The next 

generation of refinements to this history will assuredly 

make comparable contributions by reducing the uncertainty in earthquake locations and magnitudes. Modem 

seismologic methods for extracting new information on 

the mechanisms of earthquakes have already proved 

practical for many events from the early instrumental 

period. Systematic treatment of the fall instrumental 

catalog with these methods will provide a new basis for 

understanding the tectonics of the plate boundary and 

the mechanics of earthquakes. 

CATALOG OF MAJOR EARTHQUAKES, 

1769-1989 

CATALOG COMPILATION 

The publication of Edward S. Holden's catalog of 

Pacific coast earthquakes in 1898 represented the first 

systematic scientific inquiry into the seismic history of 

California and surrounding regions. This catalog, and its 

extension by McAdie (1907). formed the primary basis for 

the monumental catalog of Townley and Allen (1939) 

covering the years 1769-1928. These catalogs provide 

detailed descriptive accounts of virtually all the earthquakes that are now known from this period, and all 

subsequent analyses of seismicity up to the modem 

instrumental period build on these foundations. 

Recent studies of preinstrumental seismicity have 

focused on quantification of the historical record. The 

catalog presented here relies heavily on the research of 

Tousson Toppozada and his associates (Toppozada and 

others, 1981; Toppozada and Parke, 1982), who developed extensive new information on seismic intensities 

from newspaper accounts and other original sources, and 

determined locations and magnitudes from the resulting 

isoseismd maps. In addition, several special studies of 

important events by other workers have contributed to 

the catalog. 

The development of practical seismographic instnmentation around the turn of the 20th century led to the 

rapid growth of seismologic data, particularly for those 

events large enough to register at teleseismic distances 

on the early instruments. The publication of the Circulars 

of the Seismolo~cal Committee of the British Association 

for the Advancement of Science (1899-1912) and their 

continuation as the International Seismological Summary 

from 1913 on indicate a detection threshold of about M-6 

for the Western United States as early as 1898. Data 

from these and other sources enabled Gutenberg and 

Richter (1954) to systematically catalog seismicity from 

1904 onward. 

Modern seismographic instrumentation first installed 

in California in 1910 ushered in the era of earthquake 

EXPLANATION 

Mimmlude . > 5.5 

A > 6.0 

T 2 6.5 

> 7.0 * > 7.5 

A' 

1400 

2 1200 tu 

k ,004- 

0 

^- 

Z em- 

"i " m 5 " 

0 400 

A oYEAR 

I I I I I I I - I A .I IT <s - A A . a. A I T 

*. 1- A T 

- 1906 rupture - 

...a 

. "A - A * t' ' . A. A* .r .T 4 ..I 

A 

A . A Ãˆ - 

- TÃ‘Ã A A .* * - 

A A. 

- A TA A - T - . A '7. A T A 

-7 . A Â * a A Â - A . T " 

I I I I T . . '! T 

I I . I A I 

FIGURE 6.17.-Space-time diagram of seismicity since 1850 along the 

San Andreas fault system between head of the Gulf of California (A) 

and Punta Gorda (A'). Change in seismicity rate along northern 

section of the San Andreas fault associated with 1906 earthquake (fig. 

6.13) may also be tentatively identified along 1857 earthquake 

rupture. Persistent activity characterizes south third of the plate 

boundary since 1890. spanning the entire interval of reliable earthquake reporting in this region. 

1 1 1870 1880 1890 1900 1910 1920 1930 1840 1950 1960 1970 1980 1W 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

observation at regional distances. The BuUetins of the 

Seismographic Stations of the University of California, 

117' 116' 119 

EXPLANATION 

Mogmtude 

+ 25.0 Ã‘Ã Flint- Doited 

25.5 when: concealed 

A >6.0 

T 2 6.5 

Â 2 7.0 * > 7.5 

Berkeley, from 1910 to the present form the principal 

source for events in northern California and adjoining 

areas. Routine epicentral determinations and magnitude 

assignments for earthquakes in the southern California 

region date from 1932 and are taken from the catalog of 

the Seismological Laboratory of the California Institute 

of Technology, Pasadena. Additional instrumental results come from various other sources, chiefly the U.S. 

Geological Survey and the University of Nevada, Reno. 

The resulting catalog of major earthquakes in California, western Nevada, and northernmost Baja California 

from 1769 to 1989 (table 6.1) contains 206 entries. This 

catalog omits several earthquakes listed in earlier catalogs where this or other recent studies have failed to 

corroborate previous interpretation as significant events 

or even, in some cases, their occurrence. 

QUANTIFICATION OF EARTHQUAKES AND 

MAGNITUDE SCALES 

Physical measures of the complex mechanical event 

producing the earthquake take many forms, including the 

dimensions of the faulted region, the amount of slip, and 

the strength of the radiated elastic waves. To relate the 

characteristics of one event to another, the observed 

quantities must generally be summarized through the use 

of either an empirical relation, such as magnitude, or a 

quantity derived from a physical model, such as seismic 

moment. Both procedures have their place, and the 

choice of one metric over another depends principally on 

the purposes of the comparison and the availability of 

common data. 

Because no single procedure for determining magnitude can he applied to the entire historical record, the 

catalog must be quantified by using various magnitude 

scales. Each scale is briefly described below to define its 

origin and to clarify its relation to the other scales. I 

emphasize that each scale has a particular range of 

validity and that different magnitude scales will, in 

general, yield slightly different values for the same 

event. Such differences in magnitude seem to provide a 

never-ending source of interest and controversy for the 

news media, who commonly lump aU scales together 

under the heading of "Richter scale."To the seismologist, 

4 

FIGURE 6.18.-Where will the next great earthquake strike along the 

San Andreas fault system? Numerous lines of evidence point to its 

longdormant southernmost segment (6) as having the highest 

potential. Large earthquake activity in this region shares many 

similarities with activity in the San Francisco Bay region before 1906 

earthquake (A). In both cases, absence of activity directly on the San 

Andreas fault is pronounced, and a high regional level of activity is 

concentrated along other major branches of the fault system. 
6. EARTHQUAKE HISTORY. 1769-1989 

such differences are neither surprising nor controversial 

and can, in fact, provide information on the underlying 

physical processes of the earthquake source. 

I also emphasize that intensity scales characterize the 

effects of the earthquake at a particular location and are 

not magnitude scales. Strictly speaking, intensity values 

(or, for that matter, instrumentally measured values of 

ground-motion parameters) describe the vibratory motions that are the actual earthquake as observed at a 

particular location, whereas magnitude values describe 

the faulting event that generates the earthquake. 

THE RICHTER SCALE (M,J 

The original magnitude scale of Richter (1935) was 

introduced for the purpose of providing an objective 

measure of the energy of each earthquake in the initial 

listing of earthquakes in the southern California region 

compiled by the Seismological Laboratory in Pasadena. 

Rather than attempting to measure the energy of the 

earthquake source directly, he chose to construct an 

empirical scale derived from a simple measure of the 

complex seismic waveform. Using only the maximum 

excursion of the seismogram as measured on a single type 

of instrument, the Wood-Anderson seismograph, he 

defined the local magnitude of an earthquake as 

where the empirical function A) depends only on the 

epicentral distance of the station, A. The zero point was 

arbitrarily set by Richter to avoid negative magnitudes 

in the course of routine work. Use of common logarithms 

means that two earthquakes located at the same distance 

from a station and having peak amplitudes differing by a 

factor of 10 will differ by 1 magnitude unit. In practice, 

readings from all observing stations are averaged after 

adjustment with station-specific corrections to obtain the 

ML value. Although Richter (1935) predicted that the 

local-magnitude scale "cannot hold to any high accuracy," 

history has proved it to be a powerful quantitative tool 

for ordering the relative sizes of earthquakes. 

Several points about ML should be emphasized. First, 

it is strictly defined only for the southern California 

region, although its applicability to coastal central and 

northern California has since been shown. Recent studies 

of the A) curve suggest that it will require revision and 

regionalization. Second, because ML has no actual physical units associated with it, other empirical magnitude 

scales may be freely adjusted to coincide with it. The 

local-magnitude scale has, in fact, been used as the basis 

for establishing essentially all other magnitude scales. 

Finally, because ML is derived from measurements taken 

from a single, band-limited seismograph, Mr values 

saturate once an earthquake becomes large enough. 

Thus, the "correct" Richter magnitude ML=6.9 for the 

great 1906 earthquake obtained by Jennings and Kanamori (1979) reflects the amplitude of seismic waves at 

periods near 1 s but not the total energy of this 

earthquake. Uniformly valid characterization of the 

"size" of an earthquake requires use of magnitude scales 

based on longer-period measures of the event. 

SURFACE-WAVE MAGNITUDE (M.) AND 

BODY-WAVE MAGNITUDE (nib) 

The successful development of the local-magnitude 

scale encouraged Gutenberg and Richter to develop 

magnitude scales based on teleseismic observations of 

earthquakes. Two scales were developed, one based on 

surface waves, Ms, and one on body waves, m,,. 

Surface waves with a period near 20 s generally 

produce the largest amplitudes on a standard long-period 

seismoeranh. -., and so the amnlitude of these waves is used 

to determine Ms. using an &ation similar to that used 

for Mr. 

The body-wave magnitude, mb, which was developed 

specifically to treat deep-focus earthquakes, presents yet 

another alternative scale for magnitude determination. 

Although it presently is the most commonly reported 

teleseismic magnitude, current practice in its determination differs from that employed by Gutenberg, and so it 

is omitted from table 6.1. As a short-period magnitude, 

modern mi, values measure the same part of the earthquake energy spectrum as Mr. 

The magnitudes listed by Gutenberg and Richter 

(1954) that appear in table 6.1 as MSx are essentially Ms 

according to Geller and Kanamori (1977); magnitudes 

attributed to Richter (1958) are based on M, or Ms. 

Useful estimates of Ms can be obtained from many 

different types of low-period seismographs, including 

the undamped instruments deployed by Pklne beginning 

in 1897. Abe and Noguchi (1988) constructed estimates of 

Ms from Milne seismograms to resolve a longstanding 

controversy concerning an apparent peak in global seismicity between 1904 and 1912. Abe (1988) later used the 

Milne data to determine magnitudes for smaller earthquakes in California between 1898 and 1912. His procedures have been used to compute Ms for additional 

California events occurring between 1898 and 1934, 

which are listed in table 6.1. 

SEISMIC INTENSITY AND EARTHQUAKE MAGNITUDE (M,) 

Before the development of seismographs in the late 

19th century, descriptions of the effects of earthquakes 

provided the only means for assessing earthquake size in 

all but the rare cases where surface faulting was well 
178 THE SAN ANDEEAS FAULT SYSTEM. CALIFORNIA 

described. A robust method for relating the area undergoing shaking of a given intensity or greater to ML was 

developed by Toppozada (1975) for California and western Nevada. Using these relations, Toppozada and others 

(1981, 1982) successfully assigned an intensity magnitude, MI, to many earthquakes. The isoseismal maps 

developed in the course of their research also generally 

provide our best estimates of epicentral locations. New 

MI values have been determined for several events, 

using the same procedures as part of this study. 

SEISMIC MOMENT (Md, RADIATED ENERGY, AND 

MOMENT MAGNITUDE (M) 

Magnitude scales based on finite-bandwidth seismographs approach a maximum near which events of clearly 

different size or energy are indistinguishable. Saturation 

of ML is apparent for both the 1906 and 1952 earthquakes 

listed in table 6.1. Recent work by Hutton and Boore 

(1987) suggests that the local-magnitude scale may begin 

to saturate at about ML=6. Such saturation, which is 

understood to arise from the scaling law of the seismic 

spectrum (AM, 1967), occurs when the peak of the energy 

spectrum lies below the frequency range of the WoodAnderson seismograph. 

By using the well-known properties of the seismic 

spectrum, magnitude scales can be constructed with 

uniform validity. One such scale, Mw, proposed by 

Kanamori (1977) is based on the seismic energy radiated 

in the form of elastic waves by the source. Another nearly 

equivalent magnitude scale, M, the moment magnitude, 

is based on the seismic moment, Ma= yAv, (AM, 1966), 

where A is the area of the earthquake rupture surface, u 

is the average fault displacement, and u. is the shear 

modulus of the crustal volume containing the fault. 

Hanks and Kanamori (1979) took advantage of the nearly 

identical relations between Ma and both ML and Ms to 

define M=S'slogio Mo-10.7, where My is measured in 

dyne-centimeters. 

These two magnitude scales, though closely related, 

are not identical. Singh and Havskov (1980) showed that 

MW=%(1og,,, I&+logia Aulp-12.1), where Au is the 

stress drop. Earthquake stress drops generally fall in a 

narrow range over the entire magnitude spectrum, and 

so with Aulp-lo4 (Kanamori, 19771, Mw,=M. One advantage to M for the purpose at hand is itsdependence on 

onlv the static fault offset and ruDture area. which can be 

determined for the 1857 and 1872 earthquakes 

SUMMARY MAGNITUDE (M) 

To construct a single, summary-magnitude scale, M, to 

characterize the relative size of all the events listed in 

table 6.1.1 use each of the scales described above, being 

careful to consider such factors as the historical period 

and event location, as well as the quality of individual 

determinations. Where choices between several magnitude estimates are possible, the summary magnitude, M, 

is weighted toward long-period estimates of magnitude. 

Specifically, Ms and MGR are selected when judged 

reliable (110 events). Many local magnitudes have thus 

been superseded by surface-wave magnitudes; this effect 

is most noticeable for the largest events, where saturation of ML becomes important. ML is the principal 

contributor to 20summary magnitudes, half of which also 

have reoorted M values that amee well. For all but two 

events before 1898 (1857 and-1872) and for two 20thcentury events, M is based on M,. In effect, the summary 

magnitude is an intensity magnitude before 1898 and a 

teleseismic surface-wave magnitude thereafter. 

If M is to be uniformly validity across the entire 

timespan of the catalog, MI must be an unbiased estimator of Ms or MGR. To test this absence of bias, the 

correlation between MI and the two surface-wave magnitudes has been examined for 23 events with reliable Ms 

or MGR estimates, and an MI value determined from the 

isoseismal maps of Toppozada and others (1981) and 

Toppozada and Parke (1982). This comparison shows that 

although the two magnitude scales are well correlated, 

MI systematically underestimates Ms and MGR by 

0.3Â±0. units. If the sample is restricted to MsS6.5 

(%=IS), the bias is 0.2k0.3 units. To further investigate 

this apparent bias, ML was compared with MI for 10 

common events, for which the bias was 0.26Â±0. units. 

As a final check, the difference between Me or MGR and 

MI for the 12 events listed in table 6.1 also used by 

Toppozada (1975) to develop MI relations was found to be 

0.10Â±0.1 units. 

On the basis of these results, the summary magnitudes 

from MI values of Toppozada and others (1981) have been 

adjusted upward by 0.15 units and then rounded to the 

nearest quarter magnitude unit. Thus, events of MI=5.7 

become M=5%, and those of MI=5.8 become M=6. No 

magnitude adjustment exceeded a quarter unit. MI 

values from other sources have simply been rounded to 

the nearest quarter unit, because they average 0.2 units 

higher than the values of Toppozada and others (1981) 

and Toppozada and Parke (1982), where comparisons can 

be made. Summary magnitudes for events before 1860 

have not been adjusted upward, owing to the imprecision 

of the original estimates. 

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62 P. 

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TABLE 6.1 

Earthquake origin times, magnitudes, and locations before 1990 principally derived from interpretations of felt reports by Toppozada and others (1981); after 1900, data principally derived from 

Gutenberg and Richter (1954) and bulletins of the California Institute of Technology, University of 

California, Berkeley. University of Nevada, Reno. and U.S. Geological Survey. See text for discussion 

of summary mamitude M. Other maimitude scales are Mr. ". local-maenitude scale of Richter 119351: 

. . 

MC.* m&tudesfromGutenbergand Richter(1954~. generally baaednn2(t-asurfacewaves: Ms. 20- 

durface-wave mwtude; .M,, magnitude eslimated from felt area at various intensity levels, and M. 

moment magnitude, defined as M=?ilog,, Mn- 10.7, where .no is in dyn-ntimetera-parenthetical 

values based on seismogram envelope. Ms values before 1985generally derivedfrom undamped Milne 

seismographs, using the formula of Abe (1988)-parenthetical values based on one or two amplitudes; 

Ms values since 1968 measured from vertical seismograms. Absence of reported amplitudes on Milne 

seismographs (*) suggests Mp<6. 
184 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

TABLE 6.1. -Major California and Nevada arthqmikes, 1769-1989 

Dale and origin M Lu N. and Long W. Locality ML &-1 Ms MI M 

lime (Gw 

176907/2R 6? W 118" Los Aneelea BasinÃ 6 

-, . . ,. . . -. . . . . ..- 

1872/11/12 WOO 6 39"? 1171 

1873/11/23 0900 6% 42' 124' 

Los Angeles 

mywart fa,," 

San Francisco PenVal- Lake, B.C .- Los Angeles =p'on---------- 

Sanhcis0F.mim"b 

Great Fort Tejon earthquake------ 

Western Nevada or eastern Siena Nevada--- 

San Jo* **ionSari Bed&, region 

Southern Sam Cmz Mountains--------, 

East of Sari Francisco Bay 

Southern Sana Cruz Mountains-.------- 

Western Sim Joaquin Valley--------.- 

Virginia City, Nev 

myward hilt 

OIimglKiuse faull, Nev. 

owem valleyowens valley 

Imperial Valley (?)- 

Austin, Ne". legion (?)Ã‘ 

cmen, cit" ~ --~~--~-, 

Honey Luke-.. 

Imperial Valley la Colorado Rim delta (?)Ã 

hta - **- 

partiel* 

Westem sari Jonquil ValW 

Susanvfl&-- 

Bishop region 

San Jacinw or Elsinore fault region (?>---- 

pajm Gap 

Cape Men&.h 

Colorado River delta te&m------.-.-- 
6. EARTHQUAKE HISTORY, 1769-1989 185 

TABLE 6.1.-Major Calfomha and Nmwia earthguakes, l97%ms-Continued 

Dan and origin M Lal N. and Long W. Localily ML Mc-, Ms MI M 

time (Gl>TD 

San Jacimo or Elsinore fault region (?>---- 

Kc., canyonLyde Creek regionCape Mendocin0 regionEm of San Dice 

Souiheasiem Sierra New&------- 

Lyde Creek regionsan jacinio and Heel 

parkfield 

Colorado River delta region----.------ 

San Jnse~ - - ..... 

Great 1% eiÃ§lh<l 

Imperial Vdky 

DeBh V* legionÃ‘Ã 

Cape Mend~~& 

WiUonvUle 

Wea of Cape Maid&~----------- 

Glen Ivy Hot Spring+ 

west of EurekaÃ‘Ã‘ 

West of Eureka 

MO"~~ Bay 

Monterey BayWest of Cape Mlmino-------- 
186 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

TABLE 6.1.-Mojm CIllifomill and Nevada eartfuwafces, 1979-1989-Continued 

Date md origin M IAN. md Long W. Locality Mi. MM Ms Mr M 

lime (GMT~ 

1927Wl8 02:07 6 3790' 118W 

192l/11KM 1350 7.3 34'42' 12W48' 

193206/06 W44 6.4 40-45' 12493' 

1932/12/21 0630 72 38'45' 118- 

1933/01/05 0651 5.9 38-46' 11l044' 

Bishop - s,*- of - EmekflÃ ---- - - Ne" - Cedar Mountain. 5.7 

Laoina Salads, 6.5 

Colorado Rivet dclb---------- 7.1 

Colorado Rivet delm. --------- 6.0 

West of Cape Mendocin0 - 

BuckRid- 6.0,s.g 

west of ape ~~~d~~i~o~~ 6.0 

Caipen& 5.9 

Tom's Pl---.---- -. 5.3 6.0 

Tom*s p,- 6.0 

W- of Cape Men&no 6.4 

Maoix 6.2 

Desert Hot S- 6.5 

Ver& Mv 6.0 

West of - 59 

Pi"', Mountain 5.9 

Rain& main, NC" 

SliUwam Nev 

SdUwaOr Nev 

East of Sanm Tomas. 

East of Sanm Tomas. 
6. EARTHQUAKE HISTORY, 1769-1989 187 

TABLE 6.1.-Mcjw Califmia and Nevada earthquakes, 1979-1989-Continued 

Duelidorigin M LaN.andLmgW. Locality ML Ma-, Ms MI M 

lime (GMT) 

1956V2JJ9 1924 6.1 31-45' 11535- Sal MigwI. B.C.- 6.1 - - - 6.0 

195602/14 1833 6.3 31'30' 115'30' San Mi&. 63 - - - 6.2 

1956CM5 0120 6.4 31"30' 115'30' San MigwI. 6.4 - - - 6.2 

1956/KM1 16^8 6.0 40W 125"46' WestofCapMendocmo- 6.0 - - - - 

19W12/13 13-15 6.0 31' 115Â Western lihare. Golf of Cilifonia.--.--- 6.0 - - - - 

Dixie VtBey. Ncv.- -- 6.3 

muiz. NW 6.1 

Wea of Cape Mendodno~------- 6.2 

partteld 5.5.56 5.7 

Gulf at CaliiorniaÃ‘Ã‘Ã‘- 6.3 

Coyote- 5.8 

Imperial Valley 6.6,6.5,6d 

Livemu- 5.8 

Mammoth LateÃ‘------6.1.6.4.6 

Milmmodi Lakes 6.05.9 

-,,I, ~akes--.-. 6.1,6.6.6A 

Mammoth LateÃ‘- 61.6A. 5.8 

Vinoria. B.C 6.1.6.1 

West ofEd 69 

W~landÃ‘Ã‘Ã‘-Ã‘Ã‘Ã‘-Ã‘Ã‘ 5.6 

Nmh of Sanm Barbara Island---------. 5.3 

-oth - 5.9.6.0,59 

w6.7.6.1.6.1 6.0. 5.6 

Hill 63. 

190 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

earthquakes have been documented, some on the same 

fault strand, and the entii zone of plate-boundary 

deformation is exceptionally broad. M~or changes in 

fault strike also play a role in redistributing plateboundary deformation and diffusing it over a wider zone. 

Compressional bends intrcduce upliit, crustal thickening, 

and subsidiary reverse faulting, such as in the Transveme Ranges of southern California. Extensional bends 

are characterized by subsidence, basin filling, and, possibly, volcanism, as occurs in the Imwrial Valley. Extensional and compressional features are more localized 

at the smaller-wale diwontinuities and changes in fault 

strike that occur throughout the San Andreas system. 

Crustal movementa observed at the surface reflect 

deformation processes occurring at depth in the lithosphere. Both laboratory rock-mechanics experiments and 

studies of exhumed fault zones define the nature of these 

processes, which, in turn, constrain the classes of largescale faulting models consistent with surface measurementa. In the cml and brittle seismically active parts of 

the crust, elastic processes are dominant, the frictional 

strength of active faults increases linearly with depth, 

and faulting is controlled by Coulomb failure. The transition from brittle seismic behavior to ductile seismic 

deformation ms inthe midcrust. Although it is 

generally agreed that this transition occurs as a result of 

increasing temperature, its precise mechanism is weertain. If deformation in the midcrust is concentrated 

within a narrow vertical shear mne lying heath the 

seismically xtive fault plane, then the brittle/ductile 

transition may reflect either the increasing importance of 

ductile or cataclastic flow at depth (Sibson, 1982) or a 

thermally controlled transition from unstable to stable 

frictional sliding (Brace and Byerlee, 19m Tse and Rice, 

1986). However, if ductile deformation is broadly distributed in the midcrust, then the cyclic buildup and relief of 

stresses in the brittle seismogenic crust is controlled by 

the stress transfer between the elastic lithosphere and 

ductile "asthenosphere" and the flow properties of the 

latter. 

Both the steady, aseismic movements within the San 

Andreas plate-boundary wne and the mismic strain 

release in large earthquakes are well within the range of 

detectability of repeated geodetic-survey measurements. 

The purpose of this chapter is to summarize the salient 

features of these observations, demonstrating the constraints they place on the amount of presentday plate 

motion occurring across the San Andreas plate-boundq 

zone and showing how measmmenta shed light on the 

mechanics of the cycle of strain accumulation and release. 

The emphasis is necessady on movements close to the 

main strands of the San Andreas fault system, where 

observations are most numerow, although some networks extend as far as 100 km from the majorfaults. The 

measurements include triangulation, repeated observations of the angular separation of permanent survey 

markers, for which useful data date back tn about 12350, 

when gold was ht discovered in California; trilateration, repeated linelength measurements made by lwr 

ranging since about 197Q and local measurementa of 

aseismic fault slip made periodically or recorded continuously over apertures of about 10 to 100 m since about 

1960. 

OBSERVATIONS OF CRUSTAL DEFORMATION 

The fom here is on the spatiai and temporal patterns 

of interearthquake horizontal crustal movements in California that owe their origins to relative motion between 

the Pacific and North American plates, movements that 

supply the strain energy which is stored in crustal mks 

and ultimately released in large shallow-focus earthquakes. Observations of purely coseismic crustal deformation are not explicitly considered in this chapter; such 

movements are now well-understood consequences of slip 

on approximately the upper 10 to 15 h of vertical 

strike-slip faults. These models and their pdicted 

deformation patterns are discussed within the context of 

the entire earthquake deformation cycle in the next 

section. Readers interested in the coseismic movements 

observed for specific San Andreas earthquakes are ref& to the rep* by Lawson (1908) and Thatcher 

(1975) (1% San Francisco earthquake), Zhang and 

others (1W) (1940 El Centro earthquake), and Segall 

and Harris (1987) (I!%% Park6eld earthquake). 

Vertical crustal movementa can locally he substantial, 

at least when averaged over recent geologic time (see 

Yeats, 197t Goie, 1%). Deformation from reversefaulting earthquakes has also been well documented in 

seven4 events (for example, 1952 Kern County earthquake by Stein and Thatcher, 198c 1971 San Fernando 

earthquake by Castle and others, 197% 19B Coaliiga 

earthquake by Stein, 19B). Nonetheless, vertical movements are second-order features along most of the San 

Andrew fault system, and so they are not considered 

further in this chapter. 

Furthermore, in this chapter there is no review of 

measurement techniques, methods of analyzing and reducing data, or the mathematical and computational tmls 

used in modeling deformation processes. Interested readers are referred to the reports by Bomford (1980) and 

Savage and Prescott (1973) for descriptions of horiwntaldeformation-surveying methods and their precision, to 

those by Prewott (1976, 1981), Thatcher (1979), and 

Segall and Harris (1987) for discussions of data-analysis 

methods, and to the references cited below in the section 

entitled "Mechanics of Deformation" for details of the 

mathematical techniques used in model formulation. 
mtmnporary mtal nwvewnts in Cal@rnia are cowentrated within a plate-hndary defomtim zone that is typically 

50 to 200 km wide, centwed appmximkly on the San Andreas 

fault. Obseruatim of coseismic, postseismic, and inhs&ic 

movements define the earthquake defmation cycle and constrain 

models of strain accumulation and release fw strike-slip plate 

b~&ries. 

7. PRESENT-DAY CRUSTAL MOVEMENTS AND 

THE MECHANICS OF CYCLIC DEFORMATION 

time. ~&n~ this internal, several hundred kilometem of the earthquak~deformation cycle. 

righblateral offset has accumulated amss the San An- Several fault-zone features result in measurable deforCONTENTS 

vw 

Ob~e,."~ti~~ of -tal def,-tio" ..................... IW 

Shear strain on the San Andreas fault system---------- 191 

slip, inkgated di~placz-t rates, ad &%C. 

~~hh ~~&m PI& mti~n--------------------- 191 

Detailed dispIacemen&rate wtbme----------------- 193 

Me&,nies of d&,-tior,- ............................ ~XI 

Thid- and thin-lithosphere mod&------------------ 195 

S~li~nstitutive-law fault modele-------------- 201 

&mmq ........................................ zw 

Futm dhction8 for -~h ......................... 203 

R&erencw ................................... 2m 

INTRODUCTION 

CmM movements measured in California today Sample deformation processes that have continued through at 

least the past 5 Ma of Pliocene, Pleistocene, and Holocene 

plate-boundary faults, which are either freely slipping 

and without major seismic activity, or are in locked 

frictional contact and slip episdidly in repeated mat 

earthquakes. Aseismic fault slip (creep), as occurs across 

the San Andreas fault in central CaliforNa (fig. 7.1), 

causes no mstal deformation beyond a growing offset 

moss the fault, although this offset may be distributed 

across a wne as broad as a few tens or hundreds of 

metem. Whem the plateboundary hdt is alternately 

locked aseismically in its up~ 10 km or so and abruptly 

slipping in great earthquakes, deformation extends several tens of kilometem into the plate interiors. Between 

large events, elastic strains build up in this zone and are 

episodically released every few hundred years. Subsequent postearthquake recovery processes redistribute 

the strains aseismically for years to decades after a major 

shock, and this deformation @ually merges into the 

steady accumulation of elastic-strain energy that persists 

until the frictional strength of the fault is again exceeded. 

This sequence of interearthquake strain accumulation, 

coseismic strain release, and pastseismic readjustment is 

thus a recnrring process, here referred to inclusively as 

4 FIGURE 7.1.-A wall and sidewalk in HoUister, Calif., me ht and offset by mep on the Cahverae fault A dip rate of 6 to 6 rndyrchamkrkes 

much of the Cabera8 fault. View eatward dong north side of Sixth Street. Photopph by R.E. Wallace, U.S. Geological Survey. 

1% 

THE SAN ANDREAS FAULT SYSTEM. CALIFURNU: US. GEOLOGICAL. SURVEY PROFESSIONAL PAPER 1515 

dreas fault system, and many thousands of great earthquakcm similar to the historical events of IS7 and 19M 

have undoubtedly occurred. The obsewed deformation 

results from relative right-lateral translation of the 

Pacific and North American plates far from the main 

mation spread over an extremely broad plate-boundary 

wne. This deformation occurs where the San Andreas 

fault system comprises several subparallel splays, as in 

both the San Francisco Bay region and sonthem CaliforNa. There, hth aseismic slip and major strike-slip 
7. PRESENT-DAY CRUST& MOVEMENT3 AND THE MECHANICS OF CYCLIC DEFORMATION 191 

Horizontal interearthquake deformation is summarized below in rates of both displacement and shear 

strain. For both of these parameters, the componenb 

parallel to major active faults are the most significant and 

best illustrate the dominant pattern of presenbday 

tectonic movements, and so in this chapter these components are commonly shown exclusively. For example, 

although three independent tensor components are needed to completely characterize the horizontal-deformation 

field, in California the only significantly nonzero strainrate component is commonly the shear strain parallel tn 

the local trend of faults in the San Andreas system. Here, 

I consider only the component of maximum horizontal 

shear-strain rate, which, within observationa~ uncertainty, almost invariably parallels the San Andreas fault or 

one of its major strands. 

SHEAR STRAIN ON THE SAN ANDREAS FAULT SYSTEM 

Rates of contempomy shear strain are displayed in 

several complementary ways in figures 7.2 through 74 

details of each rate determination are summarized in 

table 7.1. Although only the magnitude of the maximum 

shear-strain rate is shown in each figure, the orientation 

of the maximum-horizontal-mntraction axis is listed in 

table 7.1. Note that for each of the strain rates shown in 

figures 7.2 through 7.4, aseismic fault slip contributes 

only negligibly, if at all, to the measured deformation. 

Further details on each strain field determination can be 

found in the references cited in table 7.1. 

Shear-strain rates peak at 0.4 to 0.6 pram (fig. 7.2) 

across the currently locked northern and southern sections of the San Andreas fault. Significant but slightly 

lower strain rates of 0.3 to 0.4 pradlyr are observed 

across right-lateral strike-slip faults in the northern 

Califomia Coast Ranges (north of lat 38' N.) east of the 

San Andreas fault, as well as across the San Jacinto fault 

in southern California. Shear-strain rates resolvably 

greater than zero are observed as far as about 80 km from 

the San Andreas fault itself. 

In addition, significant deformation is occurring across 

active faults in east-central California. In the mite 

Mountains, along the southern California-Nevada State 

line, small but resolvable strain rates (0.06*0.02 prad/ 

yr) have been measured, and the orientation of the strain 

field indicates crustal extension perpendicular to northsouth-striking normal faults in the area. Somewhat 

higher deformation rates are observed farther south, 

where right-lateral strain is murring parallel to the 

Owens Valley fault, site of the M=8 earthquake of 1872 

(see chap. 6). 

Shear-strain rate is plotted as a function of perpendicular distance from the San Andreas fault in figure 7.3. 

Deformation rates peak at the fault and decrease to half 

their maximums at a distance of about 30 km hm the 

fault. Most of the deformation is encompassed witbin a 

zone about 100 km wide centered on the fault ('â€ ̃Sa 

Andreas boundary deformation zone"), as discussed below. However, the reader may coniirm that this total lies 

in the range of about 30-40 Wyr by drawing a smooth 

curve through the data plotted in figure 7.3 and integrating this curve (that is, measuring and summing the total 

area underneath the curve) from -60 to +fX km. 

Maximum shear-strain rates at the Sari Andreas fault 

temd to be higher across the 1906 earthquake rupture in 

northern California (approx 0.6 prad/yr) than in southern 

Califomia (0.4 prad/yr), although the Carrim Plain data 

violate this generalization. Rather high deformation 

rates are also observed to 60 km east of the San 

Andreas fault in the northern California Coast Ranges. 

Shear-strain rates at various locations on the two 

currently locked sections of the San Andreas fault are 

plotted versus time since the most recent great earthquake at each ldty in figure 7.4. Most of these data are 

derived from triangulation measurements, many of which 

were first made in the late 19th or early 20th century. 

Thus, these determinations are much less precise than 

those listed in table 7.1 and plotted in figures 7.2 and 7.3, 

most of which are from the post-1970 period, Nonetheless, it is clear from figure 7.4 that deformation rates on 

the fault are much higher in the years to tens of years 

immediately after a great earthquake than they are later. 

Although it may be questionable to lump values from 

northern and southern California together on a single 

plot, the temporal decline in shear-strain rate shown in 

figure 7.4 depends only on about the first 70 years of data 

plotted, all of which come from the 1% rupture on the 

northern section of the San Andreas fault. 

ASEISMIC SLIP, INTEGRATED DISPLACEMENT RATIS, AND 

PACIFIC-NORTH AMERICAN PLATE MOTION 

Rates of surface aseismic slip (fault creep) at repre3entative pointa on the San Andreas fault system are 

Lsted in table 7.2 and plotted in figure 7.5. All fault 

Segments displaying measurable aseismic slip are repre- 

~nted, but the detailed distribution along each segment 

IS not shown; interested readers are referred to the 

references cited in table 7.2 for more details. Figure 7.6 

also displays the mtes of relative right-lateral displace 

ment integrated across geodetic networks of 50- to 

140-km aperture that span the San Andreas and related 

laults in seven areas of Califomia, for several of which 

:he detailed displacement-rate pattern is shown in figires 7.6 and 7.7. 

With the notable exception of the central, creeping 

3ection of the San Andreas fault, aseismic slip at the 

3urface represents only a very small proportion of the 
192 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

total right-lateral displaement across the San Andreas adjacent to the fault in this region, and so aU the relative 

fault svstem. On the 160-km-lone central section of the date motion taken up by the San Andreas system is here 

San a id re as fault, maximum fazb~reep rates average 

30 mdyr, close to the gecdetidy derived displacement 

rate of 3.32 1 mdy-r obtained over a 60-h aperture that 

spans the fault and the California Coat Ranges to the 

southwest. These data are the strongest evidence that no 

significant strain is accumulating in the crustal blceks 

being accommodated-birigid-block translation across the 

fault. Just north of this segment, on the southern section 

of the Calaveras fault, a signScmt amount of rightlateral slip at a rate of about 13 mdyr, occm a fault 

creep. Elsewhere in California, however, measured 

aseismic-slip rates range from 2 to 6 nun&, and creep 
7. PRESENTDAY CRUSTAL MOVEMENTS AND THE MECHANICS OF CYCLIC DEFORMATION 193 

commonly occurs only on restricted segments of otherwiselocked faults (for example, the Garlock and San 

Jacinb faults). 

The integrated righblateral displacement rates shown 

in figure 7.5 firmly constmin the proportion of PacificNorth American relative plate motion accommodated 

across the San Andreas fault system in California. In 

northern, central, and southern California, maximum 

rates range from 3.3 to 37 dyr. Global reconstructions 

of the motions of the major tectonic plates over the past 

3 Ma, as well as analyses of magnetic-anomaly lineation6 

at the mouth of the Gulf of California, point to a relative 

Pacific-North American platemotion rate of 49*3 mdyr 

(DeMets and others, 1987). The San Andreas fault 

system thus accounts for '70 to 80 percent of the relative 

plate motion, although the San Andreas fault itself does 

not everywhere take up all of this motion, and deformation is typically distributed across a boundary zone about 

100 h wide. 

Precisely how much additional relative plate motion i6 

accommodated across other faults in California is uncertain, although the amount is probably very little. According to Minster and Jordan (1987), very long baseliie 

interferometric WLBI) surveying results indicate that 

oblique extension of the Basin and Range province, 

directly east of California, is occurring at a rate of 10*2 

dyr with an orientation of N. 56' ?loo W. Dependiig 

)n the exact rate and orientation of this extension, as well 

IS on the precise direction of relative Pacific-North 

herican plate motion, the residual "missinf plate 

notion being accommodatd in California on faults other 

han those of the San Andreas system ranges from 

~egligibly small to possibly as much as I0 dyr. Thus, 

dthough the geodetic coverage in California is far from 

!omplete (see figs. 7.2, 7.5), all or most of the mne of 

!ign$cant plate-boundary deformation apparently has 

~een encompassed. 

DETNLED DISPLACEMENT-RATE PA'ITERNS 

Considerable detail on the distribution of deformation 

n the San Andreas boundary zone is provided by the 

-ather complete geodetic coverage available in the San 

hcisco Bay region and southern California. In the 

nethod wed to reduce these data, geodetic-line-length 

:hangea a~ used to determine station-displacement rates 

dative to a point at the center of gravity of the network. 

Fault-normal displacements are permitted by this meth- 

>d, but their values are minimized in the inversion 

DISTANCE FROM SAN ANOREAS FAULT, IN KILOMETERS 

FIGURE 1.3.-Shea~t.rainrate vemwp$qxndicuhrdistanc?hm the 

Sari Andm fault. Dot, determination fmm northem Wfomiq 

tide, determination from emthem Womia 1-w error bm shown 

for ~ference 

0 30 60 90 120 150 

TIME SINCE LAST GREAT EARTHQUAKE, IN YEARS 

FIGURE 1.4.-Sh~trainratevem8 time sincehtg~at earthquake 

@ears 1% and I857 in imeG heavy line indicate8 extent of fault 

mptm) on the Sa Andrea8 fault. CP, Cmko PM CV, hchelh 

Valley: FR, Fott Rmq M, Mojave; PA, Point Arena; PD, Palmdale: 

PR, Point Reyw SC, Sheiter Cove: SF, Sari fianci8co. htiom of 

hm northem Califomiq cirdea, determination8 hm southem 

Cdifoda; dWhd NNe, Uppl3xh8te fit ta data (Thatcher, 19s). 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

TABLE 7.1.-Shear-strain rates in Califmia 

[Maximum CompresBirn med from north, with clockwise directiw poaitive1 

Location Maximum 

Area compression Interval Reference 

 at N. ~ongw. direction* 0 

Round Valley .......................... Point Reyes ............................ Santa Rosa (Rodgers Creek 

f%,,l+j . - - .. , . The Geysers (Maacama fault). 

