Earthquake Hazards in 

the Offshore Environment 

... 
• 
Earthquake Hazards in 

the Offshore Environment 

By ROBERT A. PAGE and PETER W. BASHAM 

U.S. GEOLOGICAL SURVEY BULLETIN 1630 
DEPARTMENT OF THE INTERIOR 

DONALD PAUL HODEL, Secretary 

U.S. GEOLOGICAL SURVEY 

Dallas L. Peck, Director 

UNITED STATES GOVERNMENT PRINTING OFFICE 1985 

For sale by the 

Superintendent of Documents 

U.S. Government Printing Office 

Washington, D.C. 20402 

library of Congress Cataloging-in-Publication Data 

Page, Robert Alan, 1938-

Earthquake hazards in the offshore environment. 

(U.S. Geological Survey Bulletin 1630) 

Bibliography: p. 56-67 

Supt. of Docs. no.: 119.3:1630 

1. Earthquakes-United States. 2. Earthquakes-Canada. 

3. Earthquake engineering-United States. 4. Earthquake 

engineering-Canada. 5. Oil-well drilling, submarineUnited States. 6. Oil-well drilling, submarine-Canada. I. 

Basham, P.W. II. Title. Ill. Series: United States. Geological Survey. Bulletin 1630. 

QE75.B9 no. 1630 557.3s 85-600173 

[QE535.2.U6] [363.3'495] 
CONTENTS 

Abstract 1 

Introduction 2 

Offshore earthquakes of the United States and Canada 2 

Pacific margin 2 

Atlantic and Arctic margins 6 

Earthquake effects 8 

Tectonic deformation 8 

Surface faulting 8 

Subsidence and uplift 1 0 

Seismic shaking 11 

Sea-floor failures 14 

Turbidity currents 18 

Tsunamis 19 

Local tsunamis 19 

Distant tsunamis 2 0 

Assessment of earthquake potential 21 

Seismic record 22 

Historical seismicity 22 

Location of offshore earthquakes 2 2 

Ocean-bottom seismographs 24 

Geologic record 24 

Seismotectonic models 2 9 

Active western margin 30 

Passive eastern margin 31 

Induced seismicity 32 

Evaluation and mitigation of earthquake hazards 35 

Surface faulting 3 5 

Subsidence and uplift 3 7 

Seismic shaking 39 

Site-specific evaluation 41 

Areal evaluation 4 3 

Construction of probabilistic ground-shaking maps 44 

Western Canada 46 

Gulf of Alaska 4 6 

American Petroleum Institute guidelines 47 

Sea-floor failures 4 9 

Areal evaluation 51 

Site-specific evaluation 51 

Turbidity currents 54 

Tsunamis 55 

Conclusion 56 

References cited 56 

Intensity, magnitude, and seismic moment 67 

FIGURES 

1. Map showing epicenters of large historical earthquakes of the United 

States and Canada 3 

2. Index map of the United States and Canada 4 

3. Index map of California 5 

4. Index map of southern Alaska and the Gulf of Alaska 7 

5. Photograph showing fence offset by faulting during the 1906 San Francisco 

earthquake 8 

6. Photograph showing bedrock scarp formed by faulting during the 1964 Alaska 

earthquake 9 

7. Photograph showing section of gas-transmission line