Napa (West Napa fault) ............ San Francisco ueninsula ............ 

Owens Valley .......................... Hollister (E. of Calaveraa 

fault). 

San Luis network: 

we~t ................................... C5rizo network: 

B ....................................... 

c ....................................... 

D ....................................... Tehachapi network: 

Garlock fault ........................ San Andreas fault .................... Cajon network ......................... Los Padres network ................. Barstow ................................. Palmdale ................................ Mojave network: 

Wl ..................................... W3 ..................................... Eastern ............................... Anm network: 

A ....................................... 

B ....................................... 

c ....................................... 

D ....................................... 

E ....................................... 

F ....................................... 

1942-64 Prescott (1985). 

1972-82 Prescott and Yu (1986). - - - - 

-- - - -- - - 1970-80 Ppeacott and othera (1981). 

1972-79 Savage and Lfaowaki (1984). 

1974-79 Savage and Ltoowski (1980). 

1971-78 Savage and others (1979). 

1977-83 N.E. King (unpub. data, 

1988). 

1977-83 N.E. King (unpnb. data, 

1983). 

1973-83 King and Savage (1984). - - - - 1974-84 Savage and othera (1986). - - 19794% King (1985). 

1971-80 Savage and others (1981). 

1934-82 Sauber and others (1986). - - - - - - - - 

process (see Prescott, 1981). Gross departures from this 

constraint would be revealed by notable disagreements 

between observed and predicted line-length changes, but 

no such discrepancies were found for the results presented here. Because the fault-normal displacement rates are 

small and show no consistent trends, they are not plotted 

on the profiles presented here. 

Displacement rates in the San Francisco Bay region 

are plotted in figure 7.6. The distribution of deformation 

varies considerably across the San Andreas boundary 

zone from north to south of the San Francisco Bay. In the 

north bay, the integrated right-lateraldisplacement rate 

across the network of 27Â± nunlyr (fig. 7.65) indicates 

that not all of the boundary zone has been captured 

within its 110-km aperture. Within about 5 km of the San 

Andreas fault, rapid change in the gradient of deformation rate indicates that interearthquake strain is concentrated close to the fault. Outside this near-fault region, 

deformation southwest of the fault appears to be negligible. Northeast of the fault, however, the persistence of 

significant movements right to the edge of the profile 

suggests that the 5- to 10-mm/yr deficit in boundary-zone 

deformation across this profile is being accommodated to 

the east of the Green Valley fault. Across the central and 

south bay (fig. 7.60, movements are more evenly 

distributed through the network, and the integrated 
7. PRESENT-DAY CRUSTAL MOVEMENTS AND THE MECHANICS OF CYCLIC DEFORMATION 

TABLE 7.2.-Representative aseismic-81ip mlw on f&of the San Ad- system 

Fault 

Hqyward ................................ Northern section 

of the Calaveras. 

Southern section 

of the Calaveras. 

San Andreas ........................... Do.. .................................... Do.. .................................... Gariock .................................. San Andreas ........................... San Jadnto 

(Coyote Creek strand). 

Imperial ................................. 

Site MN. LwW. (z) Reference 

Hayward network .................... 37.6" E2.1' 6 Preacott and Lisowski (1983). 

Camp Parks ............................ 37.7" 121.9" 3 Do. 

Sail Felipe .............................. 37.0" 121.5" 13 Lisowski and Preacott (1981). 

Cienega Wiiery ....................... 36.7" 121.5- 13 Schuiz and others (1982). 

Eade Ranch ............................ 36.4" 121.0" SO Buford and Hmh ll'ISO). 

Partfleld (Diirham Ranch) ......... 35.9" 120.4" 13 Do. 

Cameron ................................ 35.1" 118.3' 4 Louie and others (1985). 

India Hilln .............................. 33.7" 116.2" 2 Do. 

Bailey's Wdl ........................... 33.0" 116.0" 5 Do. 

Intentate Highway 80 .............. 32.8' 115.5" 5 Do 

displacement rate of 37Â± mm!yr across the south bay 

suggests that the entire boundary zone has been 

spanned. Closer examination of the profile, however, 

reveals several zones of locally high deformation gradient, one across the San Andreas fault, where it resembles 

that observed near the fault in the north bay. In addition, 

rapid changes in the profile across the Hayward and 

Calaveras faults reflect aseismic slip at rates of 3 to 6 

mm/yr on these faults (see fig. 7.5). 

In southern California (fig. 7.7), deformation across the 

San Andreas boundary zone notably broadens from the 

Salton Sea, in the south, northwestward to the Big Bend 

region of the San Andreas fault north of Los Angeles (see 

fig. 7.5). At the south end of the Salton Sea, all of the 

boundary-zone deformation, 35Â± mm/yr, occurs within 

an area about 50 km wide (profile S, fig. 7.70 that 

rapidly broadens to more than 100 km wide north and 

west of the Salton Sea (profile N, fig. 7.70 and, possibly, 

broader still by about 60 ton farther northwest (fig. 

7.7B). North and west of Los Angeles, networks of 

1- aperture capture only 18Â± mm!yr of the total 

right-lateral-displacement rate (fig. 7.5). The profiles in 

figures 7.7B and 7.7C also show that in contrast with the 

northern section of the San Andreas fault, deformation 

gradients across the fault are smoother, and deformation 

is not so closely concentrated near the fault. 

MECHANICS OF DEFORMATION 

The observations described in the previous section 

point to a range of mechanical behavior for the faults 

comprised by the San Andreas system, from freely 

sliding with only minor accompanying seismicity, to 

completely locked from the surface to seismogenic depths 

except for abrupt slip during infrequent great earthquakes. 

On the 160-km-long central section of the San Andreas 

fault, virtually all fault slip occurs aseismically. Slip rates 

measuredat or near the fault are close to the average rate 

for the entire San Andreas boundary zone (fig. 7.61, no 

strain is detectable in the crustal blocks adjacent to the 

fault, and historical earthquakes of Ms5Yz have not 

occurred. Abundant minor seismicity (see fig. 4.10) 

contributes only negligibly to the slip budget, and except 

for a few small patches of fault surface that are in 

fictional contact between these small earthquakes, the 

first-order steady-state model for this segment involves 

rigid translation of adjacent fault blocks across the freely 

sliding plane of the San Andreas. 

A transitional behavior applies to those fault segments 

where steady-state fault creep is observed at the surface 

but historical or prehistoric earthquakes of M26 have 

been documented. Examples include the Parkfield and 

Coachella Valley segments of the San Andreas fault, the 

Hayward fault, and the Imperial fault. On these segments, during the interseismic phase of the earthquake 

cycle, the fault is inferred to be ireely slipping in its upper 

few kilometers, in locked frictional contact at seismogenic 

depths (approx 3-10 km), and once again freely slipping 

at greater depths (fig. 7.8A). The result of this slip 

distribution is interseismic creep at the surface fault 

trace and elastic-strain accumulation in the adjacent 

blocks (figs. 7.8B. 7.80. 

THICK- AND THIN-LITHOSPHERE MODELS 

The most extreme features of locked fault behavior are 

currently observed on the two San Andreas fault segments where great earthquakes have occurred in historical time, in 1857 and 1906 (see fig. 5.11 for locations and 

coseismic-slip distributions). On these segments, no 

aseisinic slip is observed at the Earth's surface, the two 

faces of the fault are in locked frictional contact to depths 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

of 10 to 15 km, and interearthquake slip is either 

extremely small or absent. At greater depths, the 

mechanics of fault movement is uncertain, but two 

models bound the range of expected behavior (fig. 7.9). In 

the first, the thick-lithosphere model, the depth D of 

coseismic faulting is much less than the thickness H of 

elastically strong lithosphere. Interearthquake defomation then predominantly results from episodic or steady 

aseismic slip on the downward extension of the seismogenic fault zone, and any effects of the underlying weak 

asthenosphere can be safely neglected. In the second, the 

thin-lithosphere model (fig. 7.9), coseismic faulting depth 

is comparable to elastic-plate thickness. In this model, 

transient postseismic and steady interseismic flow in the 

PACIFIC 

OCEAN 

u 

100 KILOMETERS 

FIGURE 7.6.-Sketch map of Wornia, showing rates (red strands of the San Andreas fault system. Values inmillimeters 

numbers) of aseiamic slip (fault creep) and relative right- per year. Locked (no surface slip) segment of major fault, red 

lateral-displacement rates (blue numbers) near arrows, which line; other Quaternary fault, black line; faults dotted where 

indicate direction of relative movement along major active concealed. See table 7.2 for details. 
N. 33' W. VELOCITY. 

IN MILLIMETERS PER YEAR 

8 A A 

8 

SAN ANDREA5 

FT n/Ã ^Iff""-;= 10 

"AWARD 

FAULT 

8 
198 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

I 

100 50 0 50 100 

N. 51" E. DISTANCE. IN KILOMETERS 

FIGURE 7.7.-Displacement rates in southern California. A, Relative right-lateral displacement rates determined from repeated 

geodetic-survey measurements during 1972-87. Error ellipses 

show 95-~ercent-confidence limits for each determination (Prescolt and others. 19871 B, Relative station velociueii parallel to 

approximate trend of the San Andrean fault (N. 39 W )plotted 

30 I I 

75 50 25 0 25 50 75 

N. 50" E. DISTANCE. IN KILOMETERS 

against distance perpendicular to this trend, for stations in 

northern part of map in figure 7.7A. l-Ã§ error bars are indicated. 

Pemendicular velocity comoonent is neeliciile and is not plotted 

here. C, same as fig& 7.7~ for stations largely to north (N) and 

south (S) of the Salton Sea. 
7. PRESENT-DAY CRUSTAL MOVEMENTS AND THE MECHANICS OF CYCLIC DEFORMATION 199 

asthenosphere provide the dominant mechanism for interearthquake strain accumulation. 

Note that in the context of these two models, the terms 

"lithosphere" and "asthenosphere" are linked to mechanical properties of the Earth's crust and upper mantle: The 

lithosphere is the strong elastic layer near the Earth's 

surface, and the asthenosphere is the region of ductile 

flow that lies beneath. Their boundary may thus lie well 

above the thermal boundary layer that separates the 

moving plates from the convecting mantle. If so, then at 

least the upper part of the "asthenosphere" forms part of 

the tectonic plate and moves with it. 

Displacement rate 

I * Distance 

Shear-strain rate 

FIGURE 7.8.-Elastic-W-space model showing fault creep at surface, 

locked (nonslipping) fault at depth, and freely sliding zone at great 

depth (A). Displacement rate (B) and strain rate (0 are plotted 

against distance from fault. 

Thus considered, the boundary between "lithosphere" 

and "asthenosphere" defines the zone of decoupling 

between surface tectonic processes and those that occur 

in the ductile region beneath. The location of this 

boundary is thus of central importance to the broad-scale 

tectonics of the San Andreas fault, the nature of the 

earthquake-generation process and its thennomechanical 

implications (see chap. 9), and the relation between 

8hallow structural features and those inferred at depth 

(see chap. 8). I explore below the influence of this 

boundary location on cyclic earthquake-related deformation at the currently locked transform fault zones in the 

San Andreas, illustrating the contrasting mechanical 

behavior of the thick- and thin-lithosphere models. 

All of the models considered here are two dimensional, 

and so neither slip nor mechanical properties vary along 

fault strike. Each model consists of only a single planar, 

vertical strike-slip fault. However, because the medium 

properties are linear elastic and (or) viscoelastic, the 

effects of multistranded fault zones can be obtained by 

simply superposing the deformation due to slip on individual fault segments. Furthermore, all of the twodimensional models discussed here have also been 

considered in three dimensions, and so complexities 

arising from changes in fault strike, variations in slip 

along strike, and the finite extent of faulting can be 

incorporated straightforwardly as necessary. Similarly, 

except for the transition from elastic lithosphere to 

viscoelastic asthenosphere, depth variations in material 

properties are not considered, although, again, solutions 

have been obtained for faulting in plane-layered elastic 

THICK- THINLITHOSPHERE MODEL LITHOSPHERE MODEL 

D/H <-<- 1 D/H =I 

FIGURE 7.9.-Thick- and thin-lithaanhere models. D. denth of eoseismic 

faulting; H. thickness of elasticalli strong lithoephe;. Small arrows 

alone fauh and larger arrows indicate direction of relative movement 

of fault and tectonic plates, respectively, 
200 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

and viscoelastic media. Although fault end effects and 

changes in slip and geometry along fault strike can be 

locally important, these effects, as well as those due to 

depth-varying material properties, are generally second 

order relative to the deformation features described 

here. More important are the effects of the several 

subparallel strands that compose much of the San Andreas fault system along its two currently locked sections. In these sections, the interseismic deformation due 

to each major fault strand contributes significantly to the 

observed displacement pattern, and as a rule the effects 

of faults lying off the San Andreas proper cannot be 

safely ignored in matching models to data across the 

entire San Andreas boundary zone. 

The simplest form of the thick lithosphere model, first 

proposed by Savage and Burford (1970), is illustrated in 

figure 7.10. In this idealized model, interearthquake 

strains accumulate uniformly throughout the deformation cycle and have precisely the same spatial pattern as 

coseismic strains, except that the sense of movement is 

reversed. The cycle consists of coseismic slip Au extending from the surface to depth D and steady interearthquake aseismic slip at a constant rate ti (=AdT, where T 

is the earthquake recurrence interval) beginning at z=D 

and extending to great depth. For this model of interearthquake deformation, simple expressions relate 

COSEISMIC INTERSEISMIC 

STRAIN RELEASE STRAIN ACCUMULATION 

DISLOCATION 

MODEL 

Slip An 

Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘i 

fp rate Au/T 

FIGURE 7.10.-Elastic-half-spaee model for earthquake cycle. Au, 

coseibmic slip; D, depth of coseismic slip; T, earthquake recurrence 

interval: u,. horizontal disvlacement parallel to fault: duJdy. shearstrain component parallel to fault; y,distance from fault. 

surface-displacement rates u,(y) and shear-strain rates 

bZyfy) to the fault parameters ti and D and the distance y 

from the fault trace: 

d 

and = - [*.(vJ\ = - - dv 

A principal utility of this model is the ease with which 

approximate values of displacement and strain rate can 

be computed, commonly as a preliminary step to more 

detailed computations that employ complex models which 

nonetheless show many of the same general features. For 

example, using typical San Andreas values of Au=4 m, 

T=200 yr, and D=15 km, then ti-20 mm/yr, and the 

engineering shear-strain rate (twice the tensor strain 

rate bZy) at the fault trace (y=O) is about 0.8 Wyr, a 

value dose to some of the peak strain rates plotted in 

figures 7.2 and 7.3. Furthermore, the bell-shaped distribution of secular strain across the model fault (middle 

right, fig. 7.10) generally accords with observations (fig. 

7.3), and the width of the profile is a direct measure of the 

fault-locking depth D. (Note, however, that the observations summarized in figure 7.3 include strain rates 

determined from multistranded segments of the San 

Andreas fault system, and so they are not directly 

comparable to the model calculations for a single fault 

strand illustrated in fig. 7.10.) Recalling the observations 

discussed in the section above entitled "Observations of 

Crustal Deformation," the wider zone of secular strain 

across the southern section of the San Andreas can be 

rationalized if the depth of seismic slip and, thus, the 

locking depth of the fault are simply greater in southern 

than in northern California. As it stands, this model has 

no transient effects and so is too simple to explain the 

postearthquake strain changes plotted in figure 7.4. 

However, introducing a rather straightforward modification remedies this defect while accounting for the 

observed difference in strain-rate profiles between northem and southern California. Surprisingly, these same 

features are, for different reasons, natural consequences 

of the thin-lithosphere model. 

The two models are illustrated in figure 7.11. In the 

thick-lithosphere model, postseismic effects are introduced by specifying transient postearthquake slip on a 

segment of the fault immediately beneath the coseismic 

rupture. Its time history is constrained by an exponentially decreasing slip rate (time constant a), and its 

magnitude by the requirement that the cumulative slip 

sum to the coseismic offset Au by the end of the cycle. In 

the thin-lithosphere model, the transient deformation 
7. PRESENT-DAY CRUSTAL MOVEMENTS AND THE MECHANICS OF CYCLIC DEFORMATION 201 

results from flow in the asthenosphere due to stress 

relaxation after seismic faulting in the lithosphere. Its 

time scale is controlled by the asthenosphere-relaxation 

time ~=2q/p, when q is the effective viscosity of the 

asthenosphere and u, is the average shear modulus 01 

lithosphere and asthenosphere, here taken to be equal. In 

both models, the transient motions are superimposed on 

a steady component of deformation that is due to relative 

plate motion. 

Detailed computations show that the two models 

produce surface deformations that with suitable choices 

of model parameters are observationally indistinguishable (see Thatcher, 1983). Here, the discussion is restricted to qualitative features, as summarized in figure 

7.12. Near the fault, shear-strain rates monotonically 

decrease over time and gradually approach a constant 

(fig. 7.125). while the deformation profile broadens and 

strains diffuse into the interiors of the adjacent plates as 

the cycle progresses (fig. 7.124. It is easy to match the 

observed temporal decline in strain rate with either 

model; the particular parameter combinations are themselves not unique, and a range of choices can provide 

equally good agreement. All satisfactory thin-lithosphere 

models, however, require an elastic plate only 10 to 15 km 

thick, the maximum depth of coseismic slip in the 1906 

earthquake (Thatcher, 1975). Both models predict a 

broadening of the zone of deformation that depends on 

the time interval since the latest great earthquake, and 

SO the greater width of the strain-rate profile in southern 

California can be accounted for. For example, data from 

the northern, locked section of the San Andreas fault may 

correspond to times t, to t, in figure 7.12, whereas those 

from the southern section may correspond to times ts and 

t6. 

THICK-LITHOSPHERE 

MODEL 

THIN-LTHOSPHERE 

MODEL 

Lithosphere ----------- 

Asthenosphere 

FIGURE 7.11.-Sdc features of thick- and thin-lithosnhere models. ~ ~ 

R, depth of cotteismic slip; Do, depth to bottom of zom of transient 

postficismic slip; H. thickness of elastieaUy Strong lithosphere: q, 

effective viscosity of asthenosphere; u., average shear modulus of 

lithosphere and asthenosphere. 

More complex models that combine features of both the 

thick- and thin-lithosphere models are also consistent 

with available data (for example, Li and Rice, 1987). 

Furthermore, coseismic and interearthquake fault slip 

undoubtedly vary as a function of depth, rather than 

abruptly terminating at some specified fault depth. 

Although this gradationality of the slip' distribution 

modifies the detailed patterns "of surface strain and 

displacement from those illustrated in figure 7.10, for 

example, the same qualitative features are preserved, 

and it will he difficult to distinguish between differing 

slip-depth distributions on the basis of surface-deformation measurements alone. 

In summary, at transform plate boundaries, available 

data are consistent with both thick- and thin-lithosphere 

models but do not strongly constrain either. The most 

geophysically interesting feature of both models is the 

predicted postearthquake diffusion-like spread of strain 

from the plate-bounding fault into the interiors of the 

adjacent plates. Postearthquake surveys, however, are 

sufficiently infrequent and areal coverage sufficiently 

limited that these effects, if they indeed occur, have not 

been directly observed. Details of the temporal decline in 

deformation rate near the fault are also absent. 

STRESS-SUP-CONSTITUTIVE-LAW 

FAULT MODELS 

A completely different class of large-scale-faulting 

models are now being developed to more realistically 

incorporate the fault-failure process into the earthquake 

deformation cycle (for example, Stuart, 1979; Tse and 

Rice, 1986). Instead of specifying slip on the platebounding fault, these models extrapolate from laboratory 

A B 

SHEAR-STRAIN 

RATE 

t 

+ 

2 

a 4 = ", '4 r, 

+ TIME (t,-t.) t 

DISTANCE FROM FAULT EQ EQ 

FIGUBE 7.12.-Complete earthquake-cycle model prediction^ for thick. 

and thin-lithosphere models. A, Shear-strain rate versus distance 

from fault and its temporal evolution through deformation cycle. B, 

Shear-strain rate on fault versus time for one complete cycle. EQ, 

earthquake. 
202 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

observations of the timedependent frictional properties 

of rocks (for example, Dieterich, 1979; Tullis, 1988) to 

assign slipstress-constitutive laws to the fault surface. 

As remotely applied stresses increase, each segment of 

the fault slips at a rate that depends on both the previous 

slip history and the current applied stress, and so the 

cycle of elastic-strain accumulation and release can be 

simulated. By specifying the depth dependence of fault 

frictional properties, a slip behavior nearly identical to 

tbat of the thick-lithosphere model (fig. 7.11A) follows 

naturally. A sample calculation of this type (Tse and Rice, 

1986) illustrates the method and shows typical results 

(fig. 7.13). After a large coseismic slip event in approximately the upper 10 km of the model fault, transient 

postseismic slip occurs on both the coseismic fault plane 

and its downdip extension. As slip rates decline to near 

zero on the shallow segments of the fault, interseismic 

slip at greater depths approaches nearly steady-state 

values. Finally, near the end of the cycle, the constitutive 

model predicts an increase in slip rate on the shallow 

coseismic fault segment before the next large slip instability ("earthquake"). 

Although the appropriateness of this extrapolation of 

laboratory results to large-scale faulting is a matter of 

current debate and the scaling of laboratory parameters 

to the field is uncertain, Tse and Rice's calculations 

clearly demonstrate tbat the principal observed features 

of the earthquake deformation cycle on the San Andreas 

fault can be reproduced by such models. Ongoing laboratory studies should refine and modify the stressslip-constitutive laws, and geodetic and continuous 

strain-monitoring observations of preearthquake and 

postearthquake crustal deformation can test the applicability of these postulates to large-scale faulting processes. 

SUMMARY 

Contemporary crustal movements in California are 

concentrated within a plate-boundary-deformation zone 

that is typically 50 to 200 km wide, approximately 

centered on the San Andreas fault. Integrated rightlateral displacement rates across this zone range from 33 

to 37 mm/ yr, representing about 75 percent of the 

Pacific-North American relative plate motion. Most or all 

of the rest may be taken up east of the San Andreas fault 

system in the Basin and Range province. Although 

aseismic fault slip (creep) is a locally important component of this relative plate motion, most of the geodetically 

measured deformation represents elastic strain on the 

crustal blocks adjacent to faults of the San Andreas 

system. Rates of secular (interseismic) shear strain are a 

maximum on the two currently locked segments of the 

San Andreas fault, sites of the great 1867 and 1906 

EXPLANATION 

Location Time interval 

AwB 2.4 yr 

B to C 4.6 hr 

CwC Instantaneous 

C to D 4 t 

D to E 3.5 min 

E to F 77.8 days 

F to G 3.6 yr 

G in H 77.2 yr 

tm+ 

CUMULATIVE SLIP. IN METERS 

FIGURE 7.13.-Cumulative slip versus depth for selected time intervals through deformation cycle in quasi-static faultinstability model modified 

from Tse and Rice (1986). T, recurrence interval. 
7. PRESENT-DAY CRUSTAL MOVEMENTS AND THE MECHANICS OF CYCLIC DEFORMATION 203 

earthquakes. Values range from 0.4 to 0.6 parts pel 

million per year (ppm/yr) at the fault to 0.1 ppdyr 30 tc 

80 km from it. Deformation occurring at the times ot 

large strike-slip earthquakes (coseismic strain) is concern 

trated within a few tens of kilometers of the surface fault 

rupture, indicating that earthquake fault slip is largely 

confined to the upper 10 to 15 km of the crust. Aftel 

major events, postseismic shear strain occurs at transiently high rates (more than 2 ppm/yr) that decay tc 

background interseismic rates over a time scale of years 

to tens of years. 

Ohsewations of coseismic, postseismic, and interseis. 

mic movements define the earthquake deformation cycle 

and constrain models of strain accumulation and release 

for strike-slip plate boundaries. Observations are fitted 

equally well by two contrasting models. In the first 

model, the depth of coseismic faulting is much less than 

the thickness of the elastically strong lithosphere, and 

pstseismic and interseismic deformation result from 

transient and steady aseismic slip on the downward 

extension of the earthquake fault plane. At the other 

extreme, if earthquake slippage extends through most or 

all of the elastic lithosphere, interearthquake deformation is due to transient or steady flow in the underlying 

weak substrate ("asthenosphere"). 

FUTURE DIRECTIONS FOR RESEARCH 

Although the broad outlines of current movement 

across the San Audreas boundary zone are now known 

and the main features of the cyclic deformation expected 

from great strike-slip earthquakes have been delineated, 

many issues still remain to be explored. Although all of 

the relative Pacific-North American plate motion occurring across California may have been measured geodetically, this determination is not yet definitive, and as 

much as 10 mm/yr of motion may be accommodated east 

or west of the approximately 100-km-wide zone defined 

by current measurements. Furthermore, the thickness of 

the elastically strong crust is uncertain by at least a 

factor of 3, and so major alternative models of the 

earthquake deformation cycle cannot be distinguished 

(fig. 7.9). Because surface-deformation observations cannot themselves resolve this ambiguity, other data, possibly gravity-field observations and lithosphericdeflection models (for example, McNutt, 1980), are 

needed. 

Few details exist on the preseismic and postseismic 

movements related to large plate-boundary earthquakes. 

Whether detectably anomalous crustal movements precede large earthquakes is uncertain. Theoretical models 

and fragmentary observations suggest that precursory 

slip may occur on or beneath the eventual coseismic 

rupture plane. However, except for the ohsewation that 

premonitory deformations must be small relative to 

coseismic movements (for example, Johnston and others, 

19871, precursory slip is otherwise unconstrained. Existing data are sufficient to demonstrate that postseismic 

movements, at least those from great earthquakes, are 

large-commonly, 10 to 30 percent of the coseismic 

deformation (Thatcher, 1984)-but the time scale and 

spatial distribution of these motions are not well determined at strike-slip plate boundaries. Laboratory experiments on lower-crustal rock types suggest that their 

ductile behavior is not approximated well by linear 

viscoelasticity, as assumed in the thin-lithosphere model, 

but postseismic observations are not yet sufficiently 

detailed to confirm this expectation. 

Furthermore, vertical crustal movements in California 

are not well understood. Though not dominant in California's largely strike-slip-faulting environment, vertical 

movements can nonetheless be locally important in such 

regions as the Los Angeles and Ventura Basins, the 

Transverse Ranges of southern California, and the Cape 

Mendocino area of northern California. Current and 

future work that integrates geologic and geodetic information in these regions should begin to shed light on 

long-term, secular vertical-movement patterns and their 

origins. 

Within complex, multistranded fault zones and, possibly, in simpler regions, permanent inelastic deformation 

of upper-crustal rocks may contribute signif~cantly to the 

current movement pattern. For example, at subduction 

boundaries, geologic and geodetic observations indicate a 

substantial imbalance between cumulative interearthquake strain and coseismic strain release, commonly 

reflected in long-term uplift or subsidence of coastal and 

inland regions. However, at such predominantly transcurrent boundaries as the San Andreas, the observable 

effects of inelastic strain are more subtle. The thermal 

consequences of such deformation may be the most direct 

evidence for inelastic strain (see chap. 9). However, for 

California at least, the available data are either contradictory or ambiguous, and the extent to which measured 

interearthquake movements release elastically stored 

strain is currently unresolved. 

REFERENCES CITED 

Bomford, Guy, 1980. Geodesy (4th ed): London, Oxford University 

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Brace, W.F., and Byerlee, J.D., 1966. Stickdip a; a mechanism for 

earthquakes: Science, V. 153, no. 3739, p. 990492. 

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Burford, R.O., and Harsh, P.W., 1980, Slip on the Ban Andrew faultin 

central California from alinement array surveys: Seismological 
204 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

Society of America Bulletin, v. 70, no. 4, p. 1223-1261. 

Castle, R.O., Church, J.P., Elliott, M.R., and MorriÃˆon N.L., 1975, 

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Fernando earthquake of February 9, 1971: A summary: Teetom 

physics, v. 29, p. 127-140. 

Delieta, Charles, Gordon, R.G., Stein, Seth, and Argus, D.F., 1987, A 

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Dieterich. J.H.. 1979. Modelling of rock friction. 1. Ex~erimental 

results and constitutive equations: ~oume~ of Geophysical REsearch, V. 84, no. B5, p. 2161-2168. 

Johnston, M J.S., Linde, A.T., Gladwin. M.T., and Boxherit, R.D., 

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Lisowski, Michael, and Preacott, W.H., 1981, Short-range distance 

measurements along the San Andreas fault system in central 

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Louie, J.N., Alien, C.R., Johnson, D.C., Haase, P.C., and Cohn, S.N., 

1985. Fault ali~ in southern California: Seismoloirical Society of 

America ~ullek, v.75, no. 3, p. 811-833. - 

McNutt, Marcia, 1980, Implications of regional gravity for state of 

stress in the earth's crust and upper mantle: Journal of Geophysical Research, v. 85, no. Bll, p. 63774396. 

Minster, J.B., and John, T.H., 1987, Vector cwtralnta on Western 

U.S. deformation from soace eeodeav. neotectonics. and plate 

motions: Journal of Geophysi& &arch, v. 92, no. ~6, p. 

4798-4804. - . . . -. - . 

Prescott, W.H., 1976, An extension of Frank's method for obtaining 

crustal shear strains from survey data: Seismological Society of 

America Bulletin, v. 66, no. 6, p. 1847-1853. 

- 1981, The determination of displacement files from geodetic data 

along a strike slip fault: Journal of Geophysical REsearch, v. 86, 

no. B7, p. 6067-6OT2. - 1985, An overview of the distribution of relative plate motion 

along the San Andreas fault system from Hollister, California, to 

the Mendocino triple junction, in Shearer, C.F., Minutes of the 

National Earthquake Prediction Evaluation Council, July 26-27, 

1985, Medo Park, California: U.S. Geological Survey Open-File 

Report 85-754, p. 232-246. 

Prescott, W.H., and Lisowski, Michael, 1983, Strain accumulation 

along the San Andreas fault system east of San Francisco Bay, 

California: Tectonophysics, v. 97, no. 14, p. 41-56. 

Prescott, W.H., Lisowski, Michael, and Savage, J.C., 1981, Geodetic 

measurement of crustal deformation on the San Andreas, Hayward, and Calaveras faults near San Francisco, California: Journal 

of Geophysical Research, v. 86, no. Bll, p 10853-10869. 

- 1987, Velocity field along the San Andreas fault in southern 

California[abs.]: Eos(American GeophysicalUnionTransactions), 

V. 68, no. 44, p. 1506. 

Preacott. W.H.. and Yu. S.-B.. 1986. Geodetic measurements of 

horizontal deformation in the northern San Francisco Bay region, 

California: Journal of Geophysical Research, v. 91, no. B7, p. 

7475-7484. 

Sauber, Jeanne, Thatcher, Wayne, and Solomon, S.C., 1986, Geodetic 

measurement of deformation in the central Mojave Desert, Califomia: Journal of Geophysical Research, v. 91, no. B12, p. 

12fia8-1m. .-... ~ 

Savage, J.C., and Burford, R.O., 1970, Accumulation of tectonic strain 

in California: Seismological Society of America Bulletin, v. 60, no. 

6, p. 1877-1896. 

Savage, J.C., and Lisowski, Michael, 1980, Deformation in Owens 

Valley, California: Seismological Society of America Bulletin, v. 

70, no. 4, p. 1225-1232. - 1984. Deformationin the WhiteMountainseismicgap, CaliforniaNevada, 1972-1982: Journal of Geophysical Research, v. 89, no. B9, p. 7671-7687. 

Savage, J.C., and Prescott, W.H., 19% Precisionof gdolite distance 

measurements far determining fault movements: Journal of Geophysical Roeah, v. 78, v. 26, p. 6001-CT08. 

Savage, J.C., Prescott, W.H., and Gu. Guohua, 1986, Strain accumulation in southern California 1973-1984: Journal of Geophysical 

Research, v. 91, no. B7, p. 7455-7473. 

Savage, J.C., Prescott, W.H., Lisowski, Michael, and King, N.E., 

1979. ~. Geodnlite measurements of deformation near Hollister. 

California, 1971-1978: ~ournal of Geophysical Research, v. 84, no. 

BlS, p. 7599-7615. 

- 1981, Strain accumulation in southern California, 1973-1980: 

Journal of Geophysical Research, v. 86, no. B8, p. 6991-7001. 

Schulz. S.S.. Mavko. G.M.. Biirford. R.O.. and Stuart. W.D.. 1982. 

L&-t& fault creepobservations incentral ~alifornia: JOU~ 

of Geophysical Research, v. 87, no. B8, p. 69774982. 

Segall, Paul, and Harris, R.A., 1987, Earthquake deformation cycle on 

the San Andreas fault near Parkfield, California: Journal of 

Geophysical Research, v. 92, no. B10, p. 10511-10526. 

Sibaon, R.H., 1982, Fault zone models, heat flow, and the depth 

distribution of earthquakes in the continental crust of the United 

States: Seismological Society of America Bulletin, v. 72, no. 1, p. 

151-163. 

Stein, R.S., 1983, Reverse slip on a buried fault during the 2 May 1983 

Coalinga earthquake: Evidence from geodetic elevation changes, 

in Bennett, J.H., and Sherburne, R.W., eds., The 1983 Coalinga, 

California earthquakes: California Division of Mines and Geology 

Special Publication 66, p. 151-163. 

Stein, R.S., and Thatcher, Wayne, 1981, Seismic and aaeismic deformation associated with the 1952 Kern County. California, earthquake and relationship to the Quaternary history of the White 

Wolf fault: Journal of Geophysical Research, v. 86, no. B6, p. 

49134928. 

Stuart, W.D., 1979, Strain softening prior to two-dimensional strike 

slip earthquakes: Journal of Geophysical Research, v. 84, no. B3, 

p. 1063-1070. 

Thatcher, Wayne, 1975, Strain accumulation and release mechanism of 

the 1906 San Francisco earthquake: Journal of Geophvaical Re 

Midi. v. 80. no. 35. D. 4862-4872. 

- 1979, Systematic inversion of geodetic data in central California: 

Journal of Geophysical Research, v. 84, no. B5. p. 2283-2295. 

- 1983, Nonlinear strain buildup and the earthquake cycle on the 

San Andreas fault: Journal of Geophysical Research, v. 88, no. B7, 

p. 58934902. 

Tse. S.T.. and Rice, J.R., 1986, Crustal earthquake inatability in 
7. PRESENT-DAY CRUSTAL MOVEMENTS AND THE MECHANICS OF CYCLIC DEFORMATION 205 

relation to the depth variation of frictional Blip properties: Journal Yeats, R.S., 1977, High rates of vertical crustal movement near of Geophysical Research, v. 91, no. B9, p. 9452-&172. Ventura, California: Science, v. 196, no. 4287, p. 295-298. 

TuUis, T.E., 1988. Rock friction constitutive behavior from laboratorv Zhanc. Y.. Thatcher. Wavne. and Snav. R.A.. . 1888. . Coseiamic sli~ . in - 

experimentsand its implications for an earthquake preriicHon field the 1940and 19791rn;rial valley earthquakesand its implications 

~no~ioring progmm: Pure and Applied Geophysics, v. 126, no. 1abs.l: Em (American Geophysical Vnion Transactio MB). v. 69, p. 
EXPLANATION 

Pre-late Miocene crystalline 

basement 

Late Miocene and younger metasedimentary ^^^^1 basement 

Subbasement (basaltic intrusions) 

Inferred intrusions or fractures 

underlying geothermal areas 

Brawley seismic zone 

- - - Fault- Arrows indicate direction of 

relative movement 

Principal stress direction 

SANANDREAS 

FAULT 

Salton Sea 

SAN JACINTO 
T he crust of much of California was formed at an Andean-type 

continental margin during the Mesozoic and early Cenozoic, 

and was modified by large strike-slip offsets along the Sun Andreas 

fault system during the late Cenozoic. Decoupling within the crust, 

as implied by present upper-crustal tectonic wedging in central 

California, and decoupling between the crust and mantle, as 

implied by "subduction" of lithaspheric mantle in southern California, indicates that the Sun Andreas fault system must change 

with depth in its location and (or) style of deformation. 

8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM 

SEISMIC-REFRACTION AND OTHER DATA 

By GARY S. FUIS and WALTER D. MOONEY 

CONTENTS 

INTRODUCTION 

Studies of the crustal and upper-mantle structure of 

California along the San Andreas fault system have been 

underway for more than half a century, beginning 

with the early studies by Byerly and Wilson (1935) and 

Byerly (1946) in northern California and by Gutenberg 

(1943) in southern California. Crustal profiling along and 

4 FIGURE 8.1.-Schematic block diagram of Imperial Valley region of the 

Salton Trough, with unmetamorphosed sedimentary rocks removed 

and seismic basement cut away along line approximately parallel to 

the Brawley seismic zone (BZ; see fig. 8.71. Seismic basement consists 

ofrocks with P-wavevelociti~ of 5.5 to 6.5 Ma. In thisreinon. ~"~~ there .~ ~- are two types of seismic basement; one type, on flanks of the Salton 

Trough, consists of pre-late Miocene igneous and metamorphic rocks; 

other tyw. in central part of trou~h, consists of late Miocene and 

younger metasedimentary rocks (similar in age and provenance to 

sedimentary rocks stripped off in this diagram). Pacific and North 

American plates are separating across the Brawley seismic zone, 

an inferred onshore spreading center of the East Pacific Rise. North 

and southof the Brawley seismic zone, these two platenareseparated 

from each other bv transform faults. the San Andreas and Imnerial 

faults, respectively. As the plates pullapart, subsidenceoccurswjthin 

the Brawlcy seiftmicmrw. sediment isdeposited tu fill rift from ahove, 

and mafic intrusions (basalt, diabase, and gabbrol enter rift from 

below, metamorphosing sedimentary rocks below a certain depth 

(generally appmx 5 kmincentral part ofrift). This processis repeated 

until central nart of rift. consiata entirelv --- of vouw cmst. Geomnhic -. 

locations projected downward onto seismic basement for reference. 

207 

THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA- US. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
208 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

near the San Andreas fault was first accomplished in the 

early 1960's by Eaton (1963), Healy (1963), and Roller 

and Healy (1963). Research accelerated after the 1966 

M=6.O Parkfield, Calif., earthquake to include both 

detailed crustal profiling and installation of dense seismic 

networks for the study of earthquakes (see chap. 5; Eaton 

and others, 1970). Since 1970, a wide variety of seismologic methods have been used to investigate crustal and 

upper-mantle structure in the vicinity of the San Andreas 

fault system. In this chapter, we summarize the main 

features of this structure and relate the structure to 

broad-scale tectonic processes. 

Seismologic studies of crustal and upper-mantle structure in California make use of three primary data sources: 

(1) traveltimes of local earthquakes as measured by 

permanent and temporary seismic arrays, (2) seismicrefraction and reflection profiles, and (3) teleseismic 

delay times measured by seismic arrays. Traveltimes of 

local earthauakes. in addition to containing the information needed to locate earthquakes, contain a wealth of 

information regarding the seismic-velocity structure of 

the crust and upper mantle. Velocity structure can be 

determined from these traveltimes by iteratively adjusting an initial velocity model and associated hypocentral 

parameters, using inversion methods (for example, Crosson, 1976; Eberhart-Phillips and Oppenheimer, 1984). 

The resolution of velocity structure from local earthquake data is a function of the interstation spacing of the 

network and the abundance and distribution of seismicity. Seismic refraction and reflection profiles together form 

a complementary set of seismic measurements. Seismicrefraction profiles provide the highest resolution of 

seismic P-wave velocities in the crust and upper mantle. 

The seismic-refraction method, however, generally does 

not provide the sharpest picture of lithologic interfaces, 

from which geologic structure is inferred; such a picture 

is better by seismic-reflection profiling. 

Teleseismic delav-time studies offer the most effective 

means of determining the structure of the subcrustal 

lithosphere. The method is based on interpreting relative 

arrival times of compressional waves throughout a seismic array in terms of velocity variations at depth beneath 

the array. The Earth structurein the volume beneath the 

array generally is described by a series of blocks, and 

velocity deviations are derived for each block from the 

observed delay times (Aid and others, 1977; Thurber and 

Aki, 1987). The California seismic array is ideally suited 

for such investigations because of its large areal extent 

and the length of time it has been in operation (see chap. 

5, -,. 

The primary product of the seismologic methods described above is a model of the seismic P-wave-velocity 

distribution in the crust and upper mantle. However, the 

interpretation of seismic P-wave velocities in terms of 

rock type is highly nonunique because laboratory velocity 

data indicate that numerous rock types can have similar 

velocities (for example, Birch, 1960). This interpretation 

is further complicated by the fact that, in rocks at 

pressures of less than 2 kbars (depths above 8 km), 

seismic velocities are strongly affected by the presence of 

cracks (on all scales) and porosity. In addition, rock 

velocities are affected by temperature and the presence 

of water. Thus, the interpretation of P-wave velocities in 

terms of rock types must involve other data sets, 

including laboratory velocity measurements on rocks at 

different pressures, temperatures, and water saturations, surface geologic data, well data, and other geophysical data, including gravity and magnetic data. 

Fortunately, abundant laboratory velocity data (for example, Stewart and Peselnick, 1977; Lii and Wang, 

1980), geologic data (see chaps. 1, 3), and geophysical 

data (see chap. 9) are available for California, making the 

lithosphere of this region one of the best studied in the 

world. 

In this chapter, we summarize the lithospheric structure and tectonics along the San Andreas fault system of 

California (fig. 8.2) with maps of crustal thickness and 

upper-mantle seismic-velocity anomalies, and with crustal cross sections for central and southern California. 

Structure changes more rapidly parallel to the San 

Andreas fault in southern California than in central 

California, and so we supplement the cross section for 

southern California with a map showing crustal-block 

motions and a diagram illustrating the different motion of 

the lithospheric mantle below. Seismic and other data 

currently are still not dense enough to construct a cross 

section along the San Andreas fault system itself. 

Construction of the crustal cross section for central 

California led us to a new interpretation of upper-crustal 

tectonic wedging, the mechanism whereby the Franciscan assemblage was emplaced in the Coast Ranges 

during the late Mesozoic(?) and Cenozoic. This interpretation extends that of Wentworth and others (1984) to 

include a two-part history whereby the observed structures atop the wedge, which include both extensional and 

compressional faults, were created. We further speculate 

that similar tectonic wedging occurred in southern California from the Mojave Desert to the Chocolate Mountains to emplace thi~and schist and the Pelona-Orocopia 

schist of Haxel and Dillon (1978) into rocks east of the San 

Andreas fault, 

CRUSTAL-THICKNESS MAP OF CALIFORNIA 

A contour map of crustal thicknesses in California (fig. 

8.3A) provides an overview of the geophysical setting of 

the San Andreas fault system. The seismic and gravity 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIGREFRACTION AND OTHER DATA 209 

FIGURE 8.2. -California, showing place names, geologic provinces, selected geologic units, and locations of crustal transects shown in figures 8.4 

and 8.6. The San Andrea. fault extends from the Salton Trough to triple junction at Cape Mendocino. CPF, Cem Meto fault; IF, Imperial 

fault; SCI. Sania Catslina Island. Fault with crosslixhg la trench, offshore northern California. 

data used in compiling this map (fig. 8.3B) were discussed 

by Mooney and Weaver (1989). 

Crustal thickness along the San Andreas fault increases from 16-24 km in northern California to 28-32 km in 

southern California. Thus, the crust along the San 

Andreas fault system is everywhere thinner than the 

36-km average for the conterminous United States 

(Braile and others, 1989), and in northern California it is 
210 THE SAM ANDREAS FAULT SYSTEM, CALIFORNIA 

substantially thinner than this average. To a first-order 

approximation, crustal thickness resembles the topography (see Jachens and Griscom, 1983, fig. 13). 

Cape Mendocino in northern California marks the 

change from the strike-slip regime of the San Andreas 

fault to the subduction regime of the Cascade Range. 

North of Cape Mendocino, the crust thickens eastward 

from about 16 km at the coast to about 38 km in the 

southern Cascade Range (fig. 8.3A). Near the coast, this 

thickness includes both the North American plate and the 

subducting Gorda plate. Estimates of crustal thickness in 

the northern Coast Ranges at Cape Mendocino lack 

FIGURE 8.3.-Crustal thickness (A) for California and adjacent (within enclosed 40-km contour) correspond0 to magmatic arc of 

regions, modified from Mooney and Weaver (1989), with data same age (Sierra Nevada). Southweat of the San Andreas fault in 

soma (B). Contour interval, 2 km. Northeast of the San central California, thin Andean-marginal sequence is repeated but 

Andreas fault in central California, thin crust (within enclosed shortened; crust ie relatively thin there. In southern California, 

28-!an contour) corresponds to Mesozoidearly Cenozoic forearc crustal thickne~~ ia relatively uniform (about 30 km), despite 

basin (Great Valley; nee fig. 8.2 for place names), and thick crust considerable tectonic activity throughout most of geologic time, 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIC-REFRACTION AND OTHER DATA 211 

including present mibduction of lithmpheric mantle (see below). Cascade Range continental arc. Dot pattern, area of contours on 

Estimated error in figure ASA is 10 percent, or I to I/; contour Gorda-plateMoho. IF, Imperialfault. Datasourcesin8.35include 

seismic refraction or reflection control, but detailed 

gravity models, heat-flow observations, and teleseismic 

data indicate an abrupt decrease in both crustal and 

lithospheric thickness southward of the landward projecintervals. Heavy lines, faults-dashed where the Mendodno seismi&draction profiles (dotted lines), earthquake networks 

Fracture Zone extends onshore (and beneath North American (wavy outlines), and gravity (dashed outlines). W' associated 

plate); arrows indicate direction of relative movement; malined with seismic-refraction profile in the Salton Trough indicates that 

along trench, offshore northern California. Triangles, volcanoes of only wide-angle reflections are available to constrain Moho depth. 

tion of the Mendorino Fracture Zone (see chaps. 9, 10; 

Zandt and Furlong, 1982; Jachens and Griscom, 1983). 

In central California, the crust thickens eastward from 

about 25 km near the coast to as much as 55 km in the 
212 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

Sierra Nevada, but this general landward thickening is 

interrupted by thin crust (25 km) beneath the Great 

Valley (fig. 8.314; compare Oppenheimer and Eaton, 

1984). The crust of central California represents a Mesozoic and early Cenozoic Andean-type continental margin 

(see chap. 3; Hamilton, 1969) that has been modified by 

late Cenozoic strike-slip faulting along the San Andreas 

fault system and by uplift of the Sierra Nevada. Andean 

features include a subduction complex (eastern Coast 

Ranges), a forearc basin (Great Valley), and a magmatic 

arc (Sierra Nevada). Cenozoic strike-slip faulting along 

the San Andreas fault system has moved a shortened 

Andean-marginal sequence outboard of this sequence. 

Southwest of the San Andreas fault, the batholithic 

Salinian block (western Coast Ranges) is juxtaposed, 

across other rightloblique-slip faults of the San Andreas 

fault system, against an inactive accretionary prism, or 

subduction complex (western Coast Ranges and offshore 

California). 

In southern California, the crust thickens eastward 

from about 20 km at the western margin of the California 

Continental Borderland to about 32 km in the eastern 

Transverse Ranges (fig. 8.3A). Over most of onshore 

southern California, crustal thickness is 30Â± km. Considering the complex tectonic history of this region, 

including the present subduction of lithospheric mantle 

(see below), this uniformity in crustal thickness is remarkable. 

CENTRAL CALIFORNIA 

Crustal structure in central California is grossly two 

dimensional, as can be readily inferred from the crustalthickness map (fig. 8.3A). There are five blocks or 

provinces with subparallel fault boundaries: an accretionary prism, which is partly off shore; the Salinian block, 

which underlies the western Coast Ranges, including the 

Santa Cruz Mountains and Gabilan Range; a complex 

block between the San Andreas and Calaveras faults, 

underlying the Santa Clara Valley; the Diablo block, 

beneath the Diablo Range; and the Great ValleylSierran 

block (fig. 8.2). To illustrate the crustal structure of 

central California, we have modified and reinterpreted 

the part of Centennial Continent-Ocean Transect C2 

(Saleeby, 1986) that extends from offshore California at 

Monterey Bay to the Sierran foothills near Modesto (figs. 

8.2, 8.4). Seismic control, which is exceptionally good 

along this transect, has been augmented since Saleeby's 

(1986) study primarily by analysis of seismic-refraction 

profiles in the Great Valley (fig. 8.4A). The reader is 

referred to Hill (1978) for an earlier treatment of deep 

structure along approximately this same transect. 

OFFSHORE REGION 

The offshore region of transect C2 is underlain by an 

inactive, early Tertiary accretionary prism overlapped 

by uppermost Oligocene to Holocene sedimentary rocks 

(see Saleeby, 1986). The San Simeon terrane, consisting 

of Late Cretaceous Franciscan rocks (disrupted marine 

sedimentary rocks; see chap. 3), is imbricated in this 

prism along with poorly known, lower Tertiary sediientary rocks. The prism is underlain by oceanic crust with 

an inferred age of about 26 to 20 Ma (Atwater and 

Menard, 1970; Atwater, 1989). The accretionary prism is 

juxtaposed against granitic rocks of the Saliian block 

across the (inactive) Nacimiento fault, which is overlapped by upper Tertiary sedimentary rocks. This fault, 

in turn, is offset by the (active) rightloblique-slip San 

Gregorio-Hosgri fault. The Moho is 10 km below sea level 

near the west end of the transect (Shor and others, 1971) 

and deepens to 24- to 26-km depth beneath the Gabilan 

Range and Santa Cruz Mountains (Walter and Mooney, 

1982). 

We follow D.S. McCulloch (in Saleeby, 1986) in showing steep northeastward dips on both the Nacimiento and 

San Gregorio-Hosgri faults (fig. 8.4A) that are based on 

marine reflection data. Focal mechanisms for earthquakes on the San Gregorio-Hosgri fault at this latitude 

indicate nearly pure strike slip on vertical planes; however, farther south, they indicate chiefly reverse faulting 

on northeast- or southwest-dipping planes (see chap. 5). 

SALINIAN BLOCK 

The area between the Nacimiento and San Andreas 

faults is underlain by a batholithic terrane that has been 

transported northwestward by the San Andreas (and 

other?) fault(s) by amounts estimated to range from 550 

km (see ROSS, 1978) to 2,500 km (Champion and others, 

1984). Plutonic rocks include tonalite, granodiorite, and 

quartz monzonite of mostly Late Cretaceous age (Ross, 

1978; Mattinson, 1982). Metamorphic pendants and 

screens include mostly quartz-rich clastic rocks of amphibolite fades. Ross and McCulloch (1979) postulated 

that these upper-crustal plutonic and metamorphic rocks 

are not rooted to the lower crust but are in fault contact 

with a buried terrane, possibly consisting of Franciscan 

rocks. 

The velocity structure derived by Walter and Mooney 

(1982) from Stewart's (1968) seismic-refraction measurements in the Gabilan Range and Santa Cruz Mountains 

can be subdivided into four separate crustal layers with 

velocities of 2.14.6 Ws (layer I), 5.8-6.6 Ws (layer 2), 
B. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIGREFRACTION AND OTHER DATA 

EXPLANATION FOR FIGURES 8.4A AND 8.6A 

3.w4.B P-wave velocity in kilometers pa- second at topibase 

of Iayer~Velocities in parentheses are projected .5 1 

Continental and continental rift 

Nonmarine daitic rocksÃ‘Circles conglomcnac: - curie', and dots, conglomerate and sandstone; 

darker color, metamorphosed 

P-wave velocity boundary-Heavy tick where profile 

perpendicular to page; light line where profile ii 

plane of page: shon dashed where change in gradient 

only. longdashed in tig~rcH4A for model ot Dean 

shore central ~alifornia (fig. 8.4~) where they are determined by reflection profiling 

7 Z Reflector from reflection profiling-Queried where 

uncertain 

-0-0- Density and magnetic boundary 

7 Inferred nijilin in which milkiwld and Moho 

discontinuities may exist-Queried where 

uncertain 

LVZ. Fusible loit-velocity am 

- Fault 

EXPLANATION FOR FIGURES 8.4B AND 8.6B 

LITHOLOGIC SYMBOLS FOR TECTONIC ENVIRONMENT 

Oceanic 

Bwdn floor and trench turhiditos 

m Pelagic sedlmentir; rocks 

Ophiolite, or top of ophiolhe whereundivided mafic crust is shown below 

- 

Mqmlic am 

Foreare rocks 

Intra-arc rocks 

vole 

m Rocksof mixed and (or1 intermediate composition 

Grunite and qnirtz monionile 

3 I Marine clastic -Strike-and-dip symbol, 

carbonatetearing 

vz 4 Rocks of mixed felsic M mafic composition 

( Middle crust-Inferred to be gneiss andschisi 

developedat 10- to 20-km depth at greenschist 10 

Mafic crust* undivided-Denser pattern: 

higher velocity 

MmUl~Closely ipiicedtkks. top of lithoipheric 

mantle; widely spaced ticks, top of aslhenospheric 

mantle or lithospheric mantle with partial melt 

OTHER SYMBOLS 

Melange-Symbols enclosed by zigzag lines 

indicate block type 

- Lithologic contact 

- Ã‘Ã‘Ã A Normal or strike-slip faultÃ‘Sens of motion shown: 

arrows indicate direction of relative movement: 

- A. away from observer; T. toward observer 

Mesozoic thrust fault-Sawteeth on upper plate 

Mesuuk or Cvmic thrust faiiMÃ‘Teet MI upper 

plate; mw indicates inferred motion 01 upper plate; 

dotted where interred to hate formerly existed 

- 3 Ce-k thrust fault-Teeth on upper plate; arrow 

indicates inferred mtbn of upper plate 

- Mesozoic or Cenozoic detachment faulI-Ticks 

1 upper plate 

Cenozok detachment faultÃ‘Tick on upper plate 

In lwrt 8.6, Chocolate Mountains: Mesozok 

or Cenozoic thrust or detachment fault 

("Chocolate Mounuins thrust fault"; see text) 

NOTE: Liiologic contacts and faultsÃ‘Lon dashed where not 

conslrained by figures 8.4A aid 8.6A or by earthquakes 

of chanter 5: short dashed where cradational (litholocic 

rntktonly) or projected above surface; queried where 

existence uncertain 

? Tup of brittle-ductile tnrsitii zone inferred 

fro- earthquakes (sre 1ext)Ã‘Quene where 

uncertain 

Expltiution for figurai 8.4 and 8.6. 
THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

SAN GI 

HOSGffl F, 

mIO- SAM ANDREAS 

J SYSTEM FAULT 

OFFSHORE REGION 

Areamdotted 1-mYT-w 

mb- accretlonarywedge: #refection profileDS.MoCdhxtlA LateTertiaty 

SÃ‘Ã‘b lÃ‘l -am 

convenedtod* 

lochs 

NACIMIENTO 

FAULT 

I 

Fmci-as-: 

San am terrene 

(Late 0-1 

SANTA CRUZ MOUNTAINS 

Ma" Las m-bs 

LT-b c* mmmawo mde and Tam sedimefitatv rocks 

SANTACLARA 

VALLEY 

Tertiarysafimetteyrocks SARGENT and & of 

FAULT co~~arn. 

Wwii md 

sequence 

FIGURE 8.4.-Crustfll structure of central California. A, Surface geology, 

depth-converted seismic-reflection data, and models of seismic-refraction, 

gravity, and magnetic data along western part of Centennial ContmentOceanTd C2 (see Saleeby, 1966). B, Reinterpretation of Trans& C2. 

Major features in figure 8.46 include, from west to east, (1) offehore, 

inactive. ~,~~ edv . Tertiarv accretinnarvwedoe: (21 batholithic Salinian block of ~~ 

.. . the&& Cruz Mountains, positioned between the active ~blique-alip San 

Gregorio-Hoagri and San And- faults; (3) Franciscan terranes of the 

Santa Clara Valley and Diablo Range, interpreted to compose a tectonic 

wedge; (4) Mesozoic and Cenozoic sedimentary rocks ofthe Great Valley; 

and (FA .-. rocks of the Si- footbilk.. ineludirur Jurassic and older ~ ~~~-~ ~ 

volcanic. plutonic, and Ated sedimentary &-accumulated or emplaced in an island-arc setting and Cretaceous plutonic rocks. Tectonic 

wedge in feature 3 is interpreted to have moved during the late 

~&d?) and Cenozoic, possibly in several episodes, largely along 

contact between Mesozoic crystalline rocks and overlying Meaozoic 

sedimentary rocks. Inthe eastern Great Valley, these sedimentary rocks 

are still rooted to (or depoflitionally overlie) this basement. Movement of 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIGREFRACTION AND OTHER DATA 216 

SAN ANDREAS CUAVERAS 

FAULT FAULT SYSTEM TESLA-1 

SANTA CLARA VALLEY n- GREAT VALLEY SIERRAN 

FOOTHILLS 

-me &: 

Volcanic, plutofik, and 

sdmmw, rocks Cenozoic sedmentary and volcanic rocks (Jurassic aid older) 

Great Valey sequence 

(Cretaceous and Late Jurassic) 

I 

wedgeduringpresent SanAndreastranfifonn-faulting&maybealong layers 1 through 4 is shown (fig. 8.4-4; see fig. 8.6-4). along with 

one or more thrust faults that merge with postulated decoUement in alternative model in which layer 4 is subdivided into layers 48 and 4b. 

brittle-ductile transitionzonein the crust, 'nus reinterpmtationdiffersfrom First model giva rise to interpretation a, and alternative model to 

Saleeby's (1986) in eliminating interred east-dipping sunduction zone or interpretation b Wg. 8.4B). a-j, reflectors in the eastern Diablo Range, 

thrust fault beneath the western Great Valley. Off shore, interpretation of Great Valley, and Sienan foothills; 1-4, seismic-velocity layers in the 

mmigmtdreflectionsection by D.S. McCullodi(inSaleeby, 1986)hanbeen Great Valley (fig. 8.4A) ihcuaed in text. See figures 8.2 and 8.3 far 

converted to depth section, wing assumed velocities for each interred location of Meet C2. No vertical exaggeration. 

geologic unit. In the Santa Cruz Mountains, velocity model consisting of 
216 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

6.0-6.15 km/s (layer 31, and 6.35-6.66 km/s (layer 4). 

Layer 4, middle and lower crust, can alternatively be 

modeled as two layers of velocities 6.3 km/s (layer 4a) and 

6.6-6.8 km/s (layer 4b). These layer velocities can be 

correlated to rock type using surface geologic data and 

laboratory velocity data. Layer 1 corresponds to outcrops 

of Cenozoic sedimentary rocks along the transect. Basement outcrops along or near the transect include abundant quartz monzonite (Ross, 1972). Lin and Wang (1980) 

studied the velocity behavior of a sample of quartz 

monzonite from this region as a function of pressure and 

temperature, and constructed a velocitydepth curve for 

this rock appropriate for the Coast Ranges. On their 

curve (fig. 8.5A). the rock is slightly faster than layers 2 

and 3 and slower than layer 4. Walter and Mooney (1982) 

interpreted layers 2 and 3 as granitic rocks similar to this 

0 . A 

EXPLANATION 

0 Quartz mmzonitc 

Alternate 

model 

quartz monzonite. The somewhat lower velocity of these 

two layers in comparison with the laboratory sample may 

be interpreted to result from (1) megascopic fi-actures in 

the Earth, not present in the laboratory sample; (2) a 

slightly lower content of mafic minerals (which have high 

seismic velocity) in the granitic rocks beneath the 

transect in comparison with the laboratory sample; or (3) 

both. Layer 4 is intermediate in velocity between the 

quartz monzonite sample and gabbro samples (hornblende gabbro and olivine gabbro) from the Coast Ranges. Walter and Mooney (1982) interpreted this layer to 

correspond to gneiss of intermediate composition, on the 

basis of a comparison of layer 4 with other laboratory 

data. In an alternative model, however, where middle 

and lower crust are separated as layers 4a and 4b, layer 

4h may be reasonably interpreted as gabbro (fie. 8.5A). 

~l&tive interpretations of these several crustal 

layers are also possible, given the fact that different rock 

types may have similar velocities. Stewart and Peselnick 

(1977) and Lim and Wang (1980) studied the velocity 

behavior of Franciscan rocks, also common in the Coast 

Ranges (Jennings and Strand, 1968). Two lithologic 

components of the Franciscan assemblage, unmetamorphosed and metamorphosed graywacke, produce velocitydepth curves (fig. 8.55) that bracket those for most 

other components of the Franciscan assemblage (including basalt). On the basis of velocity data alone, layers 2 

and 3 might be interpreted as Franciscan rocks, but 

surface geologic data lead us to reject this interpretation. 

On the basis of velocity data alone, however, layer 4 is 

most likely not Franciscan rocks. Thus, if the middle and 

- lower crust of the Salinian block represents a different 

s 

0 terrane from the upper crust, as postulated by Ross and 

z R 1 McCuUoch (1979), that terrane is most likely not Fran- - EXPLANATION 

0 Unmelamorphosed 

graywacke 

0 Mclamo~~h~i~d 

graywacke 

ciscan assemblage. 

In our cross section (fie. 8.45), we show alternative 

interpretations of layer 4, given the alternative velocity 

models discussed above. In one interpretation (a, fig. 

8.4B), layer 4 is entirely gneiss of intermediate composition. In a second interpretation (b, fig. 8.4B). layer 4a 

is intermediate gneiss, and layer 4b is gabbro. In 

interpretation a, no buried terranes are present; in 

interpretation h, the lower-crustal gabhro may be a 

\ buried terrane (oceanic crust) or magmatically undqlated gabhro. 

I 

4 5 6 7 8 SANTA CLARA VALLEY-SAN ANDREAS 

VELOCITY. IN KILOMETERS PER SECOND 

FIGURE 8.5.-Velocitydepth curves. A, Santa Cruz Mountains and 

Catalan Ranee (Salinian block). B, Diablo Range. Heavy curves from 

seismic results (Walter and Mooney, 1962); light curves from laboratory velocity mwements and heat-flow modeling (two different. 

geotherms assumed below about 10-km depth; Lin and Wang, 1980). 

See text for discuBsion of layers shown in figure 8.5A. 

TO CALAVERAS FAULTS 

In the Santa Clara Valley, between the San Andreas 

and Calaveras faults, Franciscan assemblage (Permanente terrane; Blake and others, 1982) is overlain by 

outliers of Late Jurassic Coast Range ophiolii and 

Upper Cretaceous Great Valley sequence (McLaughlin 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIC-REFRACTION AND OTHER DATA 217 

and others, 1988a). The Franciscan assemblage includes 

melange, volcanogenic sandstone, pillow basalt, serpentine, chert, and limestone. The Franciscan sedimentary 

rocks were deposited in equatorial waters and presumably transported thousands(?) of kilometers northward 

before accretion to the North American Continent (Blake 

and others, 1982). 

Along transect C2, the San Andreas fault juxtaposes a 

thick section of Tertiary marine sedimentary rocks on the 

southwest against slivers of Coast Range ophiolite. Great 

Valley sequence, and other Tertiary marine rocks on the 

northeast that have been imbricated along the southwestdipping Sargent fault and related thrust faults 

(McLaughlin and others, 1988a). Presumably, the granitic rocks of the Salinian block and Franciscan assemblage 

are juxtaposed at depth. In contrast, similar rocks are 

juxtaposed on either side of the Calaveras fault, including 

Coast Range opbiolite, Great Valley sequence, and, at 

depth, presumably the Franciscan assemblage. 

Aftershocks of the M=7.1 Loma Prieta earthquake of 

1989 indicate a steep southwestward dip (approx 7W) on 

the San AndreasISargent fault zone, and the main shock 

produced subequal components of strike- and reverse-slip 

motion (Plafker and Galloway, 1989). Relatively low 

elevations in this region, however, indicated that the 

motion along this fault zone in the past has been chiefly 

strike slip, and seismicity before the Loma F'rieta earthquake (Olsen and Lindh, 1985; Olsen, 1986) indicates a 

complex fault zone that may include both vertical and 

southwest-dipping fault strands (see fig. 8.4B). 

Although a steep (80'45") eastward dip on the Calaveras fault is indicated by earthquakes (see cross sec. 

D-D', fig. 5.7B; Reasenberg and Ellsworth, 1982; Oppenheimer and others, 19881, such an attitude is not 

resolvably different from a vertical dip (shown in fig. 

SAB), given errors in earthquake locations. 

A seismic-refraction profile across the Santa Cruz 

Mountains and Santa Clara Valley reveals a heterogeneous upper crust (Mooney and Colburn, 1985). Layer 

offsets and velocity changes are visible in the model for 

this profile at the Zayante-Vergeles, Sargent, and Calaveras faults but, surprisingly, not at the San Andreas 

fault. An additional discontinuity is visible at an inferred 

buried fault in the central Santa Clara Valley (fig. 8.4A). 

Vertical zones of low velocity, 1 to 2 km wide, extending 

to a depth of as much as 3 to 5 km, are visible at a few of 

these faults (Mayer-Rosa, 1973; BlUmling and others, 

1985; Mooney and Colburn, 1985). The surficial layer 

(2.1-4.5 kmh) corresponds to different rocks in different 

places (fig. 8AB). The "basement" layer has a velocity 

(5.4-6.0 kmls) appropriate for either granitic or Franciscan rocks at shallow crustal levels (Mooney and 

Colburn, 1985; see discussion above and fig. 8.5); presumably, it represents Franciscan rocks except west of 

the San Andreas fault. The higher-velocity basement (6.0 

km/a) in the eastern Santa Clara Valley may represent 

metamorphosed Franciscan rocks. A strong reflector is 

visible at 8- to 9-krn depth beneath the Santa Clara 

Valley, but the seismic velocity below it is unknown. By 

analogy with the strong midcrustal reflector in the Diablo 

Range (see below), we infer this reflector to be the top of 

accreted island-arc and (or) oceanic crust. 

Moho depth beneath the Santa Clara Valley is not 

known accurately enough to resolve whether the Moho 

steps downward to the east at the San Andreas and 

Calaveras faults or dips smoothly eastward between 

control points in the Santa Cruz MountainsIGabilan 

Range (24- to 26-km depth) and the Diablo Range (29- to 

30-km depth). McEvilly and Clymer (1975) conducted a 

seismic-reflection survey across the San Andreas fault 

south of its junction with the Calaveras and found a 

crustal thickness of 24 km with no change in thickness 

across the fault. Peake and Healy (1977), however, 

indicated a change in crustal thickness at the fault in this 

area. 

DIABLO RANGE 

The Diablo Range, the east-central Coast Ranges 

between the Calaveras fault and the Great Valley, is 

underlain chiefly by Franciscan assemblage. These rocks 

constitute at least three thrust sheets or nappes that are 

folded into an antifonn (fig. 8.4; Blake, 1981; Saleehy, 

1986). The youngest thrust sheet, the Burnt Hills terrane 

(Blake and others, 1982; Saleeby, 1986), consists of 

mid-Cretaceous blueschist-facies graywacke, arkose, 

conglomerate, argillite, and chert, approximately equivalent in age and provenance to mid-Cretaceous forearc 

sedimentary rocks of the Great Valley sequence. The 

Burnt Hills terrane is exposed in the core of the antifonn. 

The oldest thrust sheets are the Upper Jurassic (informal) Garzas tectonic melange (Cowan, 1974) and the 

Yolla Bolly terrane (Blake and others, 1982). The Garzas 

tectonic melange consists of mafic bluescbist-ampbibolite, greenstone, serpentinized peridotite, and metagraywacke; it contains fragments of Upper Jurassic rocks 

(Coleman and Lanphere, 1971; Suppe and Armstrong, 

1972) similar to those accreted in the Sierran foothills 

during the Nevadan orogeny (see below). The Yolla Bolly 

terrane lithologically resembles the Burnt Hills terrane, 

although there are some important differences. Both the 

Garzas tectonic melange and Yolla Bolly terrane crop out 

on the flanks of the antiform. The Coast Range ophiolite 

and Great Valley sequence lie above the Franciscan rocks 

on the low-angle Coast Range fault, which is complexly 

offset by the steeply dipping Cenozoic Tesla-Ortigalita 

fault on the northeast flank of the Diablo Range (fig. 8.4). 

(We follow Jayko and others, 1987, in referring to the 
218 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

"Coast Range thrust fault" as simply the "Coast Range 

fault1'-see below.) 

The velocity structure of the Diablo Range derived by 

Walter and Mooney (1982) from seismic-refraction data 

collected by Stewart (1968) includes, beneath a 3.5- to 

5.3-km/s near-surface layer, a 5.7- to 5.9-km/s layer 

beginning at 3-km depth, a 6.7- to 7.1-km/s layer beginning at 15-km depth, and the Moho at 29-km depth (fig. 

8.4A). Importantly, a strong reflection is observed from 

the layer boundary at 15 km, indicating a strong velocity 

contrast between upper and lower crust. Blumling and 

Prodehl (1983) reanalyzed the same data and derived a 

similar velocity structure, except that they interpreted 

more phases in the data and added a lower-crustal 

low-velocity layer (5.3? km/s), with its base at about 

26-tan depth. 

Seismic-reflection data have been collected in the 

eastern Diablo Range (Zoback and Wentworth, 1986) and 

compiled with other seismic data (Wentworth and others, 

1987). These reflection data include a band of strong 

reflectors in the upper crust that dips shallowly east 

(reflectors a, fig. 8.4A), a weak reflector in the middle 

crust that dips shallowly west (reflector b), and a weak 

reflector at about 30-km depth (reflector c). The shallowly west-dipping reflector b, appears to link the top of the 

Great Valley basement with the top of the 6.7- to 

7.1-kmls layer (Wentworth and Zoback, 1989). 

Between 3- and 15-km depth, seismic velocities in the 

Diablo Range are well bracketed by velocity-depth 

curves predicted for end-member rocks ofthe Franciscan 

assemblage (fig. 8.5B; Steward and Peselnick, 1977; Lii 

and Wang, 1980). Within this depth range, the observed 

velocities also are slightly lower than those predicted for 

most granitic rocks (see fig. 8.5A). Between 15- and 

20-tan depth, the observed velocities agree well with 

those predicted for gabbro (fig. 8.5B; Lin and Wang, 

1980) or, possibly, high-grade metamorphic rocks (Birch, 

1960; Christensen and Fountain, 1975). 

The 6.7- to 7.1-km/s layer in the Diablo Range may 

represent the middle and lower crust of an island arc or 

several imbricated island arcs. If so, this layer might 

include mixed intermediate and mafic plutonic rocks, 

including compositions from granodiorite to gabbro, as 

well as metamorphic rocks (see description of the Coast 

Range ophiolite by Evarts, 1977). Its relatively high 

velocity indicates that rocks of mafic composition must 

dominate or that the rocks are of amphiholite to granulite 

facies. This "island arc" interpretation is consistent with 

linking this layer to rocks beneath the Great Valley and 

thence to rocks of the Sierran foothills, which represent 

the middle and upper crust of island arc(s) (Saleeby, 

1986). The 6.7- to 7.1-km/s layer, however, may also 

represent middle and lower oceanic crust, or diabase and 

gabbro, similar to the lowest layer of oceanic crust at the 

west end of transect C2 (fig. 8.4A). The 6.7- to 7.1-km/s 

layer is too thick, however, to represent a single layer of 

oceanic crust. It consists of either several slices of 

tectonically underplated oceanic crust or of oceanic crust 

that has been augmented by mafic intrusions after 

underplating. If the low-velocity zone of Blumling and 

Prodehl(1983) is present, the 6.7- to 7.1-km/s layer may 

include oceanic sedimentary rocks tectonically underplated along with the oceanic crust. 

We show a fault contact between the Franciscan 

assemblage and the 6.7- to 7.1- km/s layer in the Diablo 

Range to reflect the eastward transport of a wedge of 

Franciscan rocks (fig. 8.45). similar to that discussed by 

Wentworth and others (1984). This interpretation departs from that of Saleeby (1986), who linked the 

shallowly eastdipping reflectors in the eastern Diablo 

Range (a, fig. 8.4A) with a hypothetical subduction zone 

or thrust fault beneath the Great Valley and Sierran 

foothills (see section below entitled "Discussion-Tectonic Wedging"). 

GREAT VALI-FY AND SIERRAN FOOTHILLS 

Rocks of the Great Valley are known from exposures in 

an upturned section on the east side of the Diablo Range 

and from wells. The upturned section rests structurally 

above the Franciscan assemblage on the low-angle Coast 

Range fault, although in many places this relation is 

obscured by younger high-angle faults. This upturned 

section includes, from oldest to youngest, Middle and 

Late Jurassic Coast Range ophiolite and a related tuffaworn unit" Unner Jurassic and Cretaceous Great Valley - . .. sequence, which is chiefly forearc flysch; lower Cenozoic 

postarc marine and terrestrial sedimentary rocks; and 

upper Cenozoic continental-arc sedimentary rocks (Maddock, 1964; Evarts, 1977; Bartow and others, 1985). 

At the latitude of transact C2, the Coast Range 

ophiolite is interpreted to be a (rifted) island-arc assemblage because it contains abundant silicic volcanic and 

intrusive rocks (Bailey and Blake, 1974; Evarts, 1977; 

Hopson and others, 1981; McLaughlin and others, 1988b). 

Its contact with the overlying sedimentary rocks, though 

faulted in most places, is believed to be fundamentally 

depositional (Bailey and others, 1970); on transect C2 it is 

demonstrably depositional (Evarts, 1977). 

The Great Valley sequence and younger rocks exposed 

in the upturned section in the eastern Diablo Range 

appear to be nearly twice as thick as the section of 

sedimentary rocks penetrated in wells farther east in the 

Great Valley (fig. 8.4). Some of this apparent westward 

thickening results from the stratigraphic addition of older 

rocks to the basal part of the section in the west; some 

may be caused by imbrication along thrust faults (Wentworth and others. 1984). Similar apparent thickening 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIGREPBACTION AND OTHER DATA 219 

west of the synclinal axis of the Great Valley has been 

documented in other localities as well. In the southern 

Great Valley, Wentworth and others (1984) indicated ar 

apparent doubling of thickness west of the axis, and in 

the northern Great Valley, an apparent trebling oi 

thickness. 

Most of the basement rocks that have been penetrated 

by wells in the Great Valley have been identified as 

granitic rocks (Saleeby, 1986). Rocks exposed in the 

Sierran foothills, east of the Great Valley, may be related 

to basement rocks beneath the Great Valley, but they are 

not so dense or magnetic (see below; fig. 8.4A). 

Deep structure along transect C2 in the Great Valley 

has been elaborated in some detail by Colburn and 

Mooney (1986). Holbrook and Mooney (1987), and Dean 

Whitman and others (unpub. data, 1985) from seismicrefraction data, and by Wentworth and others (1987) 

primarily from seismic-reflection data. Seismic velocities 

in the sedimentary section range from 1.6 to 4.1 km/s 

where these velocities can be clearly ascribed to sedimentary rocks, such as near well No. 1 (fig. 8.4A). In the 

eastern Diablo Range, velocities as high as 4.7 km/s may 

also be due to sedimentary rocks (fig. 8.4.4). East of the 

synclinal axis in the Great Valley, reflections within the 

sedimentary section are subparallel to the top of basement, which is marked by the disappearance of reflections (f, fig. 8.4.4). West of the synclinal axis, these 

reflections (d, fig. 8.4A) diverge slightly from the inferred top of basement (e, fig. 8.4A). 

Beneath the sedimentary rocks of the Great Valley are 

several layers of increasing seismic velocity: a 5.6- to 

5.7-km/s layer, 1.5 to 2 km thick (layer I), a 6.0- to 

6.3-km/s layer, 2.5 to 6 km thick (layer 2); a 6.6 to 

6.75-km/slayer, 4 to7 km thick (layer3); and a6.9- to 7.2- 

km/s layer, about 7 km thick (layer 4) (fig. 8.4.4). In 

addition, there is a thin, laterally discontinuous 7.0-km/s 

layer embedded in the top of layer 3. Well No. 1 indicates 

that layer 1 is granitic rocks. Farther west, however, this 

layer may be interpretable either as granitic rocks or as 

Franciscan assemblage, which have similar velocities at 

this depth (fig. 8.5). In the original data of Colburn and 

Mooney (1986) and Holbrook and Mooney (1987), there is 

no perceptible reflection from an interface between 

layers 1 and 2 (as there is, for example, between layer 1 

and the overlying sedimentary rocks), and so these two 

layers may, in fact, grade into one another. Layer 2 could 

then also he granitic rocks, and layers 1 and 2 together 

would constitute a velocitydepth section similar, for 

example, to upper crust of the batholithic Salinian block 

(figs. 8.4.4,8.5). Layers 3 and 4 (6.6-7.2 Ms), which are 

analogous to the lower-crustal layer in the Diablo Range 

(6.7-7.1 kmls), may represent the middle and lower crust 

of accreted island arc@) and (or) oceanic crust. The Moho 

is well documented at about 27-km depth. Deep reflection 

data beneath the Great Valley (Wentworth and others, 

1987) indicate a conspicuous east-dipping band of reflections (g, fig. 8.4A) and less conspicuous subhorizontal and 

west-dipping reflectors. 

Rocks of the Sierran foothills consist of Lower to 

Upper Jurassic mafic to felsic volcanic and plutonic rocks 

and related sedimentary rocks (argillite, chert, and 

flysch) that were accumulated or emplaced in an islandare setting (Clark, 1964; Schweickert and Cowan, 1975; 

Saleeby, 1982; Schweickert and Bogen, 1983). The basement and metamorphic wallrocks for the intrusive rocks 

are tectonically disrupted and polymetamorphosed Paleozoic ophiolitic rocks (approx 300 Ma; Saleeby, 1982). 

The island arc(s) in which the Jurassic rocks of the 

Sierran foothills were formed collapsed against the margin of the North American Continent during the Late 

Jurassic Nevadan orogeny (Jones and others, 1976). How 

this collapse occurred is problematic. Steeply east-dipping faults and upright antiiorms are seen in the Sierran 

foothills, but a study by Moores and Day (1984) of surface 

relations 300 km north of transect C2 indicates obduction 

of the arc(s) on west-dipping thrust faults. These rocks 

were intruded during the Early Cretaceous by mafic to 

intermediate plutons belonging to the western phase of 

Sierra Nevada plutonism (Evernden and Kistler, 1970). 

The deep structure of the Sierran foothills is known 

from the reconnaissance seismic-refraction experiment of 

Spieth and others (1981), the reflection profiling of 

Zoback and Wentworth (1986), and the compilation of 

reflection, refraction, and potential-field results by 

Wentworth and others (1987).~he refraction data can be 

modeled with a 6.2-km/s basement from near the surface 

to about 30-km deoth. a 6.6-kmls lower crust. and a Moho 

at 39-km depth. other models are possible, however, and 

the Moho may be as shallow as 30 km (Spieth and others, 

1981). We have projected the seismic-reflection results 

and gravitylmagnetic boundary of Wentworth and others 

(1987) from 45 to 60 km southward onto transect C2. Two 

conspicuous west-dipping sets of reflections are visible, 

as well as a few subhorizontal reflectors. The gravity1 

magnetic boundary, however, has a moderate eastward 

dip. 

Our projection of the results of Wentworth and others 

(1987) is uncertain not only because of the distances 

involved but also because their profile terminates on the 

east in an area that is anomalous both geologically and 

geophysically. In this area, batholithic rocks (trondhjemite) engulf most accreted rocks of the Sierran foothills 

(Jennings, 1977) and are associated with a gravity low 

[Oliver and others, 1980). Our projection, however, may 

be defensible as follows. (1) The batholithic rocks responsible for the gravity low probably do not extend below 

10-km depth (R.C. Jachens, oral commun., 1988); most of 

the reflectors that we have projected are largely below 
220 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

that depth. (2) The modeled gravitylmagnetic boundary 

is approximately similar in shape throughout the length 

of the Great Valley (Andrew Griscom, oral conunun., 

1988); in our projection, we have attempted to correct for 

the difference in azimuth between transect C2 and the 

profile of Wentworth and others (1987) by assuming a 

strike parallel to the Great Valley. 

Given the geologic and seismic constraints discussed 

above, we have interpreted the cross section through the 

Great Valley and Sierran foothills (fig. 8.4B), using some 

of the ideas of Wentworth and others (1984, 1987) for the 

configuration of an inferred tectonic wedge of Franciscan 

rocks, and some of the ideas of Saleeby (1986) for 

structure within crystalline rocks. The uppermost part of 

our cross section (to approx 2-km depth) on the east flank 

of the Diablo Range (fig. 8.U) was supplied by R.C. 

Evarts (written commun., 1989). Below this area, we 

have added a hypothetical westdipping thrust fault to 

bring the Great Valley sequence beneath the easternmost block of the Coast Range ophiolite and to grossly 

satisfy the velocity constraints of Dean Whitman and 

others (unpub. data, 1985, fig. 8.44). East of the Coast 

Range ophiolite, we postulate thrust faults that largely 

follow bedding planes in the upturned section of the Great 

Valley sequence, similar to those postulated by Weutworth and others (1984) for the northern Great Valley. 

These "backthrust" faults are required for emplacement 

of the wedge and help explain the thickening of the Great 

Valley sequence in the western limb of the syncline (see 

section below entitled "Discussion-Tectonic Wedging"). 

From the easternmost backthrust fault in the Great 

Valley to the San Andreas fault, we have modeled the 

discontinuity between variably reflective rocks of lower 

velocity (Franciscan assemblage, Coast Range ophiolite, 

and Great Valley sequence; 1.7-5.8 km/s) and poorly 

reflective rocks of higher velocity (mafic rocks of the 

Diablo Range and crystalline basement of the Great 

Valley; 5.6-6.8 km/s) as the floor thrust fault of the 

wedge. Wentworth (1987) presented a similar interpretation. 

The details of composition and structure in the crystalline rocks beneath the Great Valley and Sierran 

foothills are speculative. Saleeby (1986) interpreted these 

rocks to consist fundamentally of slabs or nappes 01 

island-arc and oceanic rocks obducted along west-dipping 

Nevadan thrust faults intruded by chiefly Early Cretaceous Sierran granitic plutons. We have adopted this 

basic scheme and added some details, interpreting layers 

1 and 2 in the basement beneath the Great Valley (5.5-6.3 

km/s; see above) as post-Nevadan felsic plutonic rocks, 

although, as noted above, the western part of layer 1 (5.5 

Ms) may be Franciscan assemblage. We interpret the 

east-dipping gravitylmagnetic boundary of Wentworth 

and others (1987) as the average top of mafic crust 

(pre-Nevadan gabbro, diabase, or basalt) in the inferred 

obducted sequence. Alternatively, this boundary may be 

the average top of mafic, magnetic intrusions in the crust 

(post-Nevadan gabbro) or the average base of felsic, 

nonmagnetic intrusions (post-Nevadan granitic rocks). 

At the location where this boundary was actually modeled, it may be the average base of a large trondhjemite 

intrusion. We associate the east-dipping reflections beneath the central Great Valley (g, fig. 8.U) with the 

thin, discontinuous 7.0-km/s layer of Holbrook and Mooney (1987), although the depth correspondence is imperfect, and we interpret this feature as a gabbroic dike. 

Alternatively, these east-dipping reflections may represent an east-dipping fault zone. Following Saleeby (1986), 

we correlate the upper and lower west-dipping bands of 

reflections in the eastern Great Valley and Sierran 

foothills (h, j, fig. 8.4A) with the Bear Mountain and 

Melones fault zones, which may represent Cenozoic 

reactivations of inferred west-dipping Nevadan thrust 

faults. 

DISCUSSION-TECTONIC WEDGING 

GEOLOGIC HISTORY 

Wentworth and others (1984) interpreted the juxtaposition of Franciscan assemblage and a coeval section 

consisting of Coast Range ophiolite and Great Valley 

sequence as having occurred during landward movement 

of the Franciscan assemblage as a tectonic wedge. They 

reinterpreted the "Coast Range thrust fault" of Bailey 

and others (1970), a subduction megathrust between the 

Coast Range ophiolite and the Franciscan assemblage, as 

the roof thrust of the wedge. More recently, the thrust 

nature of the "Coast Range thrust fault" has been 

reevaluated. Jayko and others (1987), testing an hypothesis by Platt (1986). produced abundant evidence that the 

contact between Franciscan assemblage and Coast 

Range ophiolite is a detachment surface along which the 

upper plate was extended duringuplift of the Franciscan 

assemblage. Their evidence is the consistent attenuation, 

as opposed to repetition, of geologic section across this 

discontinuity and associated faults above it. They proposed the term "Coast Range fault" for this discontinuity, which we adopt here. Evidence of attenuation is 

present even on transect C2, in that the two outcrops of 

the Coast Range ophiolite in the eastern Diablo Range 

(fig. 8.4A) represent an abridged section of ophiolite: The 

western outcrop is partially serpentinized ultramatic rock 

of the basal part of an ophiolite, whereas the eastern 

outcrop is the sill complex and volcanic flows of the upper 

part of an ophiolite. These two parts of the ophiolite are 

now juxtaposed across the crooked, steeply dipping 

Tesla-Ortigalita fault. Although this fault now offsets the 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIC-REFRACTION AND OTHER DATA 221 

Coast Range fault, it may represent reactivation of a 

normal fault that originally soled into the Coast Range 

fault (compare Raymond, 1973). 

The extensional nature of the Coast Range fault poses 

several problems for emplacement of the Franciscan 

assemblage as a tectonic wedge. Where is the roof thrust 

fault of the wedge? How did the Franciscan assemblage 

reach its current position with an extended overlying 

section of the Coast Range ophiolite and Great VaUey 

sequence? Was the Franciscan assemblage uplifted from 

beneath the western Great Valley? The apparent continuity between the Great Valley basement and the 6.7- to 

7.1-kmls layer in the Diablo Range indicates a negative 

answer to the last question. 

These problems can be solved if the extensional event 

was separated in time and space from the compressional 

event, or tectonic wedging. Jayko and others (1987) 

reviewed the published evidence regarding the geologic 

history of extensional faulting. In one place, the Coast 

Range fault and associated faults are overlapped by 

sedimentary rocks of Oligocene and younger age, and in 

another place by sedimentary rocks of Paleocene and 

younger age. The occurrence of detritus derived from the 

Franciscan assemblage in Paleocene and Eocene strata of 

the Coast Ranges (Dickinson, 1966; Berkland, 1973) 

indicates that the lower plate was exposed by the early 

Tertiary. Jayko and others (1987) inferred that uplift of 

the Franciscan assemblage and associated extensional 

faulting in the upper plate occurred during the Late 

Cretaceous and (or) early Tertiary. 

The history of compressional tectonics in the Coast 

Ranges is sparse and varies from place to place. In the 

northern Coast Ranges, thrust faulting and folding began 

during the early Tertiary (Blake and others, 1987; M. C. 

Blake, Jr., oral commun., 1989), and compressional 

deformation is continuing today in rocks of the northern 

Great Valley (Harwood and Helley, 1987). In the southern Coast Ranges, at least four Cenozoic deformations or 

uplifts, indicated by unconformities or eastward-rnigratingdepocenters, have ages of late Paleocene, late Eocene 

to early Miocene, late Miocene, and late Pliocene (Namson and Davis, 1988; Namson and others, 1990; Rentschler and Bloch, 1988). Modem thrust faulting and 

folding still is occurring, as indicated by the 1983 Coaliiga earthquake (see chap. B; Eaton, 1990). 

Landward movement of the Franciscan assemblage as 

a wedge may have even begun in the Mesozoic. In the 

northern Coast Ranges, several northwest-striking 

faults (Paskenta, Elder Creek, and Cold Fork faults) 

offset rocks structurally above the Franciscan assemblage (but not the Franciscan assemblage itself) and 

represent major discontinuities in the dep&itional environment of the Great VaUev seouence (Jones and Irwin. ,~~ ~-- 

1971). These faults, which havedisplacements of tens of 

kilometers to as much as 100 km, are interpreted to have 

moved primarily during the Cretaceous (Jones and Irwin, 

1971), although the latest limit on the time of movement 

is about 3.4 Ma (Hardwood and HeUey, 1987; M.C. Blake, 

Jr., oral commun., 1989). Wentworth and others (1984) 

and Jayko and others (1987) interpreted these faults as 

tear faults in the plate structurally above a wedge of 

Franciscan assemblage. 

In light of the above data and interpretations, we 

postulate (1) that uplift of the Franciscan assemblage and 

extension of the upper plate, consisting of Coast Range 

ophiolite and Great VaUey sequence, occurred during the 

Cretaceous (or, at the latest, during the early Tertiary, 

if Cretaceous movement on the Paskenta-Cold Fork fault 

system is not linked to landward wedge transport) well 

west of the present Diablo Range; and (2) that a tectonic 

wedge of Franciscan assemblage was subsequently driven landward, with the extended upper plate riding 

passively atop it. This wedge is interpreted to have 

moved along a floor thrust fault aligned with the contact 

between the Great Valley sequence and its crystalline 

basement. To the west of the present Diablo Range, 

where movement initiated, the basement was an outboard part of the Coast Range ophiolite. Beneath the 

Great Valley, where the movement is presently occurring, the basement is similar to the Coast Range ophiolite 

but contains numerous younger plutons. A roof thrust 

fault apparently developed only near the east tip of the 

wedge (fig. 8.45); presumably, erosion kept pace with 

uplift near the tip. Differential vertical or horizontal 

movements of the wedge may have produced tear faults, 

such as the Paskenta, Elder Creek, and Cold Fork faults, 

and may have reactivated extensional faults to produce 

complex faults, such as the Tesla-Ortigalita fault. 

PAST AND PRESENT TECTONIC REGIMES 

The Mendocino triple junction has moved northward 

through offshore central California during approximately 

the past 20 Ma, and subduction of the Farallon plate (or 

its derivative) was replaced by transform motion of the 

Pacific plate past North America (see chap. 3; Atwater, 

1970,1989). If tectonic wedging occurred during the late 

Mesozoic and Cenozoic, in association with all of the 

episodes of tear faulting or compression outlined above, 

then clearly it was driven during both subduction and 

transform regimes. At present, it is being driven by a 

transform regime. At least two additional arguments can 

be made that wedge motion-indeed, probably a major 

fraction of wedge motion-occurred during the subduction regime. The first argument is simply based on 

geometry: The east boundary of the Coast Ranges, 

inferred to coincide approximately with the buried tip of 

the wedge, largely parallels Mesozoic structures in the 
222 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

Sierran foothills and the Great Valley rather than the late 

Cenozoic San Andreas fault (Wentworth and Zoback, 

1989; C.M. Wentworth, oral commun., 1990). The second 

argument, developed below, is based on the total apparent displacement of the wedge. 

If the inferred tectonic wedge of Franciscan assemblage extends to the San Andreas fault, as we have 

shown (fig. SAB), then a minimum shortening of about 70 

km has occurred along faults at the top and bottom of the 

wedge in the Diablo Range. Likewise, in the northern 

Coast Ranges, the inferred tear faults in the plate above 

the wedge have a total displacement-and, thus, shortening-of many tens of kilometers (Wentworth and 

others, 1984), possibly as much as 100 km (Jones and 

Irwin, 1971). 

Although a transform regime has replaced a subduction regime in central California over approximately the 

past 20 Ma, plate-margin compression, necessary to drive 

the wedge, has persisted for only approximately the past 

6 Ma (Page and Engebretson, 1984). At about 6.5-4.6 Ma, 

transform motion was also transferred from offshore 

faults to the modem San Andreas fault system (see chap. 

3; Atwater, 1989; Humphreys and Weldon, in press). 

Present plate-margin compression is understandable 

from (1) the slight misalignment of the direction of 

relative plate motion (N. 35' W.; Minster and Jordan, 

1978) and the strike of the San Andreas fault (N. 40' W.), 

and (2) the opening of the Basin and Range province. 

Crouch and others (1984) calculated from these two 

effects a rate of shortening across the Coast Ranges that, 

integrated over the past 5.5 Ma, predicted a total 

shortening of 28 to 72 km. Most of this shortening could 

be accounted for in small fault displacements and folds 

distributed throughout the Coast Ranges (Crouch and 

others, 1984). Thus, the minimum shortening of 70 to 100 

km represented by the tectonic wedge, as discussed 

above, would appear to equal or exceed the maximum 

shortening calculated for the transform regime, a result 

suggesting that some, if not most, of the wedge motion 

occurred during the subduction regime. 

Shear coupling between the subducting plate and 

overlying accretionary prism (Franciscan assemblage) 

could conceivably drive the wedge during the subduction 

regime. Such a mechanism bas been postulated for 

southern Alaska by Fuis and Plafker (in press). To drive 

the wedge during a transform regime appears to require 

a less obvious mechanism, such as plate-margin compression combined with differing deformation in the upper 

and lower crust. Such a mechanism is developed below. 

Sibson (1982) pointed out, on the basis of strength 

considerations, that ductile flow could be expected in the 

middle crust, below the maximum depth of earthquake 

hypocenters. Several workers (Crouch and others, 1984; 

Namson and Davis, 1988; Eaton and Rymer, in press) 

have postulated a decollement near the base of the 

seismicity in the Coast Ranges (avg 15-km depth; see 

chap. 5; Wesson and others, 1973) into which thrust and 

oblique-slip faults on both sides of the Coast Ranges sole. 

They envision differential movement between upper and 

lower crust caused by differing alignment of the transform faults in these two layers, or by shortening of the 

lower crust by ductile thickening. 

We have incorporated the idea of a Coast Range-wide 

detachment in our cross section (fig. 8.45). In the Diablo 

Range, we show a young thrust fault at the base of the 

inferred tectonic wedge soling into the brittleductile 

transition zone, which in this area is, coincidentally, near 

the interface between Franciscan rocks and mafic crust. 

Although we also indicate soling of the San GregorioHosgri fault into such a zone and underthrusting of the 

Salinian block by the early Tertiary accretionary prism, 

focal mechanisms in this region indicate pure strike slip 

on the San Gregorio-Hosgi fault (see chap. 6) and argue 

against this interpretation. Such an interpretation of a 

Coast Ranges-wide midcrustal detachment requires that 

the deformational style and (or) location of the San 

Andreas fault system change from the upper to the lower 

crust. 

If we have correctly inferred the geologic history of 

wedge movement, it is remarkable that such movement 

has apparently occurred in two quite different tectonic 

regimes, a subduction regime and a transform regime. 

SOUTHERN CALIFORNIA 

The crustal structure of southern California is complicated by the Big Bend in the San Andreas fault, situated 

between the Coast Ranges and Transverse Ranges, and 

by onshore spreading centers of the East Pacific Rise, 

situated in the Salton Trough (figs. 8.2, 8.3). The Big 

Bend is thought to result from westward movement of 

the Sierra Nevada relative to the Mojave Desert, along 

the Garlock fault (Hill and Dibblee, 1963). The San 

Andreas fault crosses the Transverse Ranges, between 

the Big Bend and Salton Trough, at an angle oblique to 

relative plate motion, while somehow remaining a largely 

vertical, strike-slip fault. 

The onshore spreading centers in the Salton Trough 

are situated at echelon offsets between the San Andreas, 

Imperial, and Cerro Prieto faults (see fig. 3.8; Lomnitz 

and others, 1970). These three faults are interpreted as 

transform faults, the San Andreas links the northernmost 

spreading center in the Salton Trough with the Mendocino triple junction. A progressive decrease in spreading 

rate northward along the East Pacific Rise is inferred to 

give rise to movement on the San Jacinto, Elsinore, San 

Mi~elNewport-Inglewood, and other faults in southern 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIGREFRACTION AND OTHER DATA 223 

California and Mexico (Lonmitz and others, 1970; Elders 

and others, 1972). 

First, we discuss a transect across southern California, 

Centennial Continental-Ocean Transect C3 (Howell and 

others, 1985). Second, because of the tbree-dimensionality of the geology and tectonics in southern California, we 

include a discussion of block motions, largely from 

Weldon and Humphreys (1986). 

We modify and reinterpret the section of Centennial 

Continent-Ocean Transect C3 (Howell and others, 1985) 

that extends from Santa Catalina Island to the Colorado 

Desert (fig. 8.2). This section of the transect crosses four 

blocks or provinces, the California Continental Borderland (hereafter referred to simply as the "borderland"), 

Peninsular Ranges, Salton Trough, and Chocolate Mountains (fig. 8.6). The transect crosses the NewportInglewood, Elsinore, San Jacinto, and Imperial strikeslip faults. Constraints for the transect include surface 

geology, isotopic studies, seismic-refraction profiling 

(which is sparse, except in the Salton Trough), tomographic studies, and potential-field studies. 

BORDERLAND 

The borderland is broken up by right-slip faults into 

several northwest-trending blocks. Our cross section (fig. 

8.6) begins on the easternmost block, the "Catalina 

terrane" (HoweU and others, 1985). bounded on the east 

by the Newport-Inglewwd fault. The Catalina terrane is 

underlain, beneath patches of Tertiary volcanic rocks, by 

Franciscan assemblage, on the basis of outcrops on Santa 

Catalina Island (Platt, 1975. 1976; Jones and others, 

1976) and submarine dredge and core samples (Vedder 

and others, 1974). The block west of the Cataliiterrane, 

the "San Nicholas terrane" (Howell and others. 1985). is 

inferred to be underlain, beneath Cenozoic mahne &imentaly rocks, by rocks similar to the Great Valley 

sequence and Coast Raqge ophiolite of central California, 

possibly in fault contact with Franciscan assemblage at 

depth (Vedder and others, 1974). 

A reversed seismic-refraction profile just west of Santa 

Catalina Island indicates P-wave velocities of 5.8 km/s to 

6-km depth and of 6.7 km/s to the Moho at about 24-km 

depth (fig. 8.6A, Shor and Raitt, 1968). This velocitydepth section is similar to that for the Diablo Range of 

central California (see above), where Franciscan rocks 

are equated with the 5.8-km/s interval, and middle and 

lower crust of island arc(s) and (or) oceanic crust are 

equated with the 6.7-km/s interval. In this region, there 

is no clear evidence of landward movement of the 

Franciscan assemblage as a tectonic wedge, although 

such evidence may surface during future investigations. 

As in the Diablo Range, the lower crust must have been 

brought to its present 18-km thickness by (1) imbrication 

of slices of island-arc crust, (2) tectonic underplating of 

several thicknesses of oceanic crust, and (or) (3) magmatic underplating. Subduction continued beneath the borderland until sometime between 30 and 20 Ma (see 

Atwater, 1970). depending on the latitude to which the 

borderland is palinspastically restored. 

PENINSVLAR RANGES 

The Peninsular Ranges are underlain in the west by 

supracrustal rocks, including, from top to bottom, Cenozoic marine sedimentary rocks, Cretaceous forearc sedimentary rocks. Loweri?) Cretaceous and Upper Jurassic 

andesite (Santiago Peak Volcanics), and Middle Jurassic 

flysch (Bedford Canyon Formation) that was disrupted 

and overturned before the Late Jurassic (Lamen, 1948; 

Jennings, 1977; Criscione and others, 1978). These rocks 

are intruded by Early Cretaceous plutons of the Peninsular Ranges batholith that include chiefly tonalite and 

gabbro and show no special age trends (static magmatic 

arc; Silver and others, 1979). About 80 km east of the 

coastline, both prebatholithic and batholithic rocks 

change (fig. 8.6A): To the east, the prebatholithic rocks 

are dominantly metamorphosed clastic rocks of amphibolite grade, and the batholithic rocks are chiefly tonalite 

and granodiorite whose ages decrease progressively 

eastward (from 105 to 80-90 Ma; migrating magmatic arc; 

Silver and others, 1979). Major-element chemistry and 

oxygen isotopes indicate that deep crustal rocks in the 

west half of the batholith are dominantly primitive and 

tholeiitic but, in the east, more aluminous and oxidized 

(fig. 8.6A). Older crust that was once at the Earth's 

surface is inferred at depth in the east (Silver and others, 

1979). 

Seismic constraints for the deep structure of the 

Peninsular Ranges are sparse. Using blasts at the 

Corona Quarry in the northernmost Peninsular Ranges, 

Gutenberg (1951) and Shor (1954) obtained anurnversed 

refraction profile, extending southward to the United 

States-Mexican border, along with a reflection record at 

the blast site. Interpretation of these data by Shor and 

Raitt (1968) indicated velocities of 5.9 km/s to 8-km 

depth, 6.8 kmls to 26-km depth (with a possible lowvelocity zone in this interval), and 7.0 km/s to the Moho 

at 30- to 32-km depth (fig. 8.6A). In contrast, a study by 

Nava and Bnme (1982) using a hlast at the same quarry, 

reversed by an earthquake in Baja. Mexico, indicated a 

Moho depth of 42 km. Hem and Clayton (1986a, b) used 

as many as 600,000 arrivals from local earthquakes in 

southern California to map the velocity of the crust and 

upper mantle, using tomography. Their map indicates 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

SEA' 

IEVELÃ‘ 

10 - 

& 

1; 20- 5 

5 

z 

i 30- Â 

0 

40 - 

that the west half of the Peninsular Ranges has a higher 

average upper-crustal velocity and a lower average 

mantle velocity in comparison with the east half. Their 

map of Pn delays for the Peninsular Ranges suggests no 

crustal root and an average crustal thickness of nearly 30 

CALIFORNIA 

CONTINEtfrAL 

BORDERLAND 

ton. Gravity modeling of the Peninsular Ranges (Fuis and 

others, 1984) and isostatic calculations also indicate a 

maximum crustal thickness of 30 to 33 km. In our cross 

section (fig. 8.6B), we adopt a maximum crustal thickness 

of 33 km. 

PENINSULAR RANGES 

FIGURE 8.6.-Crustal structure of southern California. A, Surface geology, west halfinferred to be underlainat depth by mafic (islandarcor oceanic) 

isotope data, and models of seismic-refraction, gravity, and magnetic data crust and east half inferred to be underlain at depth by intermediate 

for part of CeirtennM Continent-O^an Transect C3 (see Howell and otherB, continental Precambrian(?) rocks; (3) late Cenozoic rift, the Salton 

1985). B. Reinteroretation of Transect C3. Maim features in fie-ure 8.66 Troueb. whose central Dart is inferred to be underlain bv entirely new 

include, from west to east. (1) ~ranciscanassemb~a~e overlying mafic crust mithat includes, from top to bottom, sedimentary &ks, thermally 

in the borderland; (21 Peninsular Ranges batholithic block, consisting of metamorphosed sedimentary rocks, and gabbm generated at onshore 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIC-REFRACTION AND OTHER DATA 225 

SALTON TROUGH 

An additional constraint on crustal structure is the 

modeling by Jachens and others (1986; R.C. Jachens, 

written commun., 1988) of strong magnetic and gravity 

steps (500 nT and 40 mGal, respectively) in the central 

Peninsular Ranges: A moderately east dipping boundary 

CHOCOLATE 

MOUNTAINS 

is modeled between more magnetic, dense rocks on the 

west and less magnetic, lighter rocks on the east. This 

boundary is poorly defined at the latitude of our transect; 

it correlates approximately (within 15 km or so) with the 

boundary between the east and west halves of the 

SAN JACJNTO 

FAULT 

I CenozdO 

roSr/BBsf -1-- 

<.m >.m 

D IMPERIAL 

FAULT 

BRAWLEY 

FAULT 

spreading center; and (4) Pelona-Orocopia schist of Haxel and Dillon (1978) primarily in interpreting mafic crust at shallower depths beneath the 

(similar to the Franciscan assemblage), interpreted to compose tectonic borderland and western Peninsular Ranges (5-8 km versus 11-15 km) to 

wedge. Tectonic wedge in feature 4 is postulated to have beenobducted onto better match seismic and potential-field resulk See figures 8.2 and 8.3 

continental cruet (see text); its tip would lie well east of east end of cross far location of Transect C3: see figure 8.4 for explanation. No vertical 

section This reinterpretation differs from Howell and others' (1985) exagaeration. 
226 THE SAN ANDEEAS FAULT SYSTEM. CALIFORNIA 

Peninsular Ranges batholith, as discussed above (fig. 

8.6A). In the cross section (fig. 8.65). we interpret an 

eastward deepening of mafic rocks, including prebatholithic and (or) batholithic mafic rocks (gahbro, diabase, 

and metamorphic rocks), along this magneticlgravity 

boundary. R.C. Jachens (oral commun., 1989) indicated 

that, in some places, this boundary is so planar as to be 

interpretable as a fault. As beneath the borderland, the 

mafic rocks beneath the Peninsular Ranges may have 

reached their current thickness by thrust imbrication, 

tectonic underplating, or magmatic underplating. We 

speculatively show some tectonic underplating on the 

west side. 

SALTON TROUGH 

The Salton Trough is the landward extension of a 

ridgeltransform-fault system, the East Pacific Rise, of 

the Gulf of California (see fig. 3.13). This system became 

well established during the late Cenozoic (approx 5 Ma) 

as the plate boundary jumped inland from offshore Baja 

California (Atwater, 1970, 1989; Humphreys and Weldon, in press). 

The Salton Trough is underlain by upper Cenozoic 

sedimentary rocks and minor amounts of volcanic rocks, 

which are exposed chiefly around its edge and are 

penetrated in wells. Onset of rifting and major subsidence in the Salton Trough was followed by marine 

incursion during the latest Miocene to late(?) Pliocene, as 

indicated by the Imperial Formation (Dibblee, 1954; 

Powell, 1984). The thick Cenozoic sedimentary section is 

offset by Quaternary faults, both exposed and buried, 

and is intruded by Quaternary volcanic rocks, both silicic 

rocks that form volcanoes at the two inferred onshore 

spreading centers (fig. 8.7) and mafic rocks that are 

penetrated in geothermal wells (Elders and others, 1972; 

Robinson and others, 1976). Faulting in the Salton 

Trough occurs primarily on conjugate northwest- and 

northeast-striking faults and is largely strike slip 

(Johnson and Hadley, 1976; Johnson, 1979; Fuis and 

others. 1982). North-south-striking faults, however, such 

as the north end of the Imwrial fault. the Brawlev fault. 

and north-south-striking seismicity lineaments (that outline inferred spreading centers; figs. 8.1, 8.7). have 

normal components and lead to the subsidence that 

ultimately created the Salton Trough. Earthquake hypocentral depths indicate that brittle fault motion extends 

to about 12-km depth in the Imperial Valley but deeper 

in the adjacent Peninsular Ranges along the San Jacinto 

fault (Doser and Kanamori, 1986). 

Detachment faulting on the east flank of the Salton 

Trough, in the Chocolate Mountains and other ranges, 

preceded the Pliocene and later basin-forming tectonics 

in the Salton Trough (Diion, 1976; Berg and others, 1982; 

Frost and others, 1982). Similar faulting on the west 

lank of the Salton Trough, however, may have both 

receded and overlapped in time the tectonics in the 

Mton Trough (Wallace and English, 1982; Schultejahn, 

984; Isaac and others, 1986). 

Biehler and others (1964) and Fuis and others (1982, 

,984) demonstrated from seismic surveys that the sedinentary rocks (1.8-5.5 Ws) in the central Salton 

['rough are as much as 5 km thick (fig. 8.6A). Below 5-km 

lepth, a low-velocity (5.6 km/s) "basement," which is not 

eparated from the overlying sedimentary rocks by a 

relodty discontinuity, is inferred to be metamorphosed 

greenschist fades) sedimentary rocks (Fuis and others, 

.982, 1984); this "basement" layer extends to 12-km 

lepth. High heat flow in the Salton Trough (see Lachenmch and others, 1985) is inferred to cause the metanorphisrn of the sedimentary rocks. Thus, the entire 

iection of inferred upper Cenozoic sedimentary rocks, 

netamorphosed and unmetamorphosed, is as much as 12 

mi thick. 

Below 12- to 14km depth in the Salton Trough, a 

ligh-velocity (7.1-7.2 km/s) "subbasement" that is indiÂ¥at by seismic-refraction data (fig. 8.6A) is inferred to 

>e gabbro generated at one of the nearby spreading 

enters (Fuis and others, 1982, 1984). Modeling of 

leismic-refraction and gravity data indicate that the 

Eoho in the central Salton Trough is 23 to 28 km deep 

Fuis and others, 1982, 1984). The central Salton Trough 

s interpreted to be underlain entirely by late Cenozoic 

Â¥rus (fig. 8.65). 

Buried scarps separating old crust (plutonic and metunorphic rocks; 5.9-6.0 km/s) from new cruet (sedimenary and basaltic rocks; 1.8-7.2 Ws) are visible by 

seismic methods on both sides of the Salton Trough (Fuis 

ind others, 1982, Fuia and Kohler, 1984). On the west 

tide of the rift, where the new-crust'old-crust boundary 

s ragged in outline (fig. 8.7). we interpret normal faults 

fr- 

~IGURE 8.7.-Tectonic block motion in southern California (modified 

from Weldon and Humnhrevs. . .. 1986. . and Humnhrevs . . and Weldon. in 

press). Various blocks (italicized names near motion vectors) move 

through region where the San An* fault irendfiobliquely Uiplatv 

motion, between the Big Bend and the Salton Troueh, without major 

convergence with each other. 'Rirough this region they move 

counterclockwise, following nearly concentric arcs (arcs and radii, 

thinred lines). New crust, which is fonninein wake of the Salton and 

Penis bl& in the salton Trough, is created by sedimentary-basin 

fill and gabbmic inuii&ioiw at onshore qireading centers, outlined by 

seismicity lineament*. High-velocity mantle beneath the Trantiveree 

Ranges ie interpreted as cold, sinking lithosphetic mantle, and 

low-velocity mantle beneath the Salton Trough as hot upwelling 

asthenosphere or lithospheric mantle containing partial melt (Humphrey~ and others, 1984; Humphreys and Clayton, in press). Motion 

vectors for the Mojave Desert and Sierra Nevada modified to 

incorporate results of Sauber and others (1986). 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIC-REFRACTION AND OTHER DATA 227 

(fig. 8.65); on the east side, where this boundary is inferred to have originated by pullaway from the Cerro 

linear, we interpret a strike-slip fault. In our cross F'rieto spreading center to the southeast; the fault on the 

section, faults on the west side of the Salton Trough are east side is inferred to be a largely passive suture (figs. 

EXPLANATION 

+ Mdon vectm olteclmic Mock- 0 bw.wty 

vector scale Relative 10 Nonh America 

20 mm/yr ------ Scismidty lineamen1 

NW CrmI (late c-k) - Fault-Dashed where approximately 

a o~"2spr~~~~ - located: doited where buried 

Thrust fault-Dasted where 

High-velocity mantk approximately locaied. Sawteeth on upper plate 
228 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

8.65, 8.7; Fuis and others, 1982). A similar rift configu- central and northern California after being extended 

ration is seen, for example, in the Gulf of Elat (Gulf of during uplift of the Franciscan assemblage, or (2) rocks of 

Aqaba, Red Sea; Ben-Awaham, 1985). the Great Valley sequence which were peeled up along 

CHOCOLATE MOUNTAINS 

Rocks on the east flank of the Salton Trough are 

igneous and metamorphic rocks that compose two or 

more fault-bounded packages, or tectonostratigraphic 

terranes (see Howell and others, 19%). A complex of 

metasedimentary and mafic metaigneous rocks described 

by Dillon (1975) may include two Precambrian terranes, 

the Joshua Tree and San Gabriel terranes, described 

farther north by Powell (1981). This complex is intruded 

by intermediate to felsic Mesozoic plutons and rests on 

the low-angle Chocolate Mountains thrust fault above the 

(informal) Pelona-Orocopia schist of Haxel and Dillon 

(1978; see also Haxel, 1977). The Pelona-Orocopia schist 

consists chiefly of metagraywacke and lesser metapelite, 

metabasite, metachert, marble, and serpentinite (albiteepidote-amphibolite facies) of uncertain but probable late 

Mesozoic or early Tertiary age (Conrad and Davis, 1977; 

Miller and Morton 1977, 1980). It resembled the Franciscan assemblage but lacks melange. 

Many workers have speculated on the depositional 

environment and origin of the Pelona-Orocopia schist. 

Haxel and Dillon (1978) postulated formation in an 

ensimatic rift basin with continent on both sides-not 

unlike the current Salton Trough. Powell (1981) favored 

an origin as a parautochthonous continental-marginal 

deposit. In any case, from its quartz content, the 

Pelona-Orocopia schist dearly originated near a continent and incorporated continental detritus. It was thrust 

beneath the continental metasedimentary-metaigneous 

complex some time after Mesozoic plutonism (80 Ma; 

Powell, 1981) and before Oligocene volcanism (35 Ma; 

Crowe 1978, Crowe and others, 1979). The thrust fault 

may have been reactivated one or more times as a 

low-angle normal, or detachment, fault (Frost and others, 1982). 

Evidence from refraction profiling in the western 

Mojave Desert across the Rand schist, which has been 

correlated with the Pelona-Orocopia schist (Ehlig, 1968). 

indicates relatively low-velocity crust beneath this body 

(max 6.4 kmls; Fuis and others, 1986) that we infer to be 

continental crust. We speculate that the Pelona-Qrocopia 

schist also rests on continental crust and that the Rand 

and Pelona-Orocopia schists were emplaced as a tectonic 

wedge into continental crust in a manner similar to the 

Franciscan assemblage of central and northern California. We hypothesize that the metasedimentary-metaigneous complex structurally above the schist is analogous 

to either (1) rocks of the Coast Range ophioliteJGreat 

Valley sequence which rode passively atop the wedge in 

backthrust faults during landward movement of the 

wedge. In southern California, tectonic wedging dearly 

occurred before the present transform regime, presumably during subduction of the Farallon plate (or its 

derivative). The geologic data discussed above indicate 

that the Salton Trough has undergone extension, rather 

than compression, for approximately the past 5 Ma 

(probably even longer; see Humphreys and Weldon, in 

press). 

Crustal thickness is unknown in the Chocolate Mountarns; however, the Colorado Desert, to the east and 

north, has a generally thin (26-28 km) crust (fig. 8.3) and 

a local root (82 km deep) under the Whipple Mountains 

metamorphic-core complex (Fuis, 1981; Jill McCarthy, 

written commun., 1988). 

TECTONICS-THE THREE-DIMENSIONAL PICTURE 

The geology and, presumably, the deep structure of 

southern California illustrated along transect C3 (fig. 8.6) 

is grossly two dimensional as far north as the Transverse 

Ranges. In the Transverse Ranges, the rocks on the 

southwest side of the San Andreas fault are similar to 

those in the Chocolate Mountains. These rocks are 

bounded on the south and west by older, deformed 

strands of the San Andreas fault system (fig. 8.7; Powell, 

1981). The tectonics also changes in the Transverse 

Ranges: Crustal-block motion swings to the west to 

follow the trend of the San Andreas fault, as discussed 

below. 

Using Quaternary geologic and geodetic evidence, 

Weldon and Humphreys (1986) documented complex 

motion of crustal blocks in southern California that is not 

simply predictable from the motion vectors of the Pacific 

and North American plates. These motion vectors predict a large component of convergence across the San 

Andreas fault in the Transverse Ranges between the Big 

Bend and the Salton Troueh - (fie. -- 8.7). For a total offset . 

on the San Andreas fault system of about 300 km (Hill and 

Dibblee, 1963; Crowell, 1962, 1981; Powell, 1981), a 

maximum of 45 krn of uplift in the Transverse Ranges 

would be expected (Weldon and Humphreys, 1986). 

However, the preservation in the Transverse Ranges of 

upper Cenozoic sedimentary rocks and of offset bedrock 

features on either side of the San Andreas fault argues 

against such major uplift and associated consumption of 

crust, as does the relatively minor crustal root in the 

Transverse Ranges (fig. 8.3). Weldon and Humphreys 

(1986) constructed a kinematic model in which crustal 

blocks between the San Andreas fault and a system of 

borderland and other offshore faults rotate counterclock- 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIC-REFRACTION AND OTHER DATA 229 

wise. parallel to the San Andreas fault, between the 

Salton Trough and the Big Bend (fig. 8.7). Approximately two-thirds of the relative northwestward motion of the 

Pacific plate past the North American plate is taken up 

by the San Andreas fault system, including the San 

Jacinto fault; approximately one-third of it is taken up by 

the Elsinore fault, a system of borderland faults, and 

offshore faults in central California, including the San 

Gregorio-Hosgri fault (fig. 8.7); and only a minor fraction 

of it is taken up within the blocks (see Humphreys and 

Weldon, in press). 

A marked advance in the P-wave traveltimes of 

teleseismic arrivals in southern California is associated 

with the Transverse Ranges and extends across the San 

Andreas fault (Hadley and Kanamori, 1977; Raikes, 

1980). Tomographic analysis of this anomaly indicates 

that it results from a vertical slablike region of relatively 

high velocity in the mantle which extends downward as 

far as 250 km (Humphreys and others, 1984; Humphreys, 

1985; Humphreys and Clayton, in press). The amount of 

velocity increase, a maximum of 3 percent, is most 

reasonably explained by a thermal difference in the 

mantle. This velocity increase, coupled with a velocity 

decrease in the upper 90 km or so of mantle beneath the 

Salton Trough, led Humphreys and Hager (1984 and in 

press) to infer small-scale mantle convection between the 

Salton Trough and the Transverse Ranges. This convection involves passive rising of asthenosphere beneath the 

Salton Trough and cooling and sinking of lithosphere 

beneath the Transverse Ranges. The vertical extent of 

the inferred lithospheric slab beneath the Transverse 

Ranges, 250 km, is similar to the 300-km estimate of total 

offset along the San Andreas fault system. However, 

because the cooled mantle slab extends across the San 

Andreas fault, most of the mantle seems to be moving 

independently of the crust (fig. 8.8; Hadley and Kanamori, 1977; Humphreys and others, 1984, Humphreys, 

1985; Humphreys and Hager, in press). The horizon of 

decoupling is apparently at or below the Moho because 

crustal material is not entrained in the slablike feature. 

Additional decoupling may be occurring in the crust, 

similar to that postulated for central California (Yeats, 

1981; Webb and Kanamori, 1985). Decoupling at the 

Moho requires that the defonnational style and (or) 

location of the San Andreas fault system change from the 

crust to the mantle (fig 8.8). We note that mantle drag on 

the crust is required to maintain the Big Bend in the San 

Andreas fault because plate-edge forces alone would tend 

to "short-circuit" the San Andreas fault south of the Big 

Bend and cause most plate motion to be taken up on the 

San Jacinto, Elsinore, or more westerly faults (Kosloff, 

1978; Humphreys, 1985). 

To summarize, block motions in the region between the 

Big Bend and the Salton Trough result in only minor 

interblock convergence in the crust. In contrast, major 

convergence in the lithospheric mantle is indicated by the 

presence of an inferred, sinking lithospheric slab. 

STRUCTURE OF THE UPPER MANTLE 

In addition to the Transverse Ranges and Salton 

hugh, other regions in California show mantle velocity 

anomalies that imply structure within the lithospheric 

mantle and even the asthenosphere. The seismic networks in California (see chap. 5) provide an abundant 

source of regional earthquake and teleseismic arrivals 

Slat have been used to determine this upper-mantle 

itructure. 

A detailed study of the compressional-wave velocity of 

Ae uppermost mantle in central California reveals a 

normal velocity of about 8.0 fan/s and no evidence for 

SALTON 

TROUGH TRANSVERSE 

Hole or region 

1 . LITHOSPHERIC MANTLE 

- 

QGUBE 8.8.-Motion of crustal blocks in southern California (open 

mm; see fig. 8.7) and somewhat different motion of lithospheric 

mantle below (solid arrows) (modified from Humphreys, 1985, and 

Humphreys and Hager, in press). Mantle convection cell is envisioned between the SaltonTroughand the Transverse Ranges. Crust 

and lithospheric mantle appear to be moving independently of one 

another, as the San Andreas fault trends obliquely across region of 

inferred, sinking lithaspheric mantle beneath the Transverse Ranges 

(see fig. 8.7). Small arrows, relative fault motion; sawteeth, upper 

plates of crustal thrust faults; crosslines, subduction zones in lithospheric mantle. 
280 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

velocity anisotropy (Oppenheimer and Eaton, 1984). A 

similar study in southern Californiafinds nearly the same 

average velocity, 7.95 kds, with evidence for 2-percent 

velocity anisotropy (Vetter and Minster, 1981; Hearn, 

1984). The fast direction is N. 75' W., approximately 

parallel to the San Andreas fault in southern California. 

Seismic-velocity anisotropy in the upper mantle has been 

reported elsewhere, notably in oceanic crust, and is 

commonly attributed to alignment of olivine in the mantle 

along a shear-stress direction (Bamford and others, 

1979). In southern California, this shear would presumably be that associated with the motion of crustal blocks 

above the lithospheric mantle. 

Lithospheric thickness along the San Andreas fault has 

been investigated by using delay times of teleseismic 

arrivals and thermal models (Zandt and Furlong, 1982). 

These studies indicate a lithospheric thickness of only 30 

to 60 km for much of western California, and as little as 

20 km for northern California just south of Cape Mendocino. These lithospheric thicknesses contrast with averages of60 to 80 km for the Western United States and 120 

to 170 km for the Central and Eastern United States 

(lyer and Hitchcock, 1989). The thinness of the lithosphere in northern California south of Cape Mendocino is 

due to the creation of the San Andreas fault systemitself: 

The transform fault is lengthening as the Mendocino 

triple junction migrates northward. As this junction 

migrates northward, the west edge of North America is 

sliding off the edge of the northward-moving, subducting 

Gorda plate, thereby creating a "windown where no 

snbducting lithospheric slab is present (Dicldnson and 

Synder, 1979). In this slabless window, the North American crust is initially in direct contact with the asthenosphere that has welled upward to fill the hole left by the 

Gorda plate (Zandt and Furlong, 1982). This geometry 

produces the thinnest lithosphere in California and, 

probably, in North America. In contrast, the lithosphere 

is abnormally thick (250 km) in the Transverse Ranges, 

where "subduction" of lithospheric mantle is occurring, 

as discussed above. 

Velocity anomalies appear to extend even into the 

asthenosphere beneath western California. Aid (1982) 

summarized the results of Cockerham and Ellsworth 

(1979) and Raikes (1980) in a combined velocity-anomaly 

model for a depth range of about 100-226 km in the 

mantle (fig. 8.9). Aid suggested that the low-velocity 

region in central California is hot, mobile material 

associated with the slabless window. Such an association 

appears likely for the northwest-trending prong of this 

anomaly, as refined by the recent work of Benz and 

others (1990); however, the center of the anomaly, 

located near Long Valley caldera (figs. 8.2, 8.9), apparently has a different origin. Low-velocity regions are also 

associated with the Salton Trough, where asthenospheric 

upwelling is inferred, and the eastern Mojave Desert, 

where crustal extension has occurred. The high-velocity 

region that crosses the San Andreas fault in southern 

California is similar to the one discussed above (fie. 8.7). 

SUMMARY 

The crust along the San Andreas fault system thickens 

from about 16 km at Cape Mendodno, in northern 

California, to about 30 km in southern California and thus 

is significantly thinner than the average thickness (36 

km) for the conterminous United States. Lithospheric 

thickness (20-60 km) is also substantially less along most 

~f the San Andreas fault system than is typical for 

continental areas (60-170 km). The lithosphere is thiinest at both ends of the fault system, at the Mendocino 

triple junction on the north, where the North American 

plate is sliding off the edge of the Gorda plate as it moves 

northward, and in the Salton Trough on the south, where 

0 - 200 KILOMETERS 

FIGURE 8.9.-Seismic-velocity anomalies in upper mantle (chiefly 

~8th-phere), derived from teleseismic delay-time data (Aid, 

1982), for depth ranges 125-225 km (central California; Cockerham 

and Ellsworth, 1979) and 100-180 km (southern California; Raikes, 

1980). H, high velocities (contours solid); L, low velocities (contours 

dashed). Contour interval, 2 percent. High seismic velocities that 

cross the San Andreas fault in southern California are similar in 

pattern to those shown in figure 8.7. Heavy lines, major faults. 
8. LITHOSPHERIC STRUCTURE AND TECTONICS FROM SEISMIGREFRACTION AND OTHER DATA 231 

onshore spreading centers of the East Pacific Rise are 

generating new crust in a rift between the North 

American and Pacific plates. In contrast, the lithosphere 

is abnormally thick (260 km) in the Tranverse Ranges, 

where "subduction" of lithospheric mantle is occurring. 

The crust of central California was formed at an 

Andean-type continental margin and has been modified 

by large offsets along strike-slip faults of the San 

Andreas fault system. East of the San Andreas fault, the 

Andean-marginal sequence includes a subduction complex (Franciscan rocks), a forearc basin (Great Valley 

sequence), and a magmatic arc (plutons of the Sierra 

Nevada). The subduction complex appears to have been 

emplaced as a tectonic wedge beneath sedimentary rocks 

of the forearc basin. West of the fault, displaced blocks 

constitute an Andean-marginal sequence that has been 

shortened by strike-slip faulting. 

The tectonic wedge of Franciscan rocks east of the fault 

is reinterpreted to extend from its tip beneath the Great 

Valley all the way to the San Andreas fault. This 

interpretation is motivated by the apparent continuity 

between crystalline basement rocks beneath the Great 

Valley and mafic rocks at midcrustal depths in the Diablo 

Range, beneath the Franciscan rocks. The presence of 

extended crust atop the tectonic wedge (outliers of Coast 

Range ophiolite and Great Valley sequence) has led us to 

propose the following tectonic evolution for the wedge. 

(1) Franciscan rocks were uplifted and upper-plate rocks 

(those above the subduction zone) were extended during 

the Cretaceous (or, possibly, early Tertiary) well west of 

their current position in the Coast Ranges. (2) The 

Franciscan rocks and overlying extended crust were 

subsequently forced landward during one or more episodes in the form of a wedge that largely followed the 

contact between Great Valley basement and the Great 

Valley sequence. (3) Wedge movement began during the 

subduction of the Farallon plate (or its derivative) 

beneath central California, however, it apparently is also 

occurring at present, in the San Andreas transform 

regime. Present movement is interpreted to result from 

compression across the San Andreas fault system coupled 

with differential motion between the upper and lower 

crust; this differential motion is interpreted to occur on 

thrust fault(s) at the base of the wedge that sole into the 

brittle-ductile transition zone. 

The crustal structure in southern California shares 

several features in common with central California, 

including, west of the fault, an Andean-marginal sequence that has been shortened or, at least, shuffled by 

strike-slip faulting, and, east of the fault, subductioncomplex rocks that are inferred to have moved landward 

as a tectonic wedge into the continental rocks. However, 

major differences are apparent in southern California. 

First, east of the San Andreas fault, the Andeanmarginal sequence is incomplete: A forearc basin is 

absent, and the magmatic arc is diffuse. Second, new 

continental crust bas formed in the Salton Trough, an 

active crustal pullapart basii, by a combination of rapid 

sedimentation, metamorphism, and magmatic intrusion 

at the onshore spreading centers. In addition, the motions of the crust and lithospheric mantle differ in 

southern California: The crust is moving as a collage of 

blocks, with only minor interblock convergence, whereas 

the lithospheric mantle is converging and "subducting" 

beneath the Transverse Ranges. 

The interpretations of (1) a midcrustal detachment in 

the brittle-ductile transition zone in central California 

and (2) a crust-mantle detachment in the Transverse 

Ranges of southern California would appear to require 

that the deformational style and (or) location of the San 

Andreas fault system change with depth in these regions. 

The properties of the lithosphere along the San Andrea~ fault are not at all typical of continental areas, and 

further characterization of these properties presents a 

significant scientific challenge. 

ACKNOWLEDGMENTS 

Stimulating discussions with C.M. Wentworth and 

M.C. Blake, Jr., led to some of our ideas concerning the 

configuration and emplacement history of the inferred 

wedge of Franciscan rocks in central and northern 

California, A. Griscom and R.C. Jachens provided insightful discussions regarding interpretation of potentialfield models, and R.C. Evarts assisted us in drawing a 

cross section through the uppermost crust in the northeastern Diablo Range. We are also indebted to E.D. 

Humphreys for providing preprints of manuscripts on 

lithospheric motion in southern California. 

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Washington. American Geophysical Union, p. 183-196. 


s titdies on the gravity and magnetic fields of the San Andreas 

fault system span more than SO years, but only recently have 

the fundamental data sets become adequate to provide a general 

view of the entire system. Modeling these new data defines the 

three-dimensional geometries of the faults and helps unravel the 

tectonic history of the system by "seeing thmughn the relatively thin 

cover of young sedimentary deposits and water to the older rocks 

below. 

9. CRUSTAL AND LITHOSPHERIC STRUCTURE 

FROM GRAVITY AND MAGNETIC STUDIES 

By ANDREW GRISCOM and ROBERT C. JACHENS 

CONTENTS 

. . Sources of gravity and magnetic anomalies-------- 241 

Geometry of faults in the San Andreas System------ 243 

PlanviewÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘ 243 

AttitudeÃ‘Ã‘Ã‘Ã‘Ã‘-Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘ 

ch-t&stic+Ã‘Ã‘Ã‘Ã‘Ã‘Ã ------- 249 

Offsets of anomaliesÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘ ------ 249 

Implications for plate &+~nic+~~~~-~~~~~~-- 252 

~endocjno triple junctionÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘ 252 

Salton Buttes spreading centev------------------- 255 

ReferencescitedÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘--Ã‘ 256 

INTRODUCTION 

Studies of the San Andreas fault system using the 

Earth's gravity and magnetic fields began before 1960 

but received their main impetus during the 19701s, when 

work on the possibility of predicting earthquakes on this 

system began in earnest. Early investigations focused 

4 FIGUKE 9.1.-Magnetic map of the Western United States and eastern 

Pacific Ocean, showing locations of major plate boundaries: solid line, 

present boundary; dashed line, former boundary; double line, spreading ridge; single line, transform fault; toothed line, subduction-zone 

fault or transpressional fault (sawteeth on upper plate). From 

mainly on short segments of the faults because only 

limited data were available. More extensive potentialfield data sets that have been published in recent years 

now permit the gravity and magnetic expression of the 

entire fault system to be viewed in a regional context (fig. 

9.1). 

Gravity and magnetic data reflect, respectively, the 

density and magnetization of the rocks beneath the 

surface; and, in many situations, these properties can be 

closely correlated with the rock types seen in outcrop. 

Anomalies in the Earth's gravity and magnetic fieldsfor example, local deviations of the measured fields from 

those predicted on the basis of simplified Earth modelsprimarily reflect lateral variations in density and magnetization that generally are not included in such simple 

models. These anomalies can be intemreted aualitativelv 

to infer the general spatial distribution of rock types in 

the subsurface, and quantitatively, through the use of 

efficient computerbased modeling techniques (Saltus 

and Blakely, 1983; Chuchel, 1985; Blakely and Simpson, 

1986), to determine the geometries and specific locations 

of concealed rock bodies. Although all such interpretations are nonunique, both because many different distributions of density and magnetization can give rise to 

identical anomalies and because density and magnetizaGeological Society of America (1987); used with permission. Plate 

boundaries from King (1969), McCulloch (1987), and Wilson (1989). 

Each color band represents 100 nanotealas; values ranee from low 

(blue) to high (red); white area, no data. Bathymetric contours in 

meters. 

239 

THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA; U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
240 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

tion do not uniquely define a specific rock type, the 

combined use of gravity and magnetic data with geologic, 

geochemical, andother geophysical data can be especially 

effective in limiting the number of acceptable interpretations. 

In the sections below, we first present regional gravity 

and magnetic maps covering the San Andreas fault 

system and briefly discuss the sources, compilation 

methods, and limitations of the data from which they 

were produced and, in general terms, the sources of the 

anomalies shown on them. We then summarize the 

results of individual studies of sections of the major faults 

in the system and attempt to synthesize these results in 

terms of the geometries of the faults, the structures and 

rock types in the surrounding areas that are related to 

the faults, and the properties of the fault zones. Next, we 

focus on studies that relate to movement on the faults, 

including constraints on total displacements. Finally, we 

discuss the platetectonic implications of potential-field 

investigations of the fault system. 

ISOSTATIC RESIDUAL GRAVITY MAP 

An isostatic residual gravity map of the region surrounding the San Andreas fault system is shown in 

figure 9.2. We have chosen to present the gravity data 

in this form rather than in terms of the more common 

Bouguer or free-air gravity because of the generally 

closer correlation between isostatic residual gravity and 

mapped geology (Jachens and Griscom, 1986; Simpson 

and others, 1986). Most long-wavelength anomalies (longer than approx 250 km) on a Bouguer gravity map are 

caused by deep-seated density distributions that buoyantly support the topography in a manner consistent with 

the principle of isostasy (Simpson and others, 1986). 

Bouguer gravity anomalies related to isostasy are prevalent in California because of the extreme topographic 

relief in the State (Oliver, 1980; Jachens and Griscom, 

1985), and they are particularly strong near the coast, 

where an eastward to northeastward decrease in gravity 

reflects the transition from thin oceanic crust to thicker 

continental crust. In such areas as California, the Bouguer gravity anomalies associated with isostatic support 

of topography are so strong that they tend to distort or 

even mask the lower-amplitude anomalies caused by 

density distributions in the middle to upper crust, those 

anomalies most easily correlatable with rocks exposed at 

the surface (Jachens and Griscom, 1985). Our isostatic 

residual gravity map has these long-wavelength isostatic 

effects removed, at least to first order. We emphasize 

that the anomalies remaining on our map are predominantly caused by lateral density variations in the middle 

to upper crust and, as such, do not represent areas that 

are out of isostatic balance (Jachens and Griscom, 1985). 

Our isostatic residual gravity map (fig. 9.2) is based on 

the new isostatic residual gravity map of the conterminous United States by Simpson and others (1986), who 

presented a detailed discussion of the data sets and 

procedures used to generate this map. The basic gravitydata set was compiled for the "Gravity Anomaly Map of 

the United States" (Society of Exploration Geophysicists, 1982) and includes 1 million on shore and 0.8 million 

offshore gravity observations. These data were sampled 

on a rectangular grid with a grid spacing of 4 km, 

containing Bouguer gravity values onshore and free-air 

gravity values at sea (Godson, 1986). To produce our 

isostatic residual gravity map, the offshore hair 

gravity values were converted to Bouguer gravity values. The gravitational effects of the deep density distributions that support the topography within 166.7 km of 

each grid intersection were computed according to the 

Airy-Heiskanen model of isostasy (Heiskanen and 

Moritz, 1967) using a 5- by bminute topographic-bathymetric data grid and model parameters as follows: 

topographic density, 2.67 g/cm8; crustal thickness at sea 

level, 30 km, and densit contrast across the base of the ? model crust, 0.35 glcm . Combined isostatic and topographic effects for the region from 166.7 km to the 

antipode of each grid intersection were obtained from the 

maps by Karki and others (1961). This model gravity field 

was subtracted from each Bouguer gravity grid value to 

yield a grid of isostatic residual gravity values; the 

resulting grid was contoured by computer and displayed 

in color-band intervals of 10 mGal to produce figure 9.2. 

Limitations on the use of this map stem both from 

uncertainties in the point data from which the grid was 

constructed and from characteristics generated by the 

gridding process. For onshore data, uncertainties in the 

point data values resulting from errors in observed 

gravity, elevation, terrain corrections, and isostatic reductions are estimated to be less than 2 to 3 mGal for 

most stations, possibly larger in areas of extreme topographic relief (Simpson and others, 1986). In offshore 

areas, the greatest uncertainty results from conversion 

of the original free-air gravity data to Bouguer gravity 

values, using the 6- by 5-minute average bathymetry. 

Where the sea-bottom topography is relatively gentle, 

this conversion probably results in uncertainties of about 

5 mGal, but in such areas as parts of the California 

Continental Borderland (south of lat 34' N.) and over the 

edge of the Continental Shelf, where water depths 

change rapidly, errors of several tens of milligals are 

possible. These conversion errors generally appear as 

high-amplitude, nearly circular anomalies with diameters 

of as much as 40 km. 
9. CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 241 

Although gravity coverage along most of the San 

Andreas fault system is quite dense when viewed at the 

scale of figure 9.2, sampling of these data on a 4-km grid 

means that anomalies with characteristic dimensions less 

than several times the grid spacing are not faithfully 

portrayed. Our isostatic residual gravity map (fig. 9.2) is 

sufficient for qualitative and quantitative interpretation 

at the scale shown, but for more detailed interpretations, 

especially quantitative modeling, the reader is referred 

to the original data sources, such as Oliver and others 

(1980), Roberts and others (1981), Snyder and others 

(1982), and the other reports cited throughout this 

chapter. 

MAGNETIC ANOMALY MAP 

A magnetic anomaly map of the region surrounding the 

San Andreas fault system is shown in figure 9.3. This 

map is based on the magnetic anomaly map of the 

Western United States by Bond and Zietz (1987), which 

was compiled from hundreds of magnetic surveys with 

widely differing flight heights, flightline spacings, and 

sensor types. 

In contrast to the lengthy series of reduction steps that 

were required to convert the gravity observations to the 

form shown in figure 9.2, very little was done to the 

observed magnetic data to prepare them for compilation. 

Although the original data were collected at many 

different heights, no analytic procedures were used to 

continue them to a common elevation. Instead, the 

various surveys were referenced to the International 

Geomagnetic Reference Field (IGRF) adjusted for the 

date of the survey and an arbitrary zero datum, and then 

combined manually by inspection. Long profiles of magnetic data collected under the National Uranium Re 

source Evaluation (NURE) program of the U.S. 

Department of Energy and by the U.S. Naval Oceanographic Office (NO01 served as guides for determining 

the zero datum for the various surveys. The resulting 

data are presented at a color contour interval of 100 nT 

(gammas) in figure 9.3. 

Our magnetic anomaly map (fig. 9.3) is the most 

complete compilation available for the San Andreas fault 

system and is useful for qualitatively determining the 

location, shape, and regional setting of large magnetic 

bodies. However, because of the compilation methods 

used to construct this map and because the contour 

interval is relatively coarse (100 nT), it will not, in 

general, be adequate for detailed qualitative or quantitative examination of individual anomalies. Where detailed information is required, the reader is referred to 

the original sources from which our map was compiled; a 

comprehensive listing of these sources is given by Bond 

and Zietz (1987). 

A particularly valuable source of aeromagnetic data 

over the San Andreas fault system is the profile data 

collected under the NURE program. In general, these 

data were collected along long profiles oriented east-west 

at a nominal height of 120 m above terrain and spaced 

about 5 km apart. The wide flightline spacing and low 

altitude of the survey lines preclude constructingrealiitic 

contour maps from these data in most places, but the long 

profiles are well suited for quantitative modeling. These 

data are available in the form of atlas folios or digital 

tapes for individual lo by 2Â quadrangles from the U.S. 

Department of Energy, Grand Junction, Colo. 

When interpretingkgnetic data, the inclination of the 

Earth's magnetic - field must be taken into account because the magnetization induced in the magnetic source 

rocks by this field will have a similar inclination. Along 

the San Andreas fault, this inclination ranges from 58' to 

64' downward toward magnetic north. Contoured magnetic anomalies over inductively magnetized or normally 

magnetized sources at these field inclinations will commonly display a dipole response, namely, magnetic lows 

associated with the north sides of magnetic highs. Inspection of our magnetic-anomaly map (fig. 9.3) indeed 

identifies numerous such magnetic lows on the north or 

northeast sides of major magnetic highs. In general, each 

magnetic low is located directly beyond the north or 

northeast contact of the causative magnetic mass. 

SOURCES OF GRAVITY AND 

MAGNETIC ANOMALIES 

Conspicuous features of the gravity field over the San 

Andrek fault system are linearhighsand lows that trend 

subparallel to the major faults in the system. Highs (210 

mGal) generally occur over exposed crystalline rocks of 

the Salinian block southwest of the San Andreas fault, 

over mafic granitic and metamorphic rocks of the Sierra 

Nevada and the Mojave Desert, and over Mesozoic and 

Tertiary layered rocks of the Franciscan assemblage, 

particularly in areas containing large amounts of mafic 

volcanic rocks (generally part of an ophiolite belt) or 

high-pressure metamorphic-mineral facies. Most of the 

deepest lows are caused by thick accumulations of 

low-density Cenozoic sedimentary rocks that fill tectonic 

basins adjacent to the faults and in the surrounding 

areas. Shallower lows occur over certain large serpenfinite bodies within the Franciscan assemblage, over 

felsic plutons in the granitic terranes of California, and 

over a young concealed granitic pluton associated with 

the Geysers geothermal area at lat 39' N., long 122%' 

W. (Chapman, 1976; Isherwood, 1976). 

Magnetic anomalies in the vicinity of the San Andreas 

fault system typically are caused by any one of three 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

Â¥f 8 *ve 'a liS ^ 4 ^. 4% ' 

\ , \ / \ /Â Â¥ / \ 

// 

I 

0 - 50 100 KILOMETERS 

EXPLANATION 

m 

ISOSTATIC RESIDUAL GRAVITY VALUES. IN N-"ALS 

FIGURE 9.2.-Isostatic residual gravity map of the San Andreas fault system. Contour interval, 5 mGaL Faults simplified from Jennings and 

others (19771, McCulloch (1987). and Vedder (1867). Fault mea and explanation in Bgure 9.4. 

different rock types. The strongest anomalies generally 

reflect tabular bodies of serpentiite associated with the 

Franciscan assemblage and may also reflect the ophiolitic 

rocks, especially serpentii, that locally lie above it. 

Mafic plutonic rocks, such as those exposed in the 

western Peninsular Ranges and along the west edge of 

the southern Sierra Nevada, can produce moderate to 

strong magnetic anomalies. Plutonic sources, not necessarily mafic only, probably account for most of the 

anomalies in the Salinian block, southwest of the San 

Andreas fault. Although younger volcanic rocks, in 

particular the mafic varieties, commonly are highly 

magnetic, such rocks do not cause significant magnetic 

features near the San Andreas fault as shown on our 

magnetic anomaly map (fig. 9.3) because magnetic voleaNC rocks are volumetrically unimportant at the scale of 

this map. 

In most areas, sedimentary rocks are considered 

nonmagnetic because they fail to cause aeromagnetic 

anomalies. Along the San Andreas fault system, however, several sedimentary-rock units cause magnetic anomalies as large as 160 nT. These units include rocks of 

Mesozoic and Tertiary age; other units composed primarfly of detrital serpentinite also produce anomalies of this 
9. CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 

GEOMETRY OF FAULTS IN 

THE SAN ANDREAS SYSTEM 

magnitude. None of these sedimentary units are areally 

large enough to produce magnetic anomalies visible at 

the scale of our magnetic anomaly map (fig. 9.3). 

Near the north end of the San Andreas fault, several 

magnetic anomalies project landward from the linear 

pattern of anomalies that characterizes the oceanic crust. 

These anomalies reflect remanent magnetization in the 

oceanic crust; the source rocks are primarily basaltic 

volcanic rocks. 

PIAN WBW 

(Blakely and Simpson, 1986). In general, the fault is 

expected to be situated at or near the steepest gradient 

of the anomaly. These sites are particularly helpful in 

areas where the fault trace or zone is concealed by young 

sedimentary deposits or by the Pacific Ocean. In addition, we have found that some of these data are useful in 

identifying the main strand of the fault zone where the 

presently active fault trace may not, in fact, be the 

original plate boundary. Some areas where the potentialfield data define the locations of faults are shown on 

figure 9.4 and are discussed below. 

Although the location of the San Andreas fault between Point Arena and Cape Mendocino is concealed by 

the Pacific Ocean, the aemmagnetic data show a linear - - - . . . 1 magnetic anomaly, striking northwest within the Pacific 

The usefulness of ootential-field data alone the San date. that is inferred to be obliouelv cut off bv the fault 

Andreas fault system is maximized where rock masses 

with differing physical properties are juxtaposed. Under 

these conditions, geophysical anomalies arise from which 

the location and attitude of the fault may be calculated 

about 20 km northwest of point Arena. art her north, 

the fault trace just south of Cape Mendocino has proved 

particularly difficult to locate because it may be too close 

to shore to be resolved by marine geophysical surveys. A 
THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

50 100 KILOMETERS I OMAGNETIC ANOMALY, IN NANOTESLAS 

FIGURE 9.3.-Magnetic map of the San Andreas fault system (from Bond and Ziilz, 1987). Contour interval, 100 nT. Fwlh simplified from 

JemiingB and others (19771, McCuUoch (1987), and Vedder (1987). Same symbols is in Sam 9.4. 

recent aeromagnetic map of this problematic area has, 

indeed, displayed a magnetic boundary trending dose to 

and along the shore, thus representing the likely location 

of the San Andrea fault (Griscom, 1980~1). 

Aeromagnetic surveys over the Pacific Ocean at the 

entrance to the San Francisco Bay (Brabb and Hanna, 

1981) show that the offshore extension of the Hiaroitos 

fault (an inactive fault branching westward from the San 

Andreas fault) is cut off by the offshore northward 

extension of the San Gregorio fault. The San Gregorio 

fault can be traced northward by using a detailed 

aeromagnetic map to the point where it intersects the 

San Andreas fault at Bolinas Lagoon, about 20 km 

northwest of the bay mouth (see McCuUoch, 1987, fig. 

16). 

From San Francisco southward to lat 3S015' N., the 

detailed gravity and magnetic data indicate that, in 

general, the westernmost strand of the main San Andreas fault zone is the major plate boundary. The layered 

Franciscan assemblage to the east may be less competent 

than the granitic basement of the Salinian block to the 

west, and new strands may be more likely to appear in 

the less competent rocks. An exception to this generalization is found at lat 36Â N., where a thin fault sliver of 
9. CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 

hornblende-quartz gabbro occurs at Gold Hill (Ross, 

1970) that has been used to estimate offset on the San 

Andreas fault. The magnetic anomaly associated with 

this gabbro body indicates that it is at most 10 km long by 

2 km wide (U.S. Geological Survey, 1987). 

Farther south along the San Andreas fault, a linear 

magnetic high extends along the fault approximately 

between long 116Â and 118" W. (fig. 9.3). On the basis of 

local model studies of this anomaly, Simpson and others 

(in press) show that this feature probably reflects the 

edge of an extensive block of magnetic rocks on the 

northeast side of the San Andreas fault, where the 

magnetic material is Precambrian igneous and metamorphic rocks, as well as Mesozoic plutonic rocks. Using 

detailed magnetic data (U.S. Geological Survey, 1979), 

the south border or magnetic boundary of this magnetic 

block (fig. 9.4) can be traced from west to east along a 

series of fault segments; from long 117'16' W., the 

boundary follows the southern fault trace to long 116"15' 

W., then crosses over to the northern trace along the 

short, east-west-trending fault segment, and finally continues eastward along the northern trace. These faults 

thus may represent the original fault boundary (now 

somewhat kinked) between the two plates. The geologic 

observation that rocks on the north side of these fault 

segments are native to the San Bernardino Mountains 

and contrast with compositionally different rocks on the 

south side (Matti and others, 1986) agrees with the 

magnetic interpretation. The magnetic boundary continues southeastward along the San Andreas fault in Coachella Valley to long 116W W. A possible farther 

continuation of this linear magnetic high extends southeastward at a lower amplitude and diverges eastward 

from the present San Andreas fault, generally following 

and lying northeast of the Clemens Wells fault, a possible 

earlier strand of the San Andreas fault. 

In Coachella Valley, the San Andreas fault (North 

Branch or Coachella segment) is situated along the 

northeast side of a substantial linear gravity low caused 

by at least 4.7 km of lowdensity sedimentary rocks 

9. CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 

(Biehler, 1964) filling the valley. Gradient studies on the 

relatively detailed gravity datain this valley by one of us 

(Griscom) identify numerous fault strands, including the 

northern and southern branches of the San Andreas fault 

(the latter, the Banning fault), as well as several possible 

fault segments on the southwest side of the valley. 

The Garlock fault at long 118'W. changes direction and 

forms a zone as much as 8 km wide. Models of both the 

magnetic and gravity fields calculated normal to the fault 

indicate that here the main lithologic boundary is the 

most northerly fault (fig. 9.4); the granitic rocks farther 

north are more magnetic and less dense than those to the 

south. 

ATTITUDE 

At several sites along the San Andreas fault, the 

juxtaposition of large masses with contrasting densities 

and (or) magnetizations causes characteristic potentialfield anomalies that reflect the attitude of the fault. 

Information on the dip of the fault most commonly is 

obtained through quantitative modeling of these anomalies, but in some cases, the anomalies are so diagnostic 

that qualitative interpretations suffice to indicate the 

direction and approximate attitude of the fault plane. 

The results of such interpretations at 16 sites primarily 

along the main trace of the San Andreas fault are shown 

in figure 9.4. The dip and depth-extent of the density or 

magnetization interfaces that are assumed to define the 

fault plane at these sites are somewhat uncertain because 

of the inherent ambiguity of gravity and magnetic 

interpretations, particularly those based on magnetic 

data, because rock magnetizations can have anomalous 

directions associated with their remanent components 

and because magnetic susceptibilities seldom are known 

with sufficient accuracy to serve as effective independent 

constraints. Thus, where the magnetic anomalies can be 

compatible with a vertical fault, we show the dip as 

vertical. The reader should he aware, however, that a 
248 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

dipping interface extending to shallower depth would 

also be compatible with the data in some places. 

Although major strike-slip faults probably are vertical 

over much of their reach, some have inferred dips of less 

than 90Â° as indicated by many of the attitudes shown in 

figure 9.4. Just north of Point Arena, a buried magnetic 

body truncated on the east by the San Andreas fault has 

an east boundary that dips east beneath the trace of the 

fault (fig. 9.5A); its precise dip is uncertain but probably 

falls in the range 30Â°-50' To the south, near the junction 

of the San Andreas and Calaveras faults, gravity modeling (Pavoni, 1973) suggests that the fault dips 70" SW. to 

a depth of about 6 km. A detailed study of seismicity 

along this section of the fault (Spieth, 1981) shows that 

hypocenters define a plane dipping 70Â SW., thus strongly supporting the interpretation by Pavoni (1973). Robbins (1982) also found a southwest-dipping density 

interface at this site but argued that the fault plane was 

vertical, on the basis of a magnetic anomaly that he 

believed reflected a magnetic body, extending from 3- to 

5-km depth, with a southwest edge directly beneath the 

surface trace of the fault. More recent, detailed magnetic 

measurements indicate that this magnetic body is much 

shallower (probably cropping out) than modeled by 

Robbins (1982) and thus weaken the argument for a 

vertical dip. 

Near the intersection of the San Andreas and Garlock 

faults, gravity modeling by Andrew Griscom and K.G. 

Freeman (Griscom and Oliver, 1980) suggests that the 

fault dips 55Â SW. to a depth of 6 km and thence vertically 

to a depth of at least 10 km (fig. 9.5B). About 60 km 

farther southeast, gravity data also indicate a southwesterly dip for the fault, but the angle of dip (30'40') is 

uncertain, owing to difficulty in interpreting the complex 

gravity field that results from large lateral density 

variations in the region southwest of the fault. Farther 

southeast, where the San Andreas fault splits into 

numerous branches (long 116'00' W.), gravity data on 

two branches indicate that both faults dipnortheast, with 

Precambrian crystalline rocks in the upper plate overlying young sedimentary rocks and alluvium. The gravity 

models suggest dips of 15'-25' NE. to depths of 1.5 to 2.5 

km but do not resolve the fault attitude at greater depth 

FIGURE 9.5.-Magnetic and gravity models across the San Andreas 

lilt. m, magnetization; p, density. A, Magnetic model just north of 

Point Arena (lone 123-40' W.l. B. Gravitv model near iunction of the 

San ~ndreasand~arlock faultsflone 1i9Â¡07 N.). ~.~esozoic and

the Great Valley. C, Gravity model across southern branch at long 

11ff40' W. A, Mesozoic and Precambrian crystalline basement; B, 

Tertiary and Quaternary sedimentary rocks of Coachella Valley. 

fig. 9.50. Geologic mapping, which shows part of the 

southern branch of the San Andreas fault as a northlipping thrust fault (Matti and others, 1985), and a study 

)f recent earthquakes in this area, which yielded fault- 

)lane solutions of predominantly oblique-slip motion and 

!S Observed /'\ 

9 , Â¥"Â¥Â 

SANANDREASFAULT 

Observation level 

rn=O A/". 

rn-0.43 Nm 

0 

A DISTANCE. INKILOMETERS 

c N 

SAY ANDREAS FAULT 

S 5 

,a 

B Â¥CE IN KILOMETERS 

S N 

I 

ANDREAS 

FAULT 

0 

l 0 ,o 20 

c DISTANCE. IN KILOMETERS 
9. CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 249 

including low-angle thrust solutions dipping 30Â N. 

(Nicholson and others, 19861, both support the gravity 

interpretation of northeast-dipping faults in this area, 

The inferred fault attitudes shown in figure 9.4 suggest 

a relation between attitude and plan-view geometry. 

Faults tend to be vertical except where they undergo 

abrupt changes of strike. The sinistral bends in the San 

Andreas fault near its junction with the Calaveras fault 

and in the Big Bend region southeast of its junction with 

the Garlock fault create regions likely to be subject to 

compression due to relative southward movement of the 

North American plate with respect to the Pacific plate. 

The dipping fault planes in these regions may reflect a 

thrust component of fault movement that accommodated 

the compression. Similarly, the region around the broad 

dextral bend in the San Andreas fault north of Point 

Arena may have a component of extension parallel to the 

direction of relative plate motion, and the low- angle 

eastward dip of fault plane there may reflect accommodation of the extension by low-angle normal faulting. 

The number of examples on which the above speculations are based is quite limited, and further detailed 

investigations at critical sites along the San Andreas fault 

system are needed to test the relation between fault 

attitude, change of strike, and relative plate-motion 

direction. 

FAULT-ZONE CHARACTERISTICS 

Millions of years of strike-slip movement along faults of 

the San Andreas system have produced, in many places, 

a narrow fault zone in which physical properties differ 

from those in the surrounding rock masses. These 

differences are due to the presence within the fault zone 

of fractured or pulverized rock, exotic rock slivers that 

have been transported along the fault from other places, 

and such mobile materials as fluids and sernc-ntinite that A ~~ ~ ~~~~ 

have migrated along the fault zone. A few investigators 

have used gravity and magnetic data to study the 

properties of this zone. 

Although Stierman (1984) and Wang and others (1986) 

sought to explain gravity lows along the fault as the 

result of a substantial increase in porosity by fracturing, 

gravity lows not directly associated with basins filled by 

Cenozoic sedimentary rocks along the faults are very 

uncommon. These rare lows amo&t, with few exceptions 

(such as the 10-12-mGal low studied bv Stiman. . I9ML - -,, 

to amplitudes of only a few milligals. Feng and McEvilly 

(1983) and Trehu and Wheeler (1987) inferred from 

seismic data that zones of low seismic velocity 5 to 10 km 

wide and more than 10 km deep are associated with the 

San Andreas fault zone and, presumably, with fractured 

rocks. The low-velocity zone of Trehu and Wheeler 

(1987), however, has no associated gravity low, even 

though calculations by Andrew Griscom indicate that this 

zone might be expected to produce a gravity anomaly of 

about -25 mGaI and more than 10 wide, using the 

standard velocity-density relations of Hill (1978). An 

explanation for this unexpected result can be found in the 

borehole gravity and seismic-velocity results (Schmoker, 

1977; Stiennan and Kovach, 1979) from a 600-mdeep 

borehole in diorite located 1.2 km from the San Andreas 

fault. For the lower half of this borehole, the seismic 

velocity averages only 3.1 kmls (although saturated core 

samples measured 6.6 kmls in the laboratory), and the 

average computed rock densities are as follows: bulk 

density from cores, 2.72 g/cm3; borehole density from 

gravity measurements, 2.60 g/cm3; and computed density 

from borehole velocities (density-velocity relations of 

Hill, 1978), 2.36 g/cm3. Correcting for a nearby lowdensity sedimentary section that causes a gravity gradient along the hole raises the borehole density (from 

gravity measurements) closer to the bulk density. The 

results described above indicate that macrofractures can 

cause large decreases in seismic velocity but much 

smaller decreases in density than those predicted from 

standard velocity-density relations. 

Allen (1968) pointed out a possible relation between the 

style of fault movement and the presence of serpentinite 

within fault zones of the San Andreas, Calaveras, and 

Hayward faults. He noted that serpentine is common 

within the fault zone along the creeping section of the San 

Andreas fault between Holliyterand Cholame, whereas it 

is absent along the locked segments to the north and 

south. Irwin and Barnes (1975) noted the same relation 

between serpentinite and fault creep and discussed the 

possible role of metamorphic fluids on the seismic behavior of fault segments. Hanna and others (1972) studied 

aeromagnetic data along the San Andreas fault between 

San Francisco and San Bemardino and found that the 

creeping segment of the fault is characterized by broad 

aeromagnetic anomalies, which they interpreted as reflecting large concealed masses of serpentinite. Linear 

magnetic anomalies that most likely reflect serpentinite 

also are present along the creeping section of the 

Hayward fault east of the San Francisco Bay (fig. 9.3). 

These magnetic data support the speculation that appreciable amounts of serpentinite contained within a fault 

zone can influence the style of movement on the fault. 

OFFSETS OF ANOMALIES 

Strike-slip movement on the faults of the San Andreas 

system has produced offsets in formerly continuous 

geophysical anomalies. As might be expected, on those 

faults where the geologic offset is at most a few tens of 

kilometers, it is generally easy to identify corresponding 

magnetic or gravity features that are offset by similar 
250 THE SAN ANDREAS FAU 

distances. Examples of such faults are the Elsinore fault 

and the rectilinear system of minor strike-slip faults in 

the Mojave Desert block northeast of the San Andreas 

fault. In figure 9.4, the two or more piercing points of an 

offset geophysical anomaly are labeled with the same 

letter, and the specific points being described are designated with subscripts, numbered consecutively from 

northeast to southwest across the fault system. 

Offset along the San Andreas fault system in southern 

California is believed to he approximately 300 km in a 

right-lateral sense, based on offset of the Pelona-Orocopia schist belts, together with associated characteristic 

Precambrian and Triassic rock assemblages of the thrust 

plate overlying the schist belts (Crowell, 1962; Clarke 

and Nilsen, 1973; summarized in Hamilton, 1978). Isostatic gravity highs are associated with the Orocopia 

Schist (point A, at lat 3335' N., northeast side of fault), 

with the Pelona Schist of the Sierra Pelona/Soledad area 

(points A2, A3 at lat 34%' N., between the San Andreas 

and San Gabriel faults), and adjacent to the south side of 

the Pelona Schist of the TejonIGarlock area (point A, at 

lat 34"50' N., south of the San Audreas fault and west of 

the San Gabriel fault). Point A, is not well determined. 

The offsets of the gravity highs are, respectively, 240 km 

along the San Andreas between the first two highs (A, 

and A2) and 60 km along the San Gabriel between the 

second two highs (A3 and A,), for a total of 300 km along 

the San Andreas fault system, in agreement with Crowell 

(1962). The source of the gravity highs is not obvious and 

may not be any of the rocks exposed at the surface 

(Griscom, 1980b), both because the density of the schist 

coring the antiforms is probably similar to or slightly 

lower than that of the surrounding Precambrian crystalline rocks and because other large areas of Pelona/ 

Orocopia schist do not display associated gravity highs. 

The schist is marine in origin, predominantly metagraywacke of low metamorphic grade (Haxel and Dillon, 

1978), and may be underlain by subducted oceanic crust. 

The gravity highs may indicate relatively uplifted oceanic 

crust beneath these specific antifonnal exposures of the 

schist, or else the proportion of greenstone interbedded 

with schist may increase here with depth. 

As mentioned above in the subsection entitled "Plan 

View," a linear magnetic high that extends along the San 

Andreas fault from long 116' to 118- W., a distance of 

about 200 km, indicates that a large area north of the fault 

in this region is composed of magnetic rocks, predominantly Mesozoic granitic plutons; the northwest limit of 

this magnetic area (J,) is shown in figure 9.4. A similar 

large area of magnetic basement, also predominantly 

Mesozoic granitic rocks, that extends along the southwest side of the San Andreas fault is displaced from the 

former area right-laterally approximately 300 km; the 

northwest limit of this correlative area (Jp) is also shown 

r SYSTEM, CALIFORNIA 

in figure 9.4. This second area of magnetic rocks does not 

produce a significant magnetic high directly at the fault 

because the fault is on the northeast side of the magnetic 

mass and a magnetic low should occur for this geometry. 

Several significant geophysical anomalies are found 

along the central section of the San Andreas fault north 

of its junction with the Garlwk fault. A pronounced 

gravity high is located on the northeast side of the fault 

at lat 34O55' N., where the southern "tail" of the Sierra 

Nevada is exposed. The associated rocks are hornblendequartz gabbro and anorthositic gabbro (Ross, 1970,1984) 

that also produce a substantial aeromagnetic high (point 

B,, fig. 9.4). Similar rocks (ROBS, 1970) are found within 

the San Andreas fault zone at Gold Hill (point B2 at lat 

35"50' N., too small to show at this scale) and at Logan 

(point Bg at lat 36'52' N.), where magnetic anomalies 

(U.S. Department of Energy, 1981; U.S. Geological 

Survey, 1987) indicate that the gabbro bodies are thin 

slivers within the fault zone. The Logan outcrops are 

offset about 290 km from the gabbro of the Sierran "tail." 

A major northwest-trending magnetic anomaly extends 

northwestward of Logan near the coast (from point B4 at 

lat 37W8' N.). The source rocksfor this magneticanomaly 

are interpreted to be gahhro, similar to that exposed near 

Logan (Hanna and others, 1972a), because the anomaly 

requires a source body several kilometers thick. These 

corresponding offset geophysical anomalies support the 

geologic correlations implying about 300 to 320 km of 

offset. 

The additional 100 km of granitic rocks extending 

northward from Logan to Montara Mountain (lat 3735' 

N.) along the southwest side of the San Andreas fault 

does not have any correlative rocks exposed on the 

northeast side of the fault north of the gabbro of Sierran 

"tail," but the concealed crystalline basement rocks 

beneath the sedimentary rocks of the Great Valley may 

be correlative. Indeed, recent work on the Tertiary 

sedimentary rocks that overlie this additional 100 km of 

granitic rocks on the San Francisco peninsula suggests a 

lithologic and paleogeographic correlation with similar 

sedimentary rocks of the Great Valley (San Joaquin 

Basin) that are relatively offset 320 to 330 km to the 

southeast (see fig. 3.4; Stanley, 1987). 

Movement on the San Andreas fault north of San 

Francisco (Griscom and Jachens, 1989) is complicated by 

right-lateral displacement added by the presently active 

San Gregorio-Hosgri fault, which intersects the San 

Andreas fault at San Francisco and provides an additional 115 km (Graham and Dickinson, 1978) or 150 km (Clark 

and others, 1984; Ross, 1984) of offset. The total offset on 

the San Andreas fault system here is farther complicated 

by movement on branch faults to the east (Calaveras and 

Hayward fault systems of unknown offset) and, more 

importantly, by past movement along the Pilarcitos fault, 
9. CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 251 

the presently inactive fault strand branching westwari 

from the San Andreas fault on the San Francisco peninsula. Probably most of the 300 to 320 km of displacement 

on the San Andreas fault has taken place along this 

strand because the presently active strand of the Sar 

Andreas fault that lies directly east of the Pilarcitos fault 

demonstrates only about 26 km of offset of a characteristic limestone belt within the Permanente terrane of the 

Franciscan assemblage (Bailey and others, 1964, p. 69 

M.C. Blake, Jr., oral commun., 1987). The Pilarcitos fault 

(now truncated to the northwest by the San Gregoric 

fault) may have its former extension on the ocean side 01 

the San Gregorio-San Andreas fault at about lat 3830' N. 

(see fig. 9.4), as proposed by Graham and Dickinson 

(1978). This proposed extension may have granitic rock? 

on the southwest side (more than the additional 100 km 

already discussed) that have no correlatives northeast ot 

the San Andreas fault, unless the total offset on the Sar 

Andreas system substantially exceeds 300 km or unless 

granitic rocks underlie thrust blocks of Franciscan assemblage near the south end of the Great Valley (see 

preceding paragraph). There may be other, unidentified 

faults within the Salinian block that allow for this 

additional offset. 

Offsets of geophysical anomalies along the San Gregorio-Hosgri fault support a total right-lateral movement oi 

about 100 to 130 km that has been added to the total offset 

on the San Andreas fault systemnorth ofits junction with 

the Sail Gregorio fault. An offset gravity high (Silver, 

1974) is located on the northeast side near Point Sur 

(point Ci at lat 36Â¡30 N.) and on the southwest side at 

Ano Nuevo (point C2 at lat 3T15' N.), with an offset of 

105 to 130 km as remeasured by Graham and Dickinson 

(1978). We prefer an offset of 105 km (max 115 km) 

because any larger displacement will place the offset 

extension of the Pilarcitos fault on land north of lat 3830' 

N., where no such fault is known. 

Displacements along the San Andreas fault north of lat 

3830' N. have proved difficult to measure, both because 

the rocks exposed southwest of the fault near Point 

Arena have no obvious correlatives on the opposite side 

of the fault and because most of the fault trace is 

concealed beneath the Pacific Ocean (Griscom and 

Jachens, 1989). The rocks cropping out southwest of the 

fault near Point Arena are Upper Cretaceous and Tertiary marine sedimentary rocks, together with some older 

spilitic volcanic rocks that may be part of the Franciscan 

assemblage (Wentworth, 1968). Little basement information from rock samples is available in the shelf areas west 

of the San Andreas fault between Point Arena and the 

Mendocino fault. An important well 20 km west of Point 

Arena (fig. 9.4) recovered quartz-mica schist and slate 

basement cuttings (Hoskins and Griffiths, 1971) at a 

depth of about 1.43 km. This description resembles rocks 

either from the eastern metamorphosed Franciscan assemblage south of San Francisco or from roof pendants in 

the Salinian block, implying that a major strike-slip fault 

is located between the well and Point Arena. The 

proposed Pilarcitos fault extension is thus interpreted to 

lie here between the well and the coastline on a major 

fault shown by McCulloch (1987). Location of this proposed Pilarcitos fault extension farther northwest than 

Point Arena is uncertain, although the fault presumably 

continues to the former triple junction. McCulloch (1986; 

1987, fig. 2b) described a boundary, termed the "Navarro 

discontinuity," trending east-west from the Point Arena 

area to the lower continental slope, on the basis of 

regional differences in magnetic pattern and physiography; this boundary may be the fault extension or an 

earlier strike-slip fault of this system. Griscom and 

Jachens (1989) also hypothesized a more northwestward 

extension, approximately colinear with the fault segment 

south of Point Arena, following a fault trace interpreted 

from seismic-reflection profiles (McCulloch, 1987, fig. 

14). 

Distinctive gravity and magnetic anomalies characterize the poorly known shelf area lying north of Point Arena 

and between the San Andreas fault and the proposed 

Pilarcitos extension (figs. 9.2, 9.3). The sources of these 

anomalies lie in the basement, with their upper surfaces 

at the basement interface below Tertiary sedimentary 

rocks, according to geophysical models and basementdepth calculations. A major gravity high (+20 mGal) is 

located near lat 40Â°N (point D,, fig. 9.4). We believe that 

the highdensity basement rocks which cause this high 

extend southward along the west side of the San Andreas 

fault at least as far as at Point Arena (E,, fig. 9.4), even 

though the gravity values on the map fall below 0 mGal 

along the southern part of this reach. The basement along 

the postulated southern part of the high is mantled by 1 

to 3 km of Tertiary sedimentary rocks (Hoskins and 

Griffiths, 1971), which probably cause gravity lows (-15 

to -30 mGal) that here mask the gravity high caused by 

the basement. Two magnetic anomalies on the shelf are 

truncated by the San Andreas fault at point F2 and at a 

place a few kilometers south of point Go (fig. 9.4), which 

is located where the steepest gradient on the northeast 

side of the second anomaly is truncated by the fault (see 

McCulloch, 1987, fig. 17). 

Our search for geophysical anomalies or features 

matching points Do, E2, F2, and G2 on the opposite 

(northeast) side of the San Andreas fault (see Griscom 

and Jachens, 1989) began with the observation that 

gravity highs are not characteristic of much of the 

Franciscan assemblage and are observed only extending 

along the San Andreas fault between approximately lat 

87" and 38" N. (fig. 9.2). We have selected points D, and 

E, (fig. 9.4) as the approximate limits of thegravity highs 
252 THE SAN ANDREAS FAl 

on the northeast side and propose to correlate these 

points and their connecting strip of high gravity with the 

corresponding points Dz and E, and associated gravity 

high discussed in the previous paragraph. The positions 

of these points (D and E) along the fault vary in reliability 

but are probably no more accurate than Â±2 km, point Ez 

is the most uncertain. The total offset of the gravity high 

by the San Andreas fault is thus about 260240 km. We 

have used the gravity results to explore our magneticanomaly map (fig. 9.3) for additional correlations. Only 

one correlation was found within an offset range of 

200-300 km. We suggest that point F,, marking the end 

of a truncated magnetic high passing through San Francisco, correlates with point F, and that point GI, the 

truncated end of a magnetic gradient more than 50 km 

long, correlates with the other truncated gradient at 

point Gz. The locations of points F,, Fz, and Go along the 

fault are accurate to within about Â± km. Point G, is 

located a few kilometers too far to the southeast because 

a short northwestward extension of the feature was 

recently cut off by the young segment of the San Andreas 

fault in the San Francisco peninsula area and now lies 

between the San Andreas and Pilarcitos faults. The offset 

of points F, and F, is 260 km; the offset of points GI and 

Gy is 263 km. The magnetic anomalies truncated at points 

F, and G, are associated with northwestrstriking belts of 

mafic and ultramafic rocks within the Franciscan assemblage and are best shown on the more detailed maps by 

Brabb and Hanna (1981) and Griscom and Jachens (1989). 

We conclude that the total offset of the pairs of 

corresponding magnetic features is approximately 

250Â±1 km. Of this offset, about 105 km is attributable to 

the San Gregorio-Hosgri fault, leaving only 145 km for 

the San Andreas fault south of its junction with the San 

Gregorio fault. Because the total San Andreas offset 

south of the San Francisco peninsula is considered to be 

much larger, namely, about 300 km, we suggest that the 

missing 155 km is predominantly accounted for by former 

movement on the Pilarcitos fault and its proposed northwestward extension, which is thought to intersect the 

San Andreas at about lat 38Â¡30 N., as described above. 

This early Pilarcitos fault was thus formerly the main 

strand of an earlier San Andreas fault system that lay to 

the west of both magnetic features F and G (that is, 

before they were offset by faulting). Note that this 

interpreted fault-movement history and the subsequent 

plate-tectonic analysis all depend on the correctness ot 

the correlation between the pairs of offset magnetic and 

gravity anomalies on the San Francisco peninsula and 

northwest of Point Arena. The magnetic and gravity 

anomalies northwest of Point Arena and west of the San 

Andreas fault are such conspicuous features and sc 

obviously truncated by the San Andreas fault that we 

would expect to find their counterparts somewhere or 

' SYSTEM, CALIFORNIA 

;he opposite side of the fault. Although we can find no 

alternative correlations for these anomalies other than 

;hose indicated in figure 9.4, we are aware that they may 

not correlate with any anomalies on the opposite side of 

die fault, although we consider this noncorrelation to be 

unlikely. 

Additional information on offset along the proposed 

Pilarcitos fault extension is provided by interpretation of 

two strong magnetic anomalies on the northeast side of 

the San Andreas fault in central California (Hi at lat 

SÂ¡80'-35'40 N. and lat 36'00'46'15' N., respectively). 

The source bodies for both anomalies appear to be 

truncated by the fault, and interpretations of the gravity 

and magnetic fields over both bodies suggest that they 

are composed of serpentinite (Hanna and others, 1972; 

Griscom and Jachens, in press). The most likely candidates for corresponding magnetic features on the southwest side of the fault system are the poorly defined 

anomalies at points Hz and In west of Point Arena. The 

magnetic field is poorly known in this area, and so 

anomaly locations and shapes may not be accurate, but 

the offset is approximately 435 km from points HI and I,. 

This distance can be obtained by summing an assumed 

320 km for offset on the San Andreas fault south of San 

Francisco plus 115 km offset on the San Gregorio-Hosgri 

fault. The location of point Iz supports the Navarre 

discontinuity as a possible continuation of the proposed 

Pilarcitos fault extension, or some earlier continuation. 

We suggest that the large magnetic-high area bounded 

by the 500-nT contour and located 25 km south of point 1, 

(fig. 9.3) may represent a southerly extension of anomaly 

I, which north of point I, extends for 400 km along the 

east side of the Coast Ranges. There appear to be no 

satisfactory alternative anomalies for correlation with 

points HI and I, along the southwest side of the present 

San Andreas fault near Point Arena. 

IMPLICATIONS FOR PLATE TECTONICS 

In the previous sections, we have discussed how 

potential-field data provide information on the threedimensional configuration of the San Andreas fault and 

on the various offsets along member faults of the San 

Andreas system in relation to plate tectonics. Here, we 

interpret the potential-field expression of the two ends of 

the San Andreas fault system in relation to plate tectonics and lithospheric thickness. 

MENDOCINO TRIPLE JUNCTION 

At the north end of the San Andreas fault off Cape 

Mendocino, three lithospheric plates (the Juan de Fuca, 
9. CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 253 

Pacific, and North American) meet at the Mendocino 

triple junction, where a trench meets two transform 

faults, the Sail Andreas and Mendocino faults. Along this 

trench to the north, the Juan de Fuca plate is subducting 

eastward beneath the North American plate. The geometry of this subducted plate has important implications 

(Jachens and Griscom, 1983) for an understanding of the 

Mendocino triple junction and its effects on the tectonics 

of California. During approximately the past 29 Ma, this 

triple junction has been migrating relatively northwestward along the coast of California from a latitude near 

Los Angeles (see fig. 3.11; Atwater, 1970; Atwater and 

Molnar, 1973). As this incipient San Andreas transform 

fault lengthened over time, eastward subduction continued to the north of the migrating triple junction. 

During 29-23 Ma, the major fault of the San Andreas 

system was probably situated near the base of the 

continental slope, where an accreted wedge of Miocene(?) 

sedimentary rocks (McCulloch, 1987) accumulated between lat 35O and 40Â N., presumably because of oblique 

subduction from transpressive forces between the plates. 

This now-inactive fault forms a contact between oceanic 

and continental crusts (fig. 9.4) that have major differences in magnetic properties (fig. 9.3). The oceanic crust 

displays the typical oceanic lineated or striped magnetic 

pattern striking north-south and northeast, with interruptions striking east-west or southeast that are caused 

by transform faults. The continental crust adjacent to 

this inactive fault is magnetically rather smooth and 

featureless. The magnetic boundary between oceanic and 

continental crust west of the San Andreas fault is very 

abrupt in comparison with active subduction zones (compare the magnetic expression of the Cascadia subduction 

zone off Oregon in Bond and Zietz, 1987); the oceanic 

stripes terminate at the base of the continental slope, 

even though reflection profiles show oceanic crust continuing farther east beneath the slope (McCulloch, 1987). 

The low convergence rate of oblique subduction and the 

time available since the fault became inactive may have 

allowed the concealed or subducted oceanic crust to heat 

up sufficiently beneath the continental margin to destroy 

the remanent magnetization that causes the stripes. 

During early Miocene time (23 Ma), the motion along 

the transform must have been essentially strike slip and 

was substantially transferred to the present San Andreas 

fault system in central California. Without subduction 

east of the elongating transform, an ever-enlarging 

triangular hole or window (Dickinson and Snyder, 1979) 

developed in the slab of lithosphere subducted beneath 

the continent. This window model is also applicable to the 

time interval (29-23 Ma) but needs modification to 

include effects of transpression along the earlier San 

Andreas fault. The north boundary of this window is the 

subducted south edge of the Juan de Fuca plate, and hot 

~pwelling asthenospheric material presumably occupies 

;he window. The south edge of the Juan de Fucaplate lies 

oeneath the North American plate at the shore about 20 

km south of Cape Mendocino and can be identified by an 

iast-west magnetic anomaly (Griscom, 1980a). as well as 

by the distribution of seismicity (Hntchings and others, 

1981). This position coincides with a steep gravity graiient (here called the Cape Mendocino gravity anomaly) 

;hat slopes downwardinto alarge gravity low (-50mGal) 

;o the north and east. The spatial coincidence between 

Ae position of the Cape Mendocino gravity anomaly at 

;he coast and the place where the south edge of the Juan 

ie Fuca plate passes beneath the coastline strongly 

suggests that this gravity anomaly reflects the south 

idge of the subductedplate (fig. 9.4). At least three other 

characteristics (Jachens and Griscom, 1983) of the anomaly support this interpretation. (1) The southeastward 

;rend of the gravity anomaly and then its change to 

easterly are consistent with the directions of present and 

past relative motions between the Juan de Fuca and 

Pacific plates (Nishimura and others, 1981; Wilson, 1986). 

:2) The gravity anomaly broadens and is less steep 

;owad the southeast, suggesting that its source progressively deepens in this direction; calculated depths along 

Â± anomaly to the end of the southeast-trending segment define a line plunging approximately 9' SE. with a 

iepth of only 6 km at the coastline corresponding well to 

;he 8-km depth estimated from aeromagnetic data 

[Griscom, 1980a). (3) A cross section across this anomaly, 

using the above depths together with reasonable densities and thicknesses for the subducted Juan de Fuca plate 

aid the asthenospheric window fill to the southwest, 

produces a calculated gravity model (Jachens and 

Griscom, 1983) in good agreement with the observed 

yavity field. We draw the following conclusions from the 

gravity data (Jachens and Griscom, 1983). 

1. Above the south edge of the Juan de Fuca plate, the 

North American plate must have the shape of a thin lip 

that gradually thickens eastward, attaining a thickness of possibly only about 30 km at the Coast Range 

fault; this fault marks the east limit of the Franciscan 

assemblage about 130 km inland from Cape Mendocino 

(see chap. 3). Just south of the Juan de Fuca plate, 

asthenospheric material that filled the slab window 

should lie beneath the North American plate at a 

depth comparable to that of the upper surface of the 

Juan de Fuca plate. Because the North American 

plate has been moving relatively southward across 

this boundary for many millions of years, the top of the 

asthenosphere probably is shallow beneath much of 

the Coast Ranges in central California, and the thin 

west lip of the North American plate may be decoupled from much of the mantle, although some underplating by mantle material is likely. 
254 THE SAN ANDREAS PA1 

2. For reasons similar to conclusion 1, the lithosphere 01 

southern California near the San Andreas fault systeir 

is thin and may be decoupled from much of the mantle. 

3. Relatively thin, decoupled lithosphere may explain 

why deformation along the boundary between the 

Pacific and North American plates takes place over a 

zone 50 to 100 km wide rather than being restricted tc 

the San Andreas fault, and why the plate boundary 

has been able to migrate eastward from the base of the 

continental slope to its present position at the San 

Andreas fault. It may also explain both why certain 

structural blocks southwest of the fault in southern 

California have been able to rotate clockwise by as 

much as 7O0-90" during and after the Miocene (Luyendyk and others, 1985; Hornafius and others, 1986) 

and how extensional basins formed between these 

blocks. Furthermore, it can help explain why the 

seismicity of the San Andreas fault generally does not 

extend below 12-km depth. 

4. Thin, relatively cool lithosphere of the southwardmoving North American plate has been continuously 

placed on hot upwelling asthenosphere when crossing 

the Juan de Fuca plate boundary. As pointed out by 

Lachenbruch and Sass (1980), this process can explain 

the heat-flow anomaly in the North American plate 

that peaks in the Coast Ranges about 300 km south of 

the latitude of Cape Mendocino (Lachenbruch and 

Sass, 1973). Calculations by Lachenbruch and Sass 

(1980) show that, given a velocity of 5 cmlyr for 

movement of the Pacific plate relative to the North 

American plate, the heat flow should increase by a 

factor of 2 approximately 200 km south of the edge of 

the Juan de Fuca plate because 4 Ma is required for 

the heat anomaly to reach the surface from 20-km 

depth. These various parameters agree with the 

observed heat-flow anomaly. For a heat source as 

deep as 20 km, the model requires the hot asthenosphere to accrete to the bottom of the North American 

plate and to be conveyed off southward, so that a 

continuous supply of vertically moving, hot asthenosphere be supplied to the bottom near the Juan de 

Fuca plate boundary. This hypothesized coupling 

involves a rather thin layer of accreting upper mantle 

that, in turn, is probably decoupled from underlying 

asthenosphere. The gravitationally predicted depth to 

the base of the North American plate is within the 

limits required by Lachenbruch and Sass, (1980) 

model, at least within 70 km of the San Andreas fault. 

Interpretation of geologic and geophysical data for the 

San Andreas fault system north of San Francisco 

(Griscom and Jachens, 1989) suggests that eastward 

migration of the plate boundary from its presumed 

original position at the base of the continental slope to its 

present position at the San Andreas transform fault may 

r SYSTEM. CALIFORNIA 

have occurred by means of a series of eastward jumps of 

the Mendocino triple junction covering a total distance of 

about 150 km during the past 29 Ma. Our general model 

for the history of this triple junction is one of successive 

eastward jumps, with sustained periods at each position 

while sienif~cant strike-slip motion occurred on the various transform fault systems, including the San Andreas 

fault. We are aware, however, that the picture in detail 

may have been far more complex. The present position of 

the San Andreas fault north of San Francisco is thus 

interpreted to be relatively youthful. The triple junction 

was initially situated near the base of the continental 

slope at the northwest end of the Miocene(?) accreted 

wedge (but far to the south of its present latitude); the 

basal fault (McCulloch, 1987) below the subduction complex is shown as a toothed line in figure 9.4 because of the 

thrust component in this transform fault. The triple 

junction is interpreted to have been subsequently situated at the north end of the proposed Pilarcitos fault 

extension and then to have jumped eastward a minimum 

of about 100 km to the present San Andreas fault trace at 

what is now approximately lat 3820' N. on the North 

American plate. When this jump occurred, the three 

faults that formed the junction all had to readjust; the 

simplest scenario is as follows: (1) The Mendocino fault 

was extended on strike farther eastward, for the distance 

of the jump, about 100 km, (2) a new segment of the San 

Andreas fault broke obliquely through the Franciscan 

assemblage to the northwest (severing the correlated 

geophysical anomalies described above) and extended 

from the new triple junction to the junction of the newly 

formed (or soon to be formed) San Gregorio fault with the 

Pilarcitos fault, a distance of about 250 km; and (3) the 

surface trace of the subduction zone north of the triple 

junction also jumped eastward 100 km, thus abruptly 

isolating a thin triangular slab of Franciscan assemblage 

(probably less than 15 km thick) from the North American plate. This postulated triangular slab of rocks is now 

gone, most likely subducted away. Further complexity is 

provided by the King Peak subterrane of the King Range 

terrane (McLaughlin and others, 1982), which is an 

elongate mass of turbidites, about 45 km long, just south 

of Cape Mendocino (fig. 9.4) that is believed to have been 

obductively accreted from the west during the early 

Pleistocene (McLaughlin and others, 1986). The King 

Peak subterrane may have been detached and transported northwestward from the San Francisco area (just 

south of lat 3820' N.) as part of the Pacific plate and then 

reattached to the North American plate by a very recent 

local jump of the triple junction westward less than 35 km 

(McLaughlin and others, 1982); this explanation may 

account for the anomalously higher thermal metamorphism of this subterrane relative to the terranes that are 

now adjacent to it. Recent work suggests that the triple 
9 CRUSTAL AND LITHOSPHERIC STRUCTURE FROM GRAVITY AND MAGNETIC STUDIES 255 

junction may be on shore at Cape Mendocino (Clarke, 

1988; McLaughlin and others, 1988); if so, the King Peak 

subterrane may still be essentially part of the Pacific 

plate. The tectonic interpretation detailed above also 

requires that the San Gregorio-Hosgri fault first began 

moving and joined the present San Andreas fault at 

approximately the same time as or shortly after the 

eastward jump of the triple junction, and thus cut off the 

proposed northward extension of the Pilarcitos fault, 

after which the extension became inactive. 

The proposed 150-km eastward movement of the triple 

junction can also explain the submarine topography near 

Cape Mendocino, where the Continental Shelf south of 

the Mendocino fault extends about 130 km farther west 

than that directly north of the fault. 

The timing of the jump can be estimated from the 

horizontal offset of the paired geophysical anomalies, 

about 250 km, which translates to an age of about 5 Ma, 

assuming combined strike-slip rates of 4.8 cmlyr (DeMets 

and others, 1987; Minster and Jordan, 1978) for the San 

Andreas and San Gregorio faults. This age estimate is 

crude because it assumes that no other faults were 

absorbing the relative motion between the two plates. 

For example, simultaneous movement on the HaywardCalaveras fault system will cause the computed age of 

offset to be too young. The eastward jump of the triple 

junction appears to be associated with a change in stress 

orientations in this region. The north end of the San 

Gregorio-Hosgri fault trends about 20' clockwise relative 

to the older fault traces. In addition, the northwardmigrating triple junction subsequently traced out a major 

right-lateral bend, as shown by the present position of 

the Ran Andreas fault north of Point Arena. The central 

section of this bend is about 100 km long and trends 20' 

clockwise to the older trace. This change may correlate 

with the gradual change in absolute motion of the Pacific 

plate that occurred between 5 and 3.2 Ma (Cox and 

Engebretson, 1985; Pollitz, 19861, producing a change 

from strike slip to transpression in this region and a 

clockwise rotation of 20Â (Harbert and Cox, 1986) in the 

relative-velocity vector for the plate pair, the same angle 

as the anomalous change in direction for both the San 

Gregorio fault and the right-lateral bend in the San 

Andreas fault north of Point Arena. This change in 

relative motion probably correlates with a change in 

strike direction of the subducting south edge of the Juan 

de Fuca plate, as deduced from gravity data (Jachens and 

Griscom, 1983). Before the jump, this strike was eastwest, thus permitting eastward extension of the Mendocino transform fault without interference; after the jump, 

the strike of the subducting plate edge changed to S. 60' 

E., making later eastward fault extension more difficult. 

Stratigraphic evidence for the postulated eastward 

jump of the triple junction about 5 Ma may be sought in 

the late Miocene and Pliocene stratigraphy of Deep Sea 

Drilling Project (DSDP) Site 173 (fig. 9.4). Depositional 

hiatuses occur at 5 and 4.3-3.2 Ma (Barron, 19891, 

whereas a study of both micropaleontology and tephra 

beds indicates a hiatus from about 4.4 to 2.8-Ma (SarnaWojcicki and others, 1987). McCulloch (1987, fig. 26) 

believed that the middle Pliocene deformation and minor 

erosion interpreted from reflection profiles correlate with 

this 4.4-2.8-Ma hiatus at Site 173. We suggest that the 

eastward jump of the triple junction about 5 Ma was 

shortly followed by the middle Pliocene deformation and 

by the hiatus at Site 173. These two correlative events 

were thus caused both by the jump and by the simultaneous change in the direction of relative plate motion. 

SALTON BUTTES SPREADING CENTER 

The San Andreas fault terminates to the southeast in a 

buried spreading center at the south end of the Salton 

Sea, where a row of five small siliceous volcanic domes 

("buttes") protrude above recent sedimentary deposits of 

the Salton Trough. These domes, in addition to being 

associated with a local northeast-striking magnetic high, 

are situated on the crest of a larger, northwest-trending 

magnetic high (outlined on fig. 9.4) that is interpreted 

(Griscom and Muffler, 1971) to be caused by a magnetic 

mass, 30 km long, 3 to 12 km wide, and about 4 km thick, 

with its top buried more than 2 km below the surface. 

This magnetic high is associated with a similarly shaped 

gravity high (Biehler and Rotstein, 1979). the source of 

which may partly be the magnetic mass but may also be 

the relatively high density metamorphosed sedimentary 

rocks associated with the geothermal area (Elders and 

others, 1972). The Salton Buttes spreading center probably strikes northeast because the row of domes, the local 

aeromagnetic and gravity anomalies, and the geothermal 

area all coincide and strike northeast; (2) this proposed 

position for the center bisects the larger northwesttrending magnetic high into approximately equal parts 

interpreted to be new "oceanic" crust; and (3), ideally, a 

spreading center should trend approximately normal to 

an associated transform fault. In apparent contradiction, 

the Brawley seismic zone strikes S. 20' E. from the 

Salton Sea (Johnson and Hill, 1982) and consists of 

shallow earthquakes (Severson and McEvilly, 1987) lo- 

:ated mostly within the valley fill; this seemingly anomalous direction may be due to accommodation of these 

averlying, partly decoupled materials to a series of short 

northeast-trending spreading centers between the Salton 

Sea and Cem hieto, Mexico (see fig. 3.6), on strike S. 

20' E. and 100 km distant (Fuis and Kohler, 1984; Sibson, 

1987). The large, northwest-trending magnetic mass is 

interpreted to reflect about 30 km of northwestward 

spreading along its long axis, in which the spreading was 
256 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

associated with intrusive activity that built up a 30- 

km-long strip of magnetic mafic rocks and new crust in 

the lower section of and below the sedimentary fill. This 

magnetic feature may not be directly comparable to 

oceanic-crustal anomalies because slow cooling beneath 

the fill probably results in weak remanent magnetization, 

unlike the situation for oceanic crust. This anomaly thus 

may be predominantly caused by induced magnetization. 

The gravity field of the Salton Trough, which is filled 

with great thicknesses of Cenozoic sedimentary rocks, 

varies systematically from north to south. An elongate 

gravity low of -30 to -40 mGal is associated with the 

sedimentary rocks northwest of the Salton Sea (beyond 

lat 33'20' N.). Southward along the axis of the trough, 

the gravity field increases rapidly until the south end of 

the Salton Sea, where maximum values of 0 mGal are 

obtained over the presumed spreading center described 

above. Farther southeast, to the United States-Mexico 

border, gravity values range from only -10 to -20 mGal, 

an initially surprising observation because the 3.5 km or 

more of young, unmetamorphosed sedimentary deposits 

in this area might be expected to produce anomalies 

lower than -40 mGal (Biehler, 1964; Griscom, 1980c, p. 

20), similar to the gravity expression northwest of the 

Salton Sea. Biehler (1964) offered two explanations for 

the missing low: thinner crust or local high-density 

basement beneath the trough. Seismic-refraction studies 

(Fuis and others, 1982, fig. 17A) confi the second 

explanation and show a deep "subbasement" (density, 3.1 

g/cm3) in the trough that extends below about 12-km 

depth. Using this refraction model as a constraint, a 

gravity model (Fuis and others, 1982, fig. 20) indicates 

that the crust beneath the trough is no thinner than that 

of the bordering mountains a few kilometers to the 

northeast. 

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A s the Pacific plate slides northward past the North American 

plate along the Sun Andreas fault, the frictional stress that 

resists plate motion there is overcome to cause earthquakes. 

However, the frictional heating predicted for the process has never 

been detected. Thus, in spite of its importance to an understanding 

of both plate motion and earthquakes, the size of this frictional 

stress is still uncertain, even in order of magnitude. 

10. STRESS AND HEAT FLOW 

By ARTHUR H. LACHENBRUCH and A. MCGARR 

CONTENTS 

w IntroductionÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã 261 

Estimates of average stress from fault energetic*--------- 262 

Energy balanc+Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘ 262 

Apparent stress; Seismic estimate of 7,--------------- 263 

Friction: Thermal estimate of Ã‘Ã‘Ã‘Ã‘Ã‘Ã 264 

SummaryÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã 

Estimates of average stress from laboratory measurements of 

frictionÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘ 

Rock friction and the strength of the faultÃ‘--------- 267 

The case of equal strength in all directiow---------- 268 

The of a weak direction- ..................... 269 

INTRODUCTION 

As one of the best exposed tectonic-plate boundaries in 

the world, the San Andreas fault provides an excellent 

opportunity to study the forces causing interplate motion 

and the associated great earthquakes. Thus, there is 

considerable motivation, scientific, social, and economic, 

to understand the thermomechanics of the San Andreas 

fault system, which has been the subject of intensive 

studies for the past several decades. 

Although substantial progress has been made in unraveling the complex kinematics of the San Andreas fault 

system (Atwater, 1970; Minster and Jordan, 1984; Weldon and Humphries, 1986), efforts to determine the 

stresses that give rise to San Andreas fault slip, to date, 

have not led to anything resembling scientific consensus. 

The uncertainty results from widespread disagreement 

over the implications of different methods of assessing 

the stresses. 

The question of how much shear stress acts on the San 

Andreas fault to cause dextral slip began to acquire 

definition in 1968, when the first heat-flow data adjacent 

to the fault zone (fig. 10.1) were gathered and analyzed 

by Henyey (1968). Because these data did not reveal any 

anomalous heat flow near the major active faults of the 

San Andreas system, upper bounds of about 10 to 20 MPa 

on the average frictional stress resisting fault motion 

could be calculated (Brune and others, 1969; Henyey and 

Wasserburg, 1971). These upper bounds were taken as 

evidence confirming speculation on the low strength of 

the crust based on earthquake stress drops, almost 

invariably in the range 0.1-10 MPa (for example, Chinnery, 1964, Brune and Allen, 1967). At the same time, 

however, laboratory experiments indicated typical frictional strengths for precut rock samples of about 100 MPa 

under pressure and temperature conditions thought to 

obtain in the upper crust (Byerlee and Brace, 1968, 1969; 

Byerlee, 1970). 

4 FIGURE 10.1.-Heat-flow and stress measurements are taken in wella such SK thk one being drilled by the U.S. Geological Survey at the 

Crysiallaire site. 4 km northeast of the San Andreas fault in the western Mojave Desert. Photograph by M.D. Zoback. 

261 

THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA: U.S. GEOLOGICAL SURVEY PROFESSIONAL PAPER 1515 
262 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

Over the next several years, new heat-flow measurements supported the absence of any local heat-flow 

anomaly associated with the San Andreas fault (Lachenbruch and Sass, 1973) and thus augmented the position 

for low frictional fault strength. The recognition of a 

broad heat-flow high coincident with the Coast Ranges of 

California led Lachenbruch and Sass (1973) to suggest 

that partial decoupling at the base of the seismogenic part 

of the crust might account for both the weak fault 

(minimum in shear stress at the fault trace) and the broad 

thermal anomaly. 

Additional laboratory experiments on different rock 

types, and in conditions of higher temperature and 

confining pressure than had been obtained previously, 

continued to support high frictional strength in the top 15 

to 20 km of the fault zone (Stesky and Brace, 1973). The 

experimental results are most simply characterized in 

terms of a coefficient of friction that varies little with rock 

type (Byerlee, 1978), slip rate, or slip history (Dieterich, 

1979; Ruina, 1983). As emphasized by Brace and Kohlstedt (1980) and Kirby (1980), these results still indicate 

a high-strength upper crust. 

Beginning in the late 1970's, inplace stress measurements have provided another way to assess the stress 

acting on the San Andreas fault (Zoback and others, 

1977), especially with the advent of stress measurements 

at depths approaching 1 km only a few kilometers distant 

from the fault (Zoback and others, 1980). If the observed 

depth gradient for the component of shear stress thought 

to act on the San Andreas fault could be extrapolated to 

the base of the seismogenic zone, as argued by McGarr 

and others (1982), then the corresponding frictional 

stress resisting fault motion is a factor of 3 greater than 

the upper bound from the heat-flow analyses, as presented most recently by Lachenbmch and Sass (1980). 

The most recent developments, if accepted at face 

value, could be construed as additional evidence favoring 

a low-strength San Andreas fault. Specifically, stressdirection indicators on either side of the fault have been 

interpreted to mean that there is almost no shear stress 

resolved on the fault plane, thus implying a very weak 

fault zone (Mount and Suppe, 1987; Zoback and others, 

1987). If so, then the question regarding the strength of 

the fault would be answered, and the outstanding problem would be the equally vexing one of understanding the 

nature of a remarkably weak fault zone. 

This chapter is largely a review and commentary on the 

different approaches taken to estimate the tractions 

acting on the San Andreas fault. We restrict our attention to three main methods: (1) inferring stress from the 

fault's energy budget (thermal and kinetic), (2) inferring 

fault strength from laboratory measurements of the 

stresses needed to slide rocks past one another under 

pressure, and (3) inferring stress on the fault from 

observations of the crustal state of stress. 

ESTIMATES OF AVERAGE STRESS FROM 

FAULT ENERGETICS 

ENERGY BALANCE 

In figure 10.2, an earthquake is viewed, according to 

Reid's (1910) rebound theory, as a strained patch of fault 

surface of area A that suddenly breaks, permitting points 

hitially in contact to be displaced from one another by an 

iverage amount u. The breakage is like the sudden 

hilure of an overloaded leaf spring. We are interested in 

the average shear stress acting parallel to the wall in the 

Sailed section of the fault surface. We denote its initial 

value by and its final value by r,. The inclined line in 

figure 10.3 represents the elastic unloading of the medium as the earthquake displacement increases to its final 

value u. The area under this line, which represents the 

total elastic energy released by the earthquake per unit 

area of faulted surface, can be written as 

1 where ?=z(rl+ra). 

The energy E must supply the work Ec of generating 

seismic waves and the work Em converted to heat in 

overcoming frictional resisting forces. Thus, 

where the question mark is a reminder (which we shall 

forget for the moment) that there may be other significant sinks of earthquake energy, such as the surface 

FIGURE 10.2.-Elastic-rebound theory (Reid, 1910). showing displacement near a strike-slip fault segment of area A before and after an 

earthquake with displacement Ã§ Arrows along fault indicate dkction of relative movement 
10. STRESS AND HEAT FLOW 263 

energy consumed in creating new fractures. We can now 

write 

where T,, the "apparent stress" of seismology, is the 

portion of the earthquake stress T allocated to the 

production of seismic waves, and T, is the average 

frictional resisting stress allocated to the production of 

heat. The individual areas represented by equations 3 

and 4 are shown in figure 10.3 by contrasting patterns. 

This interpretation of the areas in figure 10.3 is fairly 

general, as long as we define TI, TZ, and T,, respectively, 

as the weighted averages of initial stress, final stress, 

and friction over the faulted surface, the weighting 

function being the local fault slip (see Savage and Wood, 

DISPLACEMENT. IN MICROMETERS 

FIGURE 10.3.-Relation between resistii stress and displacement in 

the unloadingelastic medium (inclined line) duringan earthquake. As 

ilip (u) increases, stress in the rock diminishes linearly from T, to TZ, 

with average value T. Areaunder this line is total work expended per 

unit fault area; area (shaded) below curve of resisting stress (T,(H), 

with average value T,) is energy dissipated per unit fault area (EJ. 

Difference between total work expended and dissipated energy is 

work (done by apparent stress 7.1 that is available for seismic 

radiation (EÃ stippled area). Modified from a laboratory experiment 

on a large granite sample by Lcckner and Ohbo (1983). 

1971; Lachenbruch and Sass, 1980). Combining equations 

1 through 4 yields 

which states that unless the question mark represents 

something important that we've neglected, the average 

earthquake stress T is the sum of the apparent stress T., 

to be estimated from seismic measurements of Ea (eq. 3), 

and the resisting stress 1-, to he estimated from thermal 

measurements of frictional heat E. (eq. 4). 

APPARENT STRESS: SEISMIC ESTIMATE OFT. 

Seismologists (for example, Wyss and Brune, 1968; 

Savage and Wood, 1971; Wyss and Molnar, 1972) have 

defined apparent stress as 

where -q is the seismic efficiency, defined by 

where equation 7b follows from 7a according to equations 

2 through 5; that is, -q is simply the fraction of the total 

energy release, or the fraction of the average earthquake 

stress, allocated to seismic radiation. 

To estimate T,, seismologists first determine the radiated energy E, and the seismic moment Mo, defined as 

Mo = GAu , (8) 

where u is the average slip of an earthquake over a 

fault-surface area A. and G is the modulus of rigidity 

(Aid, 1966). Equations 1 and 6 through 8 then yield the 

simple relation 

A numerical estimate of can be obtained from 

equation 9 with the following commonly used formulas 

relating earthquake magnitude M to E. or Mu (Gutenberg and Richter, 1956; Hanks and Kanamori, 1979), 

log Mo = 16 + 1.5M (1013) 

log E. = 11.8 + 1.5M, 
264 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

where M,, and E, are in ergs. Substitution of equations 10 

in 9 yields 

With G=3x lo4 MPa, the value for T, is 2 MPa. Almost 

without exception, estimates of T= fall within the range 

0-5 MPa, with no indication of any systematic dependence on either earthquake size or tectonic environment 

(Spottiswoode and McGarr, 1976; Fletcher and others, 

1983; Boatwright and Choy, 1986). In short, 5 MPa 

appears to be a conservative upper bound to ia. Thus, the 

contribution of T. is small, and the average fault stress T 

can be large only if the frictional resistance T, is large (eq. 

5). If laboratory "earthquakes" are proper analogs of 

crustal earthquakes, which may or may not be the ease, 

then data for such events, including those illustrated in 

figure 10.3, indicate that T, is indeed small, only a tiny 

fraction of T_ By inducing unstable frictional failure 

(earthquakes) across a 2-m-long fault between slabs of 

granite 40 cm thick (Dietericb, 1981), Lockner and Okubo 

(1983) measured seismic efficiencies q for numerous 

events to conclude that q-O.W. If this result were true 

also for natural earthquakes-a big "if'-then for a 

typical value T of 2 MPa, the corresponding value of i, 

from equation 6, would be 40 MPa, which, as will be seen, 

is nearly 3 times higher than the limit inferred from an 

analysis of heat-flow data (Lachenbmch and Sass, 1980). 

FRICTION: THERMAL ESTIMATE OFT, 

Unlike the energy of seismic waves, which permits an 

estimate of apparent stress T, for individual earthquakes 

from measurements at distant stations, the heat energy 

of individual earthquakes is not readily analyzable to 

estimate friction because it causes a measurable temperature rise only withm a few meters of the earthquake 

source, a location inaccessible for measurement. Even for 

the largest earthquakes, these individual temperature 

pulses would be indistinguishable from background a few 

months after the event, and so timely attempts to detect 

them by drilling would be difficult (McKenzie and Brune, 

1972; Lachenbmch, 1986). However, because the frictionally generated heat diffuses quite slowly, it should 

accumulate in the vicinity of the fault, eventually building 

up the local thermal gradient until the observable heat 

loss at the Earth's surface in the fault zone exceeds the 

normal background heat flow by the rate of heat generation on the fault. Thus, in principle, the measurement of 

a heat-flow anomaly in the fault zone should permit an 

estimate of the average frictional contribution T- to the 

earthquake fault stress T (eq. 5). 

The heat-flow anomaly that we seek does not depend 

on the amount of heat E, liberated by a single earthquake 

in a restricted fault area with a displacement u (eq. 4), 

but on the long-term average rate of heat generation (Q) 

and the 10%-term average slip rate (v) from the cumulative effect of successive events. Although most fault 

displacement probably occurs within a few tens of seconds during large earthquakes every century or so at slip 

velocities greater than the average ones by a factor of 

about 108, the long-term buildup of the heat-flow anomaly 

would be indistinguishable from that caused by uniform 

slip at an equivalent average velocity because the thermal time constant for the buildup (approx lo6 yr) is large 

relative to the earthquake-recurrence interval (10'-10~ 

yr). Therefore, we view the slip as a uniform continuous 

process and introduce 

and 

Differentiation of equation 4 yields 

where v is the long-term average slip velocity, -rr is the 

(displacement averaged) frictional resistance, and Q is 

the long-term average rate of frictional heat production 

per unit area on the fault surface. 

Equation 13 refers to the entire seismogenic (brittle) 

layer (approx 10-15 km thick), not just a patch as in 

equation 5. Over this depth, it is reasonable to consider 

the long-term slip velocity v to be independent of depth, 

but generally the heat-production rate Q will not be. For 

example, if the friction T- increased linearly with depth 

(for example, because of increasing gravitational pressure on the fault, as discussed below), the heat production Q on the fault would also increase linearly, as shown 

in figure 10.45. According to heat-conduction theory, the 

temperature in the fault plane would then build up over 

time, as shown in figure 10.4C. and a heat-flow anomaly 

would develop at the surface over the fault, as shown in 

figure 10.4A. For such a distribution, a sharp heat-flow 

anomaly is seen to build up over the fault in about 1 my.; 

after several million years, it approaches a maximum 

value somewhat greater than half the average frictional 

heat production Q on the fault surface (fig. 10.45). This 

anomaly falls off over a distance from the fault of the 

same order as the depth of the seismic layer (assumed to 

he 14 km in fig. 10.4). Other reasonable distributions of 
10. STRESS AND HEAT FLOW 

frictional sources give similar results (see Henyey, 1968; 

Brune and others, 1969: Henyey and Wasserburg, 1971; 

Ladienbruch and Baas, 1973, 1980). 

The long-term slip rate v, which can be estimated from 

studies of offset (and dated) geologic features, generally 

SOURCE STRENGTH 

ranges from 2 to 4 cmlyr for motion on the main trace of 

the San Andreas fault over the past several million years 

throughout California (see chap. 7; Grantz and Dickinson, 

1968). As a useful rule of thumb, if the fault in figwe 10.4 

were slipping at an average rate (v) of 3 cmlyr and 

Y'X* Y 

CHMENSIONLESS KILOMETERS 

HORIZONTAL DISTANCE FROM FAULT 

DIMENSIONLESS TEMPERATURE 

9@Jx2K)-' 

I Model 1 

TEMPERATURE RISE, IN DEGREES KELVIN 

FIGURE 10.4.-Surface heat flow q (A) and fault-plane temper- heat generation on fault; K, thermal conductivity. Dimensionatwe 0 (C) for a linear increase in source strength from zero al results are for %=I4 km, K=2.5 WImK, and Q=42 m~/d 

at the surface to Qz at depth + (B). t, time since initiation of (equivalent to 2v=25 mmlyr, ?=50 MPa) (Lachenbmch and 

faulting; A, dimensionless time; Q, average rate of frictional Sass, 1980). 
266 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

resisted by an average frictional stress T, of 100 MPa, 

then the average rate of frictional heat production Q 

(figure 10.45) would be about 0.1 W/m2; that is, 

0.1 W/m2 = 100 MPa x 3 cmlyr. (14) 

This quantity is about twice the stable continental heat 

flow, and so, according to figure 10.4A, the corresponding heat-flow anomaly over the fault would be about 100 

percent of background after 2 or 3 m.y. of fault slip, 

whereas if the mean frictional resistance were only 10 

MPa, the corresponding heat-flow anomaly would be only 

about 10 percent of background, close to the limit of 

detection. Accordingly, if no heat-flow anomaly could be 

detected over the fault, the mean frictional resistance 

would be no more than about 10 MPa; if the mean 

frictional resistance were about 100 MPa, a very conspicuous anomaly should be observed. 

An example of heat-flow measurements near the San 

Andreas fault is shown in figure 10.5 for the Mojave 

Desert region of southern California (region 7, fig. 10.6). 

The pattern of anomaly curves from the model in figure 

10.4 is scaled for the estimated local slip velocity (25 

mm/yr; Weldon and Sieh, 1985) and for a mean frictional 

resistance of 50 MPa. Clearly, the data are incompatible 

with such an anomaly in fact, the average heat flows near 

the fault and far from the fault ("within 10 km" and 

"beyond 10 km," figs. 10.7C, 10.723) are statistically 

indistinguishable. Figures 10.7A and 10.7B show that the 

same condition prevails in the Coast Ranges of central 

California (regions 3-6, fig. 10.6). In fact, no local 

heat-flow anomaly has been confirmed anywhere on the 

main trace of the San Andreas fault (for possible exceptions, see Lee, 1983, Lachenbmch and Sass, 1988), and 

so, according to the foregoing simple considerations, the 

averaee friction on the fault. T- nrobahlv does not exceed 

SUMMARY 

In summary, we note that analysis of the kinetic 

energy of seismic waves suggests that the associated 

apparent stresses (T.) do not exceed 5 MPa. Similarly, 

analysis of long-term frictional heating and the predicted 

and observed effects on heat flow from conduction theory 

FIGURE 10.5.-Heat flow w a function of distance from the San An& fault in the Mojave sement (region 7, fie. 10.6). The~retical anomaly 

is for a slip velocity of 25 miw\r and average friction of 50 AfPa (Lachenbruch and Sass, 1988). 

125 

IT [" 100- 

s 

kJ 

a 

2 

IT 

u 

0. 

CO E 75- 

j 

5 

z 

s 

3 ". 

50- r 

25 

sw ' I I I I I 

N! 

ec I- 9* 

4 2 z 

s 

I 

-69 

I 

EXPLANATION 

I 

I 

Data from U.S. Geological Survey I 

I Data from California Institute of Technology 1 

I 

I I I I I 

60 40 20 0 20 40 

I 

DISTANCE FROM SAN ANDREAS FAULT. IN KILOMETERS 
10. STRESS AND HEAT FLOW 

suggest that the average frictional resistance T, does not 

exceed about 10 MPa. Thus, according to equation 5, the 

long-term average combined earthquake stress T probably does not exceed about 15 MPa, and, of course, it could 

be much less. The initial stress T~, or "fault strength," 

would be greater by half the stress drop (fig. 10.2; 

Lachenbmch and Sass, 1980, eq. 41a), no more than 

another 5 MPa, for an upper limit of 20 MPa. The major 

assumptions in this analysis are (1) that heat transfer is 

exclusively by conduction-that is, no significant heat is 

transferred by moving ground water; (2) that the fault 

geometry can be represented by the usual simple conventions (see figs. 10.2, 10.3, and 10.4); and (3) that an 

earthquake's energy is converted exclusively to seismic 

waves and heat-that is, no appreciable energy is consumed by creating new surfaces, phases, or chemical 

combinations (Lachenbmch and Sass, 1973, 1980). We 

shall discuss these points later. 

ESTIMATES OF AVERAGE STRESS FROM 

LABORATORY MEASUREMENTS OF FRICTION 

ROCK FRICTION AND THE STRENGTH OF THE FAULT 

We have seen that the average shear stress T on an 

earthquake fault can be viewed as the sum of a dynamic 

part T, and a frictional part T,. The dynamic part is shown 

0 - 300 KILOMETERS 

I 

FIGURE 10.6.-Locations of heat-flow measurements near the San 

Andreaa fault and of numbered reeions referred to in figure 10.7 

(Ladienbmch and Sass, 1980). ~eavyline, San Andreas dashed 

where approximately located, dotted where concealed; stippled area, 

Great Valley. 

to be small from seismic evidence, and BO the earthquake 

stress must be large or small according to the size of 7,- 

We have also seen that T, is small according to geothermal evidence. We now consider a second line of evidence 

from laboratory measurement of friction which suggests 

to many that, contrary to the geothermal evidence, T~ 

must be large. 

According to these results, rock surfaces will slide 

when the shear stress on their surface of contact exceeds 

the static frictional strength TC given by 

where a = a,, - P 

a,, is the normal pressure pushing the surfaces together, 

and Pis the fluid pressure in the pores and cracks tending 

to hold the surfaces apart; <r; is called the "effective" 

normal stress (we generally denote such effective stresses by a prime, " ' "). The proportionality constant p. in 

equation 15a is the coefficient of static friction; extensive 

laboratory experiments show that its value is generally in 

the range 0.6-0.9 for a remarkably large variety of rock 

types and surface conditions (Byerlee, 1978), although 

some studies (for example, Wang and others, 19801, 

reported substantially lower friction coefficients for some 

geologic materials, including certain types of fault gouge. 

REGIONS 3 THROUGH 6 REGION 7 

10iÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘lÃ‘Ã‘Ã‘Ã‘Ã‘Ã‘[Ã‘Ã‘Ã‘Ã‘Ã‘Ã‘16jÃ‘Ã 

SURFACE HEAT FLOW, q, IN MILLIWATTS PER SQUARE METER 

FIGURE 10.7.-Comparison of heat flow within 10 km of main trace of 

the San Andreal firnit (A, 0 and beyond 10 km (B, Dl for redm 3-6 

(A. B) and redon 7 (C, D) (see fig. 10.6 for locations); transitional 

regions 1 and 2 are not represented. Modified from Lachenbrueh and 

Sasa (1980). n, number of samples; a, mean heat flow, SE, standard 

mor. 
268 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

We presume that a fault is a fracture with little 

cohesive strength that remains inactive until the natural 

shear stress T resolved along it exceeds its frictional 

strength T~ given by equation 15. This shear stress, which 

depends on the magnitudes of the principal stresses and 

on the angular relation between the fault plane and the 

principal stress directions (fig. 10.8), is given by (Jaeger, 

1956, p. 8) 

T = (vl - ma) sin 26 , 2 

where it is assumed for convenience that the intermediate-principal-stress direction (ua) lies in the fault plane 

(true if pi is independent of the orientation of this plane). 

In figure 10.8, v, is vertical, and v1 and v3 are the 

maximum and minimum horizontal principal stresses. 6 

is the angle formed by the fault normal and the direction 

of least compression (v3); it is also the angle between the 

fault trace and the direction of greatest compression (ul). 

To express the failure criterion (eq. 15) in terms of the 

stress field and fault orientation, we note that the 

effective normal stress. o;, in equation 15 can be written 

as (Jaeger, 1956, p. 8) 

FIGURE 10.8.-Conventions for discussing orientation of fault relative 

to direction of principal ati-eases: o,>%>oo, Arrows indicate direction of relative movement along fault. 

With equations 15 through 17, the friction stress T~ that 

must be exceeded on a fault for it to slip can be 

determined if we know (1) the maximum and minimum 

principal stresses 01 and v3, (2) the fluid pressure P, (3) 

the coefficient of friction pi, and (4) the angle 6 describing 

the orientation of the fault relative to the principal-stress 

axes. 

As we increase the stress difference, in what direction 

(6) will the Earth ultimately break, and what will be the 

stress on the failure plane? Clearly, the answer could be 

influenced by the existence of planes of weakness (McKenzie, 1969); for example, major preexisting faults or 

foliated country rock might result in directions with 

anomalously low pi. 

THE CASE OF EQUAL STRENGTH IN ALL DIRECTIONS 

We first assume that no such directional strength 

variation exists, that the rock is fractured in all directions, and that all potential shear surfaces have the same 

coefficient of friction pi. In this case, the foregoing 

equations show that the trace of the favored fault plane 

will depart from the direction of maximum compression 

by an angle â‚¬l dependent only on the coefficient of 

friction. as follows: 

Note that generally eo<45" (the direction of the surface 

of maximum resolved elastic shear stress, eq. 16) because 

of the effects of normal stress on friction (Jaeger, 1956). 

With this additional relation (eq. 18), we can express the 

frictional strength T~ of a plane of orientation 6 in terms 

of the coefficient of friction and the effective-principalstress components as follows: 

- 

Tf = 4 

v; 1 + pi cot 8 

Tf 

- = pi 

vÃ 1 - pi tan 6 
10. STRESS AND HEAT FLOW 269 

To evaluate the frictional strength, the vertical stress 

is generally assumed to be a principal stress (reasonable 

because the Earth's surface supports no traction) equal to 

the rock column's weight, pas, an assumption supported 

by inplace stress measurements (McGarr and Gay. 1978). 

In this case, the vertical effective stress u; will be 

where p is the rock density, and the fluid pressure P is 

given by 

The value A=0 represents conditions in dry rock. For a 

typical open ("hydrostatic") hydrologic system, we have 

A-0.37 (= pn/proii, where p, is the density of water). As 

A-1, the fluid pressure approaches the weight of overburden, and the vertical effective stress a; vanishes (as 

discussed below, this limit probably occurs only in the 

thrusting regime, where a, is vertical). 

The curves in figure 10.9 (referred to ordinate scale at 

left margin) give the frictional strength normalized by 

the effective vertical principal stress for those cases in 

which the vertical stress is the maximum (dashed curve), 

average (solid curve), or minimum (dotted curve) principal stress, respectively. The first right-hand ordinate 

scale gives the increase in frictional strength with depth 

(T~/z) for the usual assumption of hydrostatic fluid pressure (P=p&z). For typical values of p. from Byerlee's 

results (for example, 0.6-0.9), the frictional strength for 

normal and thrust faults increases with depth at rates of 

about 5 and 20 MPa/km, respectively (fig. 10.9). The rate 

of increase for strike-slip faults lies between these limits; 

a commonly used value, 8 MPa/km, is shown by the solid 

curve in figure 10.9. For an upper-crustal fault extending 

to 14 km depth, these increases would result in average 

friction (the value at a 7-km depth) of 35,56, and 140 MPa 

for normal, strike-slip, and thrust faults, respectively 

(see second ordinate scale on right, fig. 10.9). Such 

calculations provide the basis for the expectation of high 

fault stress from the analysis of laboratory results: These 

values are substantially greater than the 20 MPa upper 

limit for initial stress suggested from the analysis of 

heat-flow data in strike-slip tectonic regimes (horizontal 

dashed line, fig. 10.9). Note that the heat-flow limit 

would require psO.2 for the assumed conditions. 

THE CASE OF A WEAK DIRECTION 

The estimates of large friction from the fault model of 

figure 10.9 depend on three principal assumptions: (1) 

that the average coefficient of friction on real faults is 

comparable to typical laboratory values (p-0.6-0.9), (2) 

that the average fluid pressure throughout the depth of 

the fault is comparable to the weight of the overlying 

column of water (A-0.37, eq. 20), and (3) that the 

coefficient of friction (p.) is the same in all directions, so 

that the fault direction (On) is determined by the applied 

stress (eq. 18) and not by the orientation of a special plane 

of weakness. Partly in response to recent reports that the 

maximum horizontal principal stress is oriented nearly 

perpendicular to the San Andreas fault (Mount and 

Suppe, 1987; Zoback and others, 1987), we drop the last 

assumption and suppose that the fault occupies a very 

weak plane (which is assumed to contain the intermediate 

principal stress). Because of the anomalous weakness of 

this plane, the friction along it could be very low, 

consistent with the heat-flow data, and faulting could 

persist there irrespective of the ambient stress field. 

2 

COEFFICIENT OF FRICTION 

FIGURE 10.9.-Variation ofnormalized fault friction (left-hand ordinate 

scale) with coefficient of friction under three conditions for vertical 

effective stress (0;) for failure at optimum angle â‚¬I Increase in fault 

friction T, with depth s under hydrostatic fluid pressure ie (riven by 

right-hand ordinate scale, and mean friction TF on a fault 14 km deep 

by scale on far right. Horizontal clashed line is upper limit of mean 

friction suggested by heat-flow constraint. 0; and 4, maximum and 

minimum horizontal effective principal stress, respectively. 
270 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

According to the friction model (eq. 15), the two factors 

that might weaken the plane are either an abnormally low 

coefficient of friction or unusually high pore pressure. 

For now, we assume that each of these conditions can 

exist regardless of laboratory or hydrologic evidence. 

The first question we consider is whether a very weak 

fault can coexist with stronger faults such that both types 

are active, as may be the case along the San Andreas fault 

(for a closely related discussion, see Sibson, 1985). To 

address this question, it is convenient to express the 

crustal strength in terms of the ratio uilu; (Brace and 

Kohlstedt, 1980). From equations 19a and 19b, the 

condition at failure (eq. 16a) for a weak plane oriented at 

an angle 0 to the direction of ul (fig. 10.8) is 

For isotropic strength, failure occurs at (eq. 18), the 

direction in which uilu; is a minimum for a given p: 

Faults at angles other than go support greater deviatoric 

stresses and the higher values of u[laÃ given by equation 

21a. 

The conditions necessary for the coexistence of active 

faults with different coefficients of friction are illustrated 

in figure 10.10, where the ratio of effective principal 

stresses at the point of failure is plotted as a function of 

the fault angle for various values of the coefficient of 

friction (eq. 21a). Suppose, for example, that the coefficient of friction is only 0.1 in the direction of the San 

Andreas fault, whereas in all other directions it is 0.6. 

Because d-ilui must be at least 3.1 to cause faulting in the 

crustal environs, the low-strength San Andreas fault 

must be oriented at 9~3.5~ or 9a81.5Â (fig. 10.10); 

otherwise, uilu; would be too low to cause slip in the 

stronger directions. In this example, then, the weak fault 

must be oriented either nearly parallel or nearly perpendicular to the direction of ul. 

In the context of the notion that the San Andreas fault 

is nearly perpendicular to the direction of u:, or at 0-90- 

in figure 10.10, we note that a very low coefficient of fault 

friction is required. The strength curves for each value of 

p have two asymptotes where uflui Ã‘ m. These asymptotes occur where the denominator of equation 21a 

vanishes; one asymptote is at 0=0, or ul parallel to the 

fault, for any value of p, and the other is at 0=200 (eq. 

81, or 0=90-tan p. Thus, the normal to any fault that 

fails in shear must be oriented at an angle of at least tan-1 

p from the direction of up For the four curves in figure 

10.10, the right-hand asymptotes are at 6=84.3, 73.3, 

i9.0, and48.0, respectively, for p=0.1,0.3,0.6, and 0.9. 

Fhus, if the fault trace makes an angle greater than 69', 

hen p must be less than 0.6 as long as the fluid pressure 

3 less than the least principal stress. 

More generally, enhanced pore pressure alone cannot 

ead to active faults nearly normal to u, unless P>uÃ§ in 

vhich case u;<0 and failure is likely to manifest itself as 

lydraulic fracturing rather than fault slip. 

SUMMARY 

To recapitulate, the simplest interpretation of earthluakes in terms of the frictional fault model and laboraory measurements of rock friction leads to fault stresses 

nany times larger than the limits suggested from heatlow and fault energetics. This interpretation depends on 

hree assumptions: (1) that the average coefficient of 

riction on real faults is comparable to typical laboratory 

ralues (p-0.6-0.9), (2) that the pore-fluid pressure 

hroughout the depth of faulting is comparable to the 

veight of the overlying column of water, and (3) that the 

ntrinsic resistance of the Earth to sliding is isotropichat is, no weak directions exist. To reduce the high 

'stirnates of friction obtained from rock mechanics to the 

ow ones obtained from heat flow. we must assume either 

imaller values of the coefficient of friction p or larger 

9. IN DEGREES 

'IGURE 10.10.-Failure criteria for various coefficients of friction as 

functions of fault angle Q. For fixed a,', (TO, and (J,, slip mil occur for 

all 0 between points where horizontal line delined by o{/(T; intersects 

curve defined by p. For example shown, if i72~'=3.1, then for 

u.=0.1, slip can occur at all angles between 3' and 82' but only at 

optimum angle 29.5' if p=0.6. Dashed curve, optimal failure angle 

â‚¬lo (q. 18). a,' and 06, maximumandmmmium horizontal effective 

principal stress, respectively. 
10. STRESS AND HEAT FLOW 

values of fluid pressure P (see eq. 15). Of particular 

interest in this connection is reported evidence that the 

trend of the San Andreas fault in California might occupy 

an anomalous weak direction. According to Mount and 

Suppe (1987) and Zoback and others (1987), the fault 

plane is nearly perpendicular to the direction of maximum compression (0-VE, fig. 10.10). a direction in 

which the resolved shear stress is very small. Such a 

condition could be consistent with the low friction required by heat flow, while permitting high stresses to 

accommodate the subsidiary faulting observed on more 

favorably situated planes. This model, however, raises 

some basic questions regarding the mechanics of faulting; 

for the fault to slin. &. it must have a low shear strenrth. -. as 

well as the low shear stress suggested by its orientation. 

If conventional friction theory applies, anomalously high 

fluid pressure along the fault cannot readily account for 

the required low friction because, unless p is unusually 

low, the fault becomes exceedingly resistant to shear 

failure as 0 begins to approach ir12 (fig. 10.10). 

ESTIMATION OF FAULT STRENGTH FROM 

INPLACE STRESS MEASUREMENTS 

In principle, measurements of the magnitude and 

orientation of crustal stress in the vicinity of the San 

Andreas fault should provide the most direct evidence of 

the forces acting to cause interplate motion there. 

However, some essential problems exist with this approach. Because we have little understanding of the 

mechanics of the system, it is difficult to interpret the 

data. We are not dealing with alaboratory experiment in 

which a sample is loaded in a testing machine whose 

characteristics are well known; in such a situation, it is 

straightforward to use gages to estimate the magnitude 

of the load. In contrast to the well-controlled laboratory 

situation, we have little idea of the nature of the forces 

applied to the Earth's crust to cause a deviatoric state of 

stress and, in the case of tectonically active areas, slip 

across major throughgoing faults. We know neither 

where the forces are applied nor what is applying them; 

moreover, there is even debate about what the state of 

stress would be if only gravity were acting (McGarr, 

1988). 

In addition to the absence of a conceptual framework, 

there are numerous experimental difficulties in determining the state of stress, that is, the magnitudes and 

orientations of the three principal stresses as functions of 

position within the crust. Data must he obtained from 

depths below the zone of weathering, in rock that is 

sufficiently strong to support deviatoric stresses. In 

granitic rocks, this requirement, in effect, necessitates 

stress measurements at depths of about 60 m or more, 

thus limiting the measurement technique to hydraulic 

fracturing, the only common procedure that can be used 

at such depths (Haimson and Fairhurst, 1970). 

The hydraulic-fracturing, or "hydrofrac,'' method involves isolating a section of a borehole and then pressurizing this cylinder by pumping in fluid until a tensile crack 

forms and propagates into the previously unfractured 

rock. By monitoring the pressure-time history of the fluid 

in the isolated section, both the maximum and minimum 

horizontal stresses can be estimated (Hubbert and Willis, 

1957; Zoback and Haimson, 1983). This approach assumes 

that one of the principal stresses my is oriented vertically 

and can be calculated from the weight of overburden (eq. 

20). The other two principal stresses are the maximum, 

OH, and minimum, oh, horizontal stresses. In contrast to 

engineering usage, the convention adopted here is for 

compressional stresses to he positive because, in the 

Earth's crust, tensional stresses are rarely encountered, 

even at the surface. 

Although the uppermost crust near the San Andreas 

fault system has not been sampled as much for stress as 

for heat flow, enough inplace stresses have been measured to provide an indication of the state of stress there 

and how it compares with crustal stresses in other 

tectonic settings. To date, 41 successful hydrofrac measurements have been made in the 12 wells shown in figure 

10.11 at depths of as much as 850 m. A total of 29 of these 

data, in wells along the Mojave reach of the fault (fig. 

10.110, were analyzed by McGarr and others (1982). 

Since that study, four stress measurements have been 

made at Black Butte (BB, fig. 10.110 in the Mojave 

Desert (Stock and Healy, 1988), the data from the Hi 

Vista well have been reanalyzed by Hickman and others 

(1988), and additional measurements have been made in 

central California (Zoback and others, 1980). Currently, 

stress measurements are being made at the Cajon Pass 

well near the southeast end of the Mojave reach of the 

San Andreas fault, with some observations at depths 

below 3 km. Because no clear picture has yet emerged 

(see Healy and Zoback, 1988), we have not incorporated 

the Cajon Pass results into this review. 

The state of horizontal deviatoric stress can be characterized in terms of two parameters: the maximum 

horizontal shear stress T,,, given by 

and the angle 0 between the trace of the fault and the 

direction of maximum horizontal compressive stress 

Under favorable conditions, both parameters can be 

determined by the hydrofrac technique. We have shown 

that if 8-45Â¡ then T,,, is entirely resolved onto the plane 

of the fault to produce its slip; as 0 approaches 0Â or 90e, 
272 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

the resolved stress on the fault becomes arbitrarily small 

irrespective of the magnitude of im (eq. 16). 

.-"- 

Hughes HI vim MOJ8 A Lamcaster 4 + 

lJ4 

ich Bune 5 Pearblossom 

'+XTLR. MOJ2 

0 - 50 KILOMETERS 

EXPLANATION 

LKD + Well and measurement site 

- Fault-Dashed where approximately located. Arrows 

indicate direction of relative movement 

FIGURE 10.11.-Sketch maps of the Gabilan Range (A), central 

01 coast (B), and western Mojave Desert (0, showing 

locations of wells where stress measurements have been taken along 

the San Andreas fault wing the hydrofracturing technique. 

Evidence regarding the actual orientation of OH relative to the strike of the San Andreas fault is contradictory. Observations favoring 0 distributed about 45", so 

as to cause dextral fault slip, were presented by McNalley and others (1978), Zoback and others (1980), Zoback 

and Zoback (19801, and Hickman and others (1988); 

however, these data, from the Mojave Desert, show 

considerable scatter. In contrast, Mount and Suppe 

(1987) and Zoback and others (1987) reviewed a broad set 

of data, including many borehole breakout orientations, 

that suggest 0-90Â¡ Oppenheimer and others (1988) 

came to a similar conclusion. An intermediate result was 

obtained by Jones (1988). who stated that us is oriented 

at 65' to the local strike of the San Andreas fault in 

southern California. Thus, currently, we know neither 

the preferred value of 0 nor whether such a value even 

exists. For the foregoing values of 0 (46', 65', or W0), the 

shear stress resolved on the fault (eq. 16) would be imp 

0.77im, or 0, respectively. In view of thisuncertainty, we 

leave 0 unspecified and describe what is known of T~ the 

upper limit to the shear stress that can be resolved on the 

fault. 

The first-order feature seen in data from the San 

Andreas fault zone (fig. 10.12) is a marked tendency for 

im to increase with depth. The solid line, a regression fit 

to all of the data, indicates a depth gradient of 8.3 

MPalkm, not significantly greater than the gradient of 

7.9 MPah reported by McGarr and others (1982) on the 

basis of 29 of the 41 data plotted in figure 10.12. We note 

that the observed depth gradient of T-, also agrees well 

with the curves for strike-slip faults (solid curves. fie. 

10.9) for a coefficient of friction of 0.6 or greater. in 

addition to the general increase in im with depth, 

considerable variation from one well to another and 

within individual wells is suggested by figure 10.12. 

Figure 10.13 shows that the departure of the measured 

values of im from the regression line in figure 10.12 does 

not vary systematically with distance from the San 

Andreas fault. The principal conclusion to be drawn from 

figure 10.13 seems to be that the magnitude of deviatoric 

stress is not measurably affected by proximity to the San 

Andreas fault. Thus, whatever effect the fault may have 

on the magnitude of the shear stress, it is either too 

subtle, too localized, or too deep to be recognized in the 

current data set. 

the Mojave Desert residuals, measured near a locked 

section of the San Andreas fault, and those in central 

California (fig. 10.11A), where the fault is creeping and 

presumably does not produce great earthquakes. If 

measurements were made to greater depths, some differences might appear, but at least in the topmost several 

hundred meters, the magnitude of shear stress seems to 

be largely independent of position along the strike of the 

San Andreas fault. 

.3 We note that there is no detectable difference between 
10. STRESS AND HEAT FLOW 

Having failed to discover any spatial relation between 

the San Andreas fault and deviatoric-stress magnitudes, 

we now consider the question of whether or not any 

detectable differences exist between the stress states 

measured near the San Andreas fault (fig. 10.12) and 

those measured elsewhere in different tectonic settings. 

A review of crustal shear stress by McGarr (1980) 

considered a large suite of stress data in "hard" rocks 

measured at depths extendmg to 3.6 km. The resulting 

regression line of 

- 

6 

c F F 

T,,, (MPaj-1.15+8.28z (ki s 0.6 - 

1 

EXPLANATION 

0.7 - Distance from 

San AndSymbol Well fauk(km) 

A MOJI 2 

B MOJ2 4 

0.8 - c XTLR 4 

D Lake Hughes 4 

E MOJ4 22 

F Hi Vista 32 

G MOJ5 34 

0.9 - H LKB 2 

I LKC 4 

J LKD 14 

K San Aide 3 

L BlackButte 18 

MAXIMUM HORIZONTAL SHEAR STRESS. Tm, IN MEGAPASCALS 

FIGURE 10.12.-Maximum horizontal shear stress asafunctionof depth 

z. Solid wasion line has been fitted to all data, and dashed line to 

measurements in crystalline rock only. See figure 10.11 for locations 

of wells. 

has a greater surface intercept but a similar depth 

gradient to the San Andreaa regression 

fitted to the crystalline-rock data in figure 10.12. The 

comparison between equations 23 and 24 is not entirely 

appropriate because the data used to develop equation 23 

revresent all three stress states: stresses measured in 

regions of strike-slip tectonics were not considered separately by McGarr (1980) (see fig. 10.9). More recently, 

however, data measured in a 2,000-mdeep well in 

Cornwall, U.K. (Pine and others, 1983), permit quite an 

interesting comoarison. For both the San Andreas and 

the cornwall data sets, most of the stress observations 

are compatible with strike-slip tectonics; that is, u, is the 

intermediate principal stress. For most of the San 

Andreas and all of the Cornwall measurements, the rock 

is granitic. In contrast to the San Andreas system, 

however, the tectonic setting in Cornwall is presently 

inactive. The deviatoric stresses at Cornwall are believed 

to be a consequence of the Alpine orogeny, which 

apparently has caused the maximum horizontal stress to 

be oriented northwestward throughout much of Europe 

(for example, McGarr and Gay, 1978). 

The 12 data sets obtained by Pine and others (1988) 

indicate a stress state (fig. 10.14) surprisingly similar to 

that of the San Andreas fault (fig. 10.12). For the 

maximum shear stress, the depth gradient of 7.52 MPal 

krn is indistinguishable from its counterpart in the 

crystalline San Andreas crust of 7.46 MPaIkm, however, 

the surface intercept at Cornwall is larger. If the state of 

deviatoric stress is much the same in Cornwall as along 

the San Andreas system, then we must conclude that the 

plate-tectonic motion in California along the San Andreas 

fault has no expression in the shallow (1-2 km deep) 

stress field. Accordingly, much of what has been discovered about continental-crustal stress in general may 

apply to the crust adjacent to the San Andreas fault. 

This generalization implies that the applied forces 

which give rise to T,,, in the vicinity of the San Andreas 

fault are not specific to the Pacific-North American plate 

boundary. In terms of observed shear stress, a major 

active plate-boundary fault would be at least as likely in 

Cornwall, U.K., as in California from what we currently 

know of stress magnitudes, at depths down to a few 

kilometers. 

DISCUSSION 

From what we currently know of crustal stress and 

heat flow, neither is influenced by proximity to the San 

Andreas fault, the most conspicuous and best studied 
274 THE SAN ANDREAS FAULT SYSTEM, CALIFORNIA 

plate-boundary fault on the continents. The measured 

horizontal shear stress increases rapidly with depth 

(approx 8 MPa/km), essentially as would be predicted 

from laboratory measurements of friction and the assumption that crustal stress is limited by the frictional 

resistance of fractures forced together by the weight of 

overlying rocks. From this consistency of independent 

observations, two large "ifs" lead to what seems to be a 

physical contradiction: (1) if these vertical stress gradients persist throughout the depth of the seismogenic 

faulting layer (approx 12-15 km), then the average of the 

maximum horizontal shear stresses throughout the layer 

is quite large (approx 50 MPa); and (2) if the direction of 

the San Andreas fault is aligned with this maximumhorizontal-shear-stress direction, then the frictional heat 

generated by such stress during the documented fault 

motion (tens of kilometers per million years) should cause 

the background heat flow to double as the fault is 

approached. In 100 heat-flow measurements over a 

1,000-km span of the San Andreas fault, no such heatflow anomaly has been observed. 

The contradiction stems from two separate lines of 

argument: (1) inplace and laboratory measurements of 

rock stress imply average fault stresses of about 50 MPa 

or more, and (2) the absence of a local heat-flow anomaly 

and the energy balance of the fault imply an average fault 

stress of about 15 MPa or less. At least one of these 

arguments must be wrong. We have outlined the major 

factors in each argument, and we shall now point out 

some possible loopholes and areas for further study. 

The energy-balance argument leading to the heat-flow 

constraint on fault stress could be invalidated if the 

neglected energy sinks turn out to be important, or if the 

heat-conduction model is unrealistic or inappropriate. 

The general energy argument assumes that fault slip 

produces only seismic radiation and heat. It supposes 

that the energy consumed by the grinding of rocks into 

fault gouge (Lachenbmch and Sass, 1980) or the heat 

absorbed by possible phase changes or chemical reactions 

is negligible, and that the energy of seismic radiation 

does not grossly differ from the estimates made by 

seismologists. The heat-conduction model assumes that 

the frictional heat production occurs in a near-vertical 

fault zone (whose width is small relative to its depth) 

extending throughout the seismogenic layer. Systematic 

nonconductive removal of frictional heat by circulating 

ground water could invalidate this model (see O'Neill and 

Hanks, 1980; Williams and Narasimhan, 1989), as could a 

grossly different fault geometry-for example, a fault 

whose lower half was continually being rejuvenated 

because of migration of the upper half away from it along 

an upper-crustal detachment surface (Namson and Davis, 

1988). All of these effects probably deserve further 

study. 

The mechanical argument leading to large fault stress 

is based on observations of inplace stress (to maximum 

depths of approx 1 km in the San Andreas fault zone and 

of approx 4 km elsewhere on the continents), on laboratory measurements of rock friction and the efficiency of 

simulated earthquakes, and on downward extrapolation 

of these results through the seismic layer, on the 

assumption that fluid pressure is normal and frictional 

properties are uniform and isotropic. The consistency 

between the most frequently measured friction coeffidents and the inplace determinations of the vertical 

gradient of maximum shear stress is reassuring (solid 

curves, fig. 10.9; fig. 10.12); however, the downward 

extrapolation of these results to depths of 10 or 15 km is 

an uncertain step, with loopholes that could invalidate 

the high-fault-stress conclusion. 

0 Mojave Desen 

Central California 

-4 I I I I I I I I I I I I 

16 12 8 4 0 4 8 12 16 30 24 28 33 36 

DISTANCE FROM SAN ANDREAS FAULT, IN KILOMETERS 

FiGL-RR 10.18.-Averages of N maximum-shcar-etreas (in,) residuals as functionn of dktance from the San Andreas fault. In calculating residual 

at a particular distance, effect of depth 2 is removed by wing solid-line regression fit to data of figure 10.12 [T_ lUPal=1.15-8.282 (lanJl. 

Error bars, lo. 
10. STRESS AND HEAT FLOW 

There are at least three such loopholes. First, the fluid 

pressure might increase with depth, as it is known to do 

in some sedimentary basins, approaching the minimum 

principal stress (Berry, 1973). Second, the friction coefficient at depth might be lower than average laboratory 

values; such lower values have been reported in some 

studies of gouge and other clay-size aggregates (Wang 

and others, 1980). Each of these effects could substantially lower the maximum stress at depth. Third, the 

frictional strength properties might be anisotropic, with 

the main trace of the fault occupying a weak direction. If 

so, the maximum stress at depth might be high, as 

maintained in the mechanical argument, but the shear 

2.2 

0 5 10 15 20 25 

MAXIMUM HORIZONTAL SHEAR STRESS. Tm. IN MEGAPASCALS 

FIGURE 10.14.-Maximum horizontal shear stress T,,, as a function of 

depth 2 in granite near Cornwall. U.K. Data from Pine and others 

(1983). 

stress resolved on the fault might be very low, as 

maintained in the thermal argument. Under these cirlumstances, the fault must be nearly parallel to a 

principal axis, as suggested by Mount and Suppe (1987) 

and Zoback and others (1987). Such a condition could be 

consistent with the low friction on the main fault required 

by heat flow, while permitting high stresses to accommodate subsidiary faulting on more favorably situated 

planes with normal frictional properties. 

As mentioned above in the section entitled "Introdncion," geothermal studies of the San Andreas fault have 

provided evidence for a very weak fault for two decades; 

;he meaning of this result depends heavily on the 

iirection of principal stresses in the fault zone and the 

nagnitude of the stress differences there. As we have 

ihown, existing evidence is contradictory, especially in 

;he Mojave Desert region. Many measurements of stress 

rear the San Andreas fault suggest that the fault trace is 

nclined at an intermediate azimuth (approx 45") to the 

arincipal-streas directions, and thereby imply that the 

'ault coincides approximately with the direction of max- 

,mum shear stress and that the heat-flow constraint could 

lot be satisfied unless horizontal shear stress (and stress 

iiierences) were low everywhere. In this case, both the 

San Andreas fault and active subsidiary faults with other 

mentations would have to be weak. If, however, the 

lorizontal-stress differences are large, the weak fault 

quired by heat flow is a "zero shear stress" boundary 

wndition on the adjacent fault blocks that requires the 

hult to be almost normal to a principal-stress direction. 

[n this case, the heat-flow constraint could be honored on 

%n anomalously weak main trace, whereas subsidiary 

aults with other orientations and normal strength could 

ilso be active. 

Thus, the occurrence of a weak direction may reconcile 

Aservations of rock mechanics with the longstanding 

mplications of thennomechanical studies. It does, howiver, raise several questions: 

1. Does the maximum-horizontal-principal-stress direction form an intermediate angle with the trace of the 

San Andreas fault (as was formerly accepted and as is 

required by isotropic frictional properties), or is the 

maximum compression nearly fault normal, as suggested by more recent observations (Mount and 

Suppe, 1987; Zoback and others, 1987)T As we have 

pointed out, there is conflicting observational evidence on this issue. 

2. If the horizontal compressive stress is nearly faultnormal, what is the physical mechanism that permits 

the fault to slip under the small shear stress resolved 

on its direction? We have shown that the mechanism 

commonly invoked to explain a weak fault-namely, 

locally elevated fluid pressure-is not likely; however, anomalously low coefficients of friction could 
276 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

account for slip under near fault-normal compression 

if the frictional fault model is valid. 

3. How would such a weak plate-boundary fault evolve, 

and would its existence imply that the resistance to 

relative motion between the Pacific and North American plates is negligible? What is the role of deconpling and basal resistance? Lachenbruch and Sass 

(197% pointed out that strong shear stresses in the 

far field which mirht drive dextral slio between the 

plates cannot be balanced by a weakfault without 

invoking unexpected strength in the lower crust and 

drag (and possible decoupling or "detachment") beneath the horizontal base of the faulting layers. If the 

faulted plate boundary should weaken as it evolves, 

then either such basal drag must develop near the 

fault,or the far-field stress must diminish to maintain 

the equilibrium condition. The best way to learn 

whether such basal tractions exist is to determine 

whether the shear stress resolved parallel to the fault 

diminishes as the fault is approached (Lachenbruch 

and Sass, 1973; McGarr and others, 1982). We have 

shown here that a transect normal to the San 

Andreas fault shows no such diminution, although the 

observations were much shallower (approx 1 km deep 

or above) than the depth of earthquakes; direct stress 

measurements at seismogenic depths (below 5 km) 

are needed. Whether or not such basal decoupling 

and drag exist near the San Andreas fault is fundamental to our understanding of its earthquakes and of 

the nature of continental transform plate boundaries 

and their resistance to plate motion. 

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Grantz, Arthur, and Diekinson. W.R., 1968, Indicated cumulative 

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earthquakes: Annali di Geofisica, Rome, v. 9, no. 1, p. 1-15. 

Haimson. B.C., and Fairhmt, Charles, 1970, In situ strew determination at great depth by means of hydraulic fracturing, chap. 28 of 

Somerton, W.H., ed., Rock mechanics, theory and practice: 

Sympoiium on Rock Mechanics, llth, Berkeley. Calif., 1969, 

Prweedii, p. 659-584. 

Hanks, T.C., and Kanamori, Hiroo, 1W9, A moment magnitude scale: 

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Healv. . J.H.. and Zoback. M.D.. 1988. Hydraulic fracturing in aitu . 

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Henyey, T.L., 1968. Heat flow near major -trike-slip fault8 in central 

and southern California: Pasadena, California Institute of Technology, Ph.D. thesis, 421 p. 

Henyey, T.L., and Wasserburg, G.J., 1971, Heat flow near major 

strike-slip faults in California: Journal of Geophysical Research, v. 

76, no. 32, p. 7924-7946. 

Hickman, S.H., Zoback, M.D., and Healy, J.H., 1988, Continuation of 

a deep borehole stress measurement profile near the San Andreas 

fault, 1, Hydraulic fracturing stress measurements at Hi Vista, 

Mojave Desert, California; Journal ofGeophysical Research, v. 93, 

no. B12. D. 15188-151%. 

Habbert. M.K. and Willis, D.G., 1%7, Mechtnics of hydraulic frucluring: Journal of Petroleum Technology, v. 9, no. 6. p. 153-168. 

Jeger, J.C., 1956. Elasticity. fracture. and flow: New York, John 

Wiley and Sow, 152 p. 

Jones, L.M., 1988, Focal mechaniamn and the state of stress on the San 

Andrea fault in southern California: Journal of Geouhysical . - Research, v. 93, no. BE, p. 8869-8891. 

Kirby, S.H., 1980, Tectonic stresses in the lithosphere: Constraints 

provided by the experimental deformation of rocks: Journal of 

Geophysical Research, v. 85, no. Bll, p. 6S53-6868. 

Lachenbmch, A.H., 1986, Simple models for the estimation and 

measurement of frictional heating by an earthquake: U.S. Geological Survey Open-File Report 86-608, 13 p. 
10. STRESS AND HEAT FLOW 

Lachenhmch, A.H., and Sass, J.H., 1973, Thermomechmicalaspectsof 

the San Andreas fault system, in Kovach, R.L., and Nur, Amos, 

eds., Proceedings of the conference on tectonic problems of the 

San Andreas fault system: Stanford, Calif., Stanford University 

Publications in the Geological Sciences, v. 13, p. 192-205. - 1980, Heat flow and energetics of the San Andreas fault zone: 

Journal of Geophysical Research, v. 85, no. Bll, p. 6185-6222. - 1988, The stress heat-flow paradox and thermal results from 

Cajon Pass: Geophysical Research Letters, v. 15, no. 9, p. 

981-984. 

Lee, T.C., 1983, Heat flowthrough the SiliiJadntofault zone, southern 

California: Royal Astronomical SocietyGeophysical Journal, v. 72, 

no. 3, p. 721-731. 

Lockner, D.A., and Okubo, P.G., 1988, Memrements of frictional 

heatingm granitei Journalof Geophysical Research, v. 88, no. B5, 

p. 4313-4320. 

McGarr, Arthur, 1980, Some constraints on levels of shear stress in the 

crust from observations and theory: Journal of Geophysical 

Research, v. 85, no. Bll, p. 6231-6238. 

- 1988, On the state of lithmpheric street in the absence of applied 

tectonic forces: Journal ofGeophysical Research, v. 93, no. Bll, p. 

13609-13617. 

McGan, Arthur, and Gay. N.C., 1978. State of stress in the earth's 

crust: Annual Review of Earth and Planetary Sciences, v. 6, p. 

405-136. 

McGtrr, Arthur, Zoback, M.D., and Hanks, T.C., 1982, Implication! of 

an elastic analysis of in situ stress measurements near the San 

Andreas fault: Journal of Geophysical Research, v. 87, no. B9, p. 

7797-7806. 

McKenzie, D.P., 1969, The relation between fault plane solutions for 

earthquakes and the directions of the principal stresses: Seismological Society of America Bulletin, v. 59, no. 2, p. 591-601 . McKenzie, D.P. and Bmne, J.N., 1972, Melting on fault planes during 

large earthquakes: Royal Astronomical Society Geophysical Journal, v. 29, no. 1, p. 65-78. 

McNally, KC., Kanamori, Him, Peehmann, J.C., and Fills, G.S., 

1978, Earthquake swarm along the San Andreas fault near 

Palmdale, southern California, 1976 to 1977: Science, v. 201, no. 

4358, p. 814417. 

Minster, J.B., and Jordan, T.H., 1984, Vector constraints on Quaternary deformation of the western United States east and west of 

the San Andreas fault, in Crouch, J.K., and Bachman, S.B., eda., 

Tectonics and sedimentation along the California margin: Society 

of Economic Paleontologists and Mineralogists, Pacific Section 

field trip guide, v. 38, p. 1-16. 

Mount. V.S.. and Suone. . . John. . . 1987. State of stress near the San - 

Andreas fault; Implications for wrench tectonics: Geology, v. 15, 

no. 12, p. 1143-1146. 

Namson, J.S., and Davis, T.L., 1988, Seismically active fold and thrust 

belt in the San Jowuin Valley, central California: Geoloeieal 

Society of America ~ulletin, v. 100, no. 2. p. 257-273. - 

O'Neill. J.R.. and Hanks. Thomas. 1980. .~~~ Geochemical evidence for ~~ ~ ~~ ~~~ 

water.& interaction along the San Andreanand Garlock faults of 

California: Journal of Geophysical Research, v. 85, no. Bll. p. 

6286-6282. 

Oppenheimer, D.H., Reanenberg, P.A., and Simpson, R.W., 1988, 

Fault-plane solutions for the 1984 Morgan Hill, California, earthquake sequence: Evidence for the state of stress on the Calaveras 

fault: JournalofGeophysical Research, v. 93, no. B8, p. 9007-9026. 

Pine, R.J., Lediiham, Peter, and Mdeld, C.M., 1983, In-titu 

stress measurement in the Cammenellis granite-11. Hydrofracture tests at Rosemanowea Quarry to depths of 2000 m: Interne 

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Geomechanical Abstracts, v. 20, no. 2, p. 63-72. 

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California earthquake of April 18, 1906: Report of the State 

Earthquake Investigation Commission: Carnegie Institution of 

Washington Publication 87, 192 p. 

Ruina, A.L., 1983, Slip instability and state variable friction laws: 

Journal of Geophysical Research, V. 88, no. BIZ, p. 10359-10370. 

Savage, J.C., and Wood, M.D., 1971, The relationship between 

apparent stress and stress drop: Seismological Society of America 

Bulletin, v. 61, no. 5, p. 1381-1388. 

Sibson, R.H., 1985, A note on fault reactivation: Journal of Structural 

Geology, v. 7, no. 6, p.751-754. 

Spottiswoode, S.M., and McGarr, Arthur, 1975, Source parameters of 

tremors in a deep-level gold mine: Seiemological Society of 

America Bulletin, v. 65, no. 1, p. 93-112. 

Stesky, R.M., and Brace, W.F., 1973, Estimation of frictional stress on 

the San Andreaa fault from laboratory measurements, in Kovach, 

R.L., and Nur, Amos, eds., Proceedings of the conference on 

tectonic problems of the San Andreas fault system: Stanford, 

Calif., Stanford University Publications in the Geological Sciences, v. 13, p. 206-214. 

Stock, J.M., and Healy, J.H., 1988, Continuation of a deep borehole 

stress measurement profile near the San Andreas fault zone, 2, 

Hydraulic fracturingstress measurements at Black Butte, Mojave 

Desert, California: Journal of Geophysical Research, v. 93, no, 

B12, p. 15196-15206. 

Wang, C.-Y., Maa, N.-H., and Wu, F.T., 1980, Mechanical properties 

of clays at high pressure: Journal of Geophysical Research, v. 85, 

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Weldon, R.J., and Humphrien, E.D., 1986, A kinematic model of 

southern California: Tectonics, v. 5, no. 1, p. 3^-48. 

Weldon, R.J., and Sieh, K.E., 1985, Holocene rate of slip and tentative 

recurrenceinterval for large earthquakes on the San Andrean fault 

in Cajon Pass, southern California: Geological Society of America 

Bulletin, v. 96, no. 6, p. 793-812. 

Williams, C.F., and Naraaimhan, T.N., 1989, Hydrogeologic constraints on heat flow along the San Andreas fault-a testing of 

hypotheses: Earth and Planetary Science Letters, v. 92, p. 

131-143. 

Wyss, Visa, and Brune, J.N., 1968, Seismic moment, stress, and source 

dimensions for earthquakes in the California-Nevada region: 

Journal of Geophysical Research, v. 73, no. 14, p. 4681-4694. 

Wyss, Max, and Molnar, Peter, 1972, Efficiency, stress drop, apparent 

stress, effective stress, and frictional stress of Denver, Colorado, 

earthquakes: Journal of Geophysical Research, v. 77, no. 8, p. 

1433-1438. 

Zoback, M.D., and Haimson, B.C., eds., 1988, Hydraulic fracturing: 

stress measurements: Washington, D.C., National Academy 

Press, 270 p. 

Zoback, M.D., Healy, J.H., and Roller, J.C., 1977, Preliminary stress 

measurements in central California using the hydraulic fracturing 

technique: Pure and Applied Geophysics, v. 115, p. 135-152 

Zoback, M.D., Tsiitahara, Hiroaki, and Hickman, S.H., 1980, Stress 

measurements at depth in the vicinity of the San Andreas fault: 

Implications for the magnitude of shear stress at depth: Journal of 

Geophysical Research, v. 85, no. Bll, p. 6157-6173. 

Zoback, M.D.. Zoback, M.L., Mount, V.S., Suppa, John, Eaton, J.P., 

Healy. J.H.. Oppanheimer, D.H., Reaeenberg, P.A.,Jonea, L.M., 

Raleigh, C.B., Wong, I.G., Scotti, Oona, and Wentwerth, C.M., 

1987, New evidence on the state of stress ofthe San Andreas fault 

system: Science, v. 238. no. 4830, p. 1105-1111. 

Zoback. M.L., and Zohack. M.D., 1W. State of stress in the conterminous United States: Journal of Geophysical Research, v. 85, no 

Bll, p. 6113-6156. 

SUPPLEMENT: ADDITIONAL READING AND SOURCE MATERIAL 

ROBERT E. WALLACE, Compiler 

To help the reader find avenues into the voluminous 

literature about the San Andreas fault system, the 

following selected references are grouped into four sets: 

(1) maps, (2) review and reference publications, (3) 

puhlications of historical interest, and (4) publications of 

general interest. Many of these selected references have 

been included because they themselves have extensive 

bibliographies. The references in section 3 are limited to 

those pre-1965, except for Hill's review (1981) on the 

history of concepts. The references in section 4 are 

directed especially toward the layman or specialists in 

disciplines other than earth science. These suggested 

additional readings should be available in most earthscience libraries or from the publisher; many can also be 

found in larger public libraries. 

MAPS 

Bond, K.R., and Zietz, Isidore, 1987, Composite magnetic anomaly map of the conterminous United 

States: U.S. Geological Survey Geophysical Investigations Map GP-977, scale 1:2,500,000. 

Dibblee, T.W., Jr., 1973, Regional geologic map of San 

Andreas and related faults in Carrizo Plain, Temblor, 

Caliente, and La Panza Ranges and vicinity, California: U.S. Geological Survey Miscellaneous Geologic 

Investigations Map 1-757, 9 p., scale 1:125,000. 

Drummond, K.J., chairman, 1981, Pacific Basin sheet of 

Plate-tectonic map of the circum-Pacific region: Tulsa, OMa., American Association of Petroleum Geologists, scale 1:10,000,000. 

Engdahl, E.R., and Rinehart, W.A., 1988, Seismicity 

map of North America: Boulder, Colo., Geological 

Society of America: scale 1:5,000,000, 4 sheets. 

Geological Society of America, 1987, Gravity anomaly 

map of North America: Boulder, Colo., scale 

1:5,000,000, 4 sheets. 

- 1987, Magnetic anomaly map of North America: 

Boulder, Colo., scale 1:5,000,000, 4 sheets. 

Goter, S.K., 1988, Seismicity of California 1808-1987: 

U.S. Geological Survey Open-File Report 88-286, 

scale 1:1,000,000. 

Shaded relief map in color with earthquake epicenters in red. 

Jennings, C.W., compiler, 1958-66, Geologic atlas of 

California: Sacramento, California Division of Mines 

and Geology, scale 1:260,000, 27 sheets. 

This atlm contoins 27geologic map sheets that togetherco~titv,te 

the %do& Map of Califamin'' at a scale of l:25O,m. I&- 

motion sheets accompany each geologic map. A8 of 1989, the map 

aeries is being brought up to date, and the Santa Rosa Regional 

Geologic map series 2A (1982, 5 sheets) and San Bemardim 

Regional Geoloaic map series SA (1987, 5 sheets) have been 

completed. 

- compiler, 1975, Fault map of California with locations of volcanoes, thermal springs, and thermal 

wells: California Division of Mines and Geology Geologic Data Map 1, scale 1:750,000. 

Thisis the principal map representation offaults in Califamin on 

a single map sheet. It contains a wealth of information and is 

indispensable to anyone interested in the San Andreas fault. An 

updated edition is in preparation. 

- compiler, 1977, Geologic map of California: California Division of Mines and Geology Geologic Data Map 

2, scale 1:750,000. 

- 1985, An explanatory text to accompany the 

1:750,000 scale fault and geologic maps of California: 

California Division of Mines and Geology Bulletin 

201, 197 p. 

A significant review of data. For example, plate 2 includes four 

maps: (1) structural provinces of Calijamia, (2) pamllelism 

between major Quatemar# faults, (3) relation of eavtlupde 

epicenters to faults, and (A) earthquake epicenters. 

King, P.B., compiler, 1969, Tectonic map of North 

America: Washington, U.S. Geological Survey, scale 

1:5,000,000, 2 sheets. 

King, P.B., and Beikman, H.M., compilers, 1974, Geologic map of the United States (exclusive of Alaska 

and Hawaii): Washington, U.S. Geological Survey, 

scale 1:2,500,000, 3 sheets. 

Real, C.R., Toppozada, T.R., and Parke, D.L., 1978, 

Earthquake epicenter map of California: California 

Division of Mines and Geology Map sheet 39, scale 

1:1,000,000. A map and tabulation of earthquakes of 

279 
280 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

magnitude 5 and greater, 1900-74. 

Simpson, R.W., Hildenbrand, T.G., Godson, R.H., and 

Kane, M.F., 1987, Digital colored Bouguer gravity, 

free-air gravity, station location, and terrain maps 

for the conterminous United States: U.S. Geological 

Survey Geophysical Investigations Map GP-953-B, 

scale 1:7,500,000, 2 sheets. 

Simpson, R.W., Jachens, R.C., Saltus, R.W., and Blakely, R.J., 1986, Isostatic residual gravity, topographic, and first-vertical-derivative gravity maps of the 

conterminous United States: U.S. Geological Survey 

Geophysical Investigations Map GP-975, scale 

1:7,500,000,2 sheets. 

Stover, C.W., 1986, Seismicity map of the conterminous 

United States and adjacent areas, 1975-1984: U.S. 

Geological Survey &ophysical Investigations Map 

GP-984, scale 1:5,000,000. 

A standard reference depicting seismicity in the United States. 

REVIEW AND REFERENCE PUBLICATIONS 

Atwater, Tanya, 1970, Implications of plate tectonics for 

the Cenozoic tectonic evolution of western North 

America: Geoloeical Society of America Bulletin. v. 

81, no. 12, p. 3613-3536. 

- 

Develops the concept of the Pacific system of plates convew.ng 

with and undemding the North American plate. 

Crowell, J.C., ed., 1975, San Andreas fault in southern 

California: California Division of Mines and Geology 

Special Report 118, 272 p. 

Society of America Bulletin, v. 60, no. 4, p. 

1151-1197. 

Ernst, W.G., ed., 1981, The geotectonic development of 

California(Rubey volume I): Englewood Cliffs, N.J., 

Prentice-Hall, 706 p. 

Possibly the best general mview in one volume of the geologicand 

tectonics relations pertinent to the San Andreas fault syatern. 

More than 1,300 references listed in the back of the wiltme provide 

an inval*mMe source for further study. 

Hart, E.W., Hirschfeld, S.E., and Schuiz, S.S., eds., 

1982, Conference on Earthquake Hazards in the 

Eastern San Francisco Bay Area, Hayward, Calif., 

1982, Proceedings: California Division of Mines and 

Geology Special Publication 62, 447 p. 

A collection of 61 papers in which the Hayward and Calaveras 

faults, and other smaller hmhes of the San Andreas fault 

system, receive cmisidemble attention. 

Hinze, W.J., Kane, M.F., O'Hara, N.W., Redford, M.S., 

Tanner, James, and Weber, Christian, eds., 1985, 

The utility of regional gravity and magnetic anomaly 

maps: Tulsa, Okla., Society of Exploration Geophysicists, 454 P. 

A set of84 pa~providing infmtion and gu1'Aince on the use 

a4 interpre(atim ofgravity and magnetic-awmlg data,, as well 

0.8 regiowal maps. 

Kovach, R.L., and NUT, Amos, eds., 1973, Proceedings 

of the conference on tectonic problems of the San 

Andreas fault system: Stanford. Calif., Stanford 

University Publications in the Geological Sciences, v. 

13, 494 p. 

AH imprtant collection of 5S papers, representing the prevailing 

stateofknwrlntgeabout thesari Avdreo~ fault systemat thefiine 

Eaton, J.P., O'NeiU, M.E., and Murdock, J.N., 1970, 

Aftershocks of the 1966 Parkfield-Cholame, California, earthquake: A detailed study: Seismological 

Aguidebook containing 29 papers and a useful preliminary fault 

and geologic map of smthem California at a awle of 1:750,000. 

Dibblee, T.W., 1967, Areal geology of the western 

Mojave Desert, California: U.S. Geological Survey 

Professional Paper 522, 153 p. 

Domments many ofthe bedrock features that characterize the San 

~ndreas fault system. 

Dicldnson, W.R., and Grantz, Arthur, eds., 1968, Proceedinga of conference on geologic problems of San 

Andreas fault system: Stanford, Calif., Stanford 

University Publications in the Geological Sciences, v. 

11, 374 p. 

AÃ imiBi>rtBÃ§ ~o!i~!iton of 47~pers, rÃ§>re~t!ttn the prÃ§iÃ§iili 

iitiite of knowledge about the San Andrm fault systemat the tima 

of its publication. - 

An insight into the longer term history of the San Andreas fault 

system as indicated by the disruption of older rocks. 

Scholl, D.W., Grantz, Arthur, and Vedder, J.G., eds. 

of its publication. 

Oliver, H.W., 1980, Interpretation of the gravity map of 

California and its continental margin: California Division of Mines and Geology Bulletin 205, 52 p. 

A general reference and bibliographic source of grow.'ty data. 

Pakiser, L.C., Jr., and Mooney, W.D., eds., in press, 

Geophysical framework of the continental United 

States: Geological Society of America Memoir 172. 

A co~~ectim of 84 dxg a i,-t mm of ~i>t,j~ctÃ§ 

incisding atreas, tectonics, regional summmX, and mathads. 

Ross, D.C., 1984, Possible correlations of basement rocks 

across the Sari Andreas, Sari Gregorio-Hosgri and 

Rinconada-Reliz-King City faults, California: U.S. 

Geological Survey Professional Paper 1317, 37 p. 
SUPPLEMENT: ADDITIONAL READING AND SOURCE MATERIAL 

1987, Geology and resource potential of the contiiental margin of western North America and adjacent 

ocean basins-Beaufort Seato Baja California (Earth 

Science Series, v. 6): Houston, Tex., Circum-Pacific 

Council for Energy and Mmeral Resources. 

A major source durn abwd the continental m@n of North 

America. Includes 83 chapters concd with topics ravgiwf 

fawn tectonic and basin evolution to hydmear&on and metallicmineral occwrrence and potential. 

Sharp, R.V., 1967, San Jacinto fault zone in the Peninsular Ranges of southern California: Geological Society of America Bulletin, v. 78, no. 6, p. 705-729. 

Stewart, J.H., and Crowell, J.C., in press, Strike-slip 

tectonics in the Cordilleran region, western United 

States in Burchfiel, B.C., Libman, P.W., and Zoback, M.L., eds., The Cordilleran orogen: Boulder, 

Colo., Geological Society of America. 

Areview of strike-slipfa%Its, imlvding the Sun Andmas fault, in 

the Western United States. 

Streitz, Robert, and Sherburne. R.W., eds.. 1980, Studies of the San Andreas fault zone in northern california: California Division of Mines and Geology Special 

Report 140, 187 p. 

A valltdbit set of 16 payers on subjects iwlding geology, 

geophysics, seismology, and engineen'wf. 

Sylvester, A.G., 1988, Strike-slip faults: Geological Society of America Bulletin, v. 100, no. 11, p. 

1666-1703. 

An excellent review of the concerst. recmitian. mechanics. ami 

behamorof^fnke-slip faults, mth nuimvs refmvce~ to the Sow 

Andreaa fault Includes an extensive bibliography. 

Ziony, J.I., ed., 1985. Evaluating earthquake hazards in 

the Los Angeles region-an earth-science perspective: U.S. Geological Survey Professional Paper 1360, 

505 p. 

A vaJuable collection of I6 papers that review the state of the art 

about geologic hazards, rnw'ng from predicting ea-e 

faulting and pund motion to Invuhlides, tsummit, and the we 

of earth-sdence information. 

PUBLICATIONS OF HISTORICAL INTEREST 

Allen, C.R., 1967, San Andreas fault in San Gorgonio 

Pass, southern California: Geological Society of 

America Bulletin, v. 68, no. 3, p. 316-349. 

Anderson, P.M., 1899, The Geology of Point Reyes 

peninsula: Berkeley, University of California Publications, Department of Geology Bulletin, v. 2, no. 5, 

p. 119-153. 

Bailey, E.H., Irwin, W.P., and Jones, D.L., 1964, 

Franciscan and related rocks, and their significance 

in the geology of western California: California 

Division of Mines and Geology Bulletin 183, 177 p. 

A classic paper in which the San And- fault receives codmble aftention and early tectonic model8 are presented, 

Branner, J.C., 1906, An authoritative opinion: Mining 

and Scientific Press, v. 92, p. 347. 

Crowell, J.C., 1952, Probable large lateral displacement 

on the San Gabriel fault, southern California: American Association of Petroleum Geologists Bulletin, v. 

36, no. 10, p. 2026-2035. 

- 1962, Displacement along the San Andreas fault, 

California: Geological Society of America Special 

Paper 71, 59 p. 

Documents large-scale strike slip on the San Andreaa jault 

system. 

Eaton, J.E., 1939, Geology and oil possibilities of Caliente Range, Cuyama Valley and Carrizo Plain, California: California Journal of Mines and Geology, v. 35, 

no. 3, p. 255-274. 

Suggests 25 miles of strike slip on the San Andrew jault. 

Gilbert, G.K., 1907, The earthquake as a natural phenomenon, in The San Francisco earthquake and fire 

of April 18,1906, and their effects on structures and 

structural materials: U.S. Geological Survey Bulletin 

324, p. 1-13. 

Thispaperprovided the firstowmewofthe Son Awlreas fault m &hem Cdt/mio From Gillmi'# field notes, heappears In have 

been thefirst to aporfci01e Ike imf s~ff"t/VQnce of Ike fault 

Hill, M.L., 1981, San Andreas fault: History of concepts: 

Geological Society of America Bulletin, pt. 1, v. 92, 

no. 3, p. 112-131. 

A revim of the early development of comqts about the sign@- 

cam of the San And- fault, especially the cmnt of strike 

slip. 

Hill, M.L., and Dibblee, T.W., Jr., 1953, San Andreas, 

Garlock, and Big Pine faults, California: A study of 

the character, history, and tectonic significance of 

their displacements: Geological Society of America 

Bulletin, v. 64, no. 4, p. 44S-168. 

This classic paper was the first to provide data indicatiw of 

hundreds ofmiles of strike slip on the San And- fault. 

Hill, R.T., 1928, Southern California geology and Los 

Angeles earthquakes: Los Angeles, Southern California Academy of Sciences, 232 p. 

Lawson, A.C., 1695, Sketch of the geology of the San 

Francisco peninsula: U. S . Geological Survey Annual 

Report 15, p. 439-473. 
282 THE SAN ANDREAS FAULT SYSTEM. CALIFORNIA 

First use of the name "San Andmmfaidt." 

- chairman, 1908, The California earthquake of April 

18, 1906: Report of the State Earthquake Investigation Commission: Carnegie Institution of Washington 

Publication 87, 2 v. 

This treatise on the 1906 eecrthquake cmtatn.l an ahndance of 

imwnt data about the San Andm fault. 

Noble, L.F., 1926, The San Andreas rift and some other 

active faults in the desert region of southeastern 

California: Carnegie Institution of Washington Year 

Book 25, p. 415-422.. 

- 1933, Excursion to the San Andreas fault and Cajon 

Pass: International Geological Congress, 16th, Washington, 1933, Guidebook 15, p. 10-21. 

- 1964, The San Andreas fault zone from Soledad 

Pass to Cajon Pass, California, [pt I 5 oft Structural 

features, chap. 4 of Jahns, R.H., ed., Geology of 

southern California: California Division of Mines 

Bulletin 170, v. 1, p. 37-18. 

Oakeshott, G.B., 1959, The San Andreas fault in Marin 

and San Mateo Counties, in Oakeshott, G.B., ed., 

San Francisco earthquakes of March 1957: California 

Division of Mines Special Report 57, 127 p. 

Pack, R.W., 1920, The Sunset-Midway oil field, California, part 1, Geology and oil resources: U.S. Geological Survey Professional Paper 116, 179 p. 

Pack, R.W., and English, W.A., 1915, Geology and oil 

prospects of Waltham, Priest, Bitterwater, and 

Peachtree Valleys, California: U.S. Geological Survey Bulletin 681, p. 119-160. 

Reed, R.D., 1933, Geology of California: Tuba, OMa., 

American Association of Petroleum Geologists, 356 

P. 

- 1943, California's record in the geologic history of 

the world, in Geologic formations and economic 

development of the oil and gas fields of California: 

California Division of Mines Bulletin 118, p. 99-118. 

Does not mention strike slip on the San Andreas fault. 

Reed, R.D., and Hollister, J.S., 1936, Structural evolution of southern California: American Association of 

Petroleum Geologists Bulletin, v. 20, no. 12, p. 

1529-1704. 

Reid, H.F., 1910, The mechanics of the earthquake, v. 2 

of The California earthquake of April 18, 1906: 

Report of the State Earthquake Investigation Commission: Carnegie Institution of Washington Publication 87, 192 p. 

Develops elastic-rebound theory. 

Schuyler, J.D., 1896, Reservoirs for irrigation: U.S. 

Geological Survey Annual Report 18, p. 617-740. 

Steinb~gge, K.V., Zacher, E.G., Tocher, Don, Whitten, 

C.A., and Claire, C.N., 1960, Creep on the San 

Andreas fault: Seiimological Society of America 

Bulletin, v. 50, no. 3, p. 389-415. 

This paper was the first to report slow, continuo~is slip (creep) on 

ffie San And- fault. 

Taliaferro, N.L., 1938, San Andreas fault in central 

California [abs.]: Geological Society of America Proceedings, 1937, p. 254-255. 

Concludes that homontal slip on the San And- foul1 is less 

than a mile. 

Vaughan, F.E., 1922, Geology of San Bernardino Mountains north of San Gorgonio Pass: Berkeley, University of California Publications, Department of 

Geological Sciences Bulletin, v. 13, no. 9, p. 319-411. 

Vickery, F.P., 1925, The structural dynamics of the 

Livermore [California] region: Journal of Geology, v. 

33, no. 6, p. 608-628. 

Wallace, R.E., 1949, Structure of a portion of the San 

Andreas rift in southern California: Geological Society of America Bulletin, v. 60, no. 4, p. 781-806. 

Wilson, J.T., 1965, Anew dass offaults and their bearing 

on continental drift: Nature, v. 207, no. 4995, p. 

343-347. 

Develops the concept of a traytom fault. 

Wood, H.O., and Buwalda, J.P., 1930, Horizontal displacement along the San Andreas fault in Carrizo 

Plain, California [abs.] : Pan-American Geologist, v. 

54, no. 1, p. 75. 

PUBLICATIONS OF GENERAL INTEREST 

Dewey, J.F., 1972, Plate tectonics: Scientific American, 

v. 226, no. 5, p. 66-68. 

A brief iÂ¥ntradnetio to the concepts of phte tectonics, including 

excellent disamm of the geometry of the global plates. 

Iacopi, Robert, 1971, Earthquake country: Menio Park, 

Calif., Lane, 160 p. 

A review and guidebook about the San And- fault, ten for 

the layman, but imludes useful maps and photographs of interest 
SUPPLEMENT: ADDITIONAL READING AND SOURCE MATERIAL 

to the scientist and lapan alike. 

Jones, D.L., Cox, Allan, Coney, Peter, and Beck, Myrl, 

1982, The growth of western North America: Scientific American, v. 247, no. 5, p. 70-84. 

Refines the concept of plate tectonics has been refined to imIude 

the accretion of smaller blocks to the main cmtimntal masses. 

The San Andm8 fault swtemis involved with the mtvmS of bath 

larger and smatlerplans, 

Jordan, T.H., and Minster, J.B., 1988, Measuringcrustal 

deformation in the American West: Scientific American, v. 269, no. 2, p. 48-58. 

Prfmidss ~fmif. ideae arid dnta aht funowestern Nmth America 

is being deformed. 

Schuiz, S.S., and Wallace, R.E., 1987, The San Andreas 

fault: Washington, U.S. Geological Survey, 17 p. 

Pamphlet providing a brief review for the lapam. 

Wesson, R.L., and Wallace, R.E., 1985, Predicting the 

next great earthquake in California: Scientific American, v. 252, no. 2, p. 35-13. 

The science of earUmmke prediction has been evolving, and the 

Saw Andreaa fault system has beena taqpztof majorintereat. This 

paper mt6ev.s some progress that has been made. 

ft U. S. GOVERNMENTPRINTINGOFFICE: 1991-We-$33 

7 EXPLANATION 

I Segment on which slip occurred during 

great earthquakes of 1857.1872. and 

1 

1906Ã‘Arrow indicate direction of 

relative movement 

t Segment on which slip occurred during 

smaller earthquake 

Segment on which fault creep occul 

SAN ANDREAS FAULT SYSTEM AND OTHER LARGE FAULTS IN CALIFORNIA 

Different segments of fault display different behavior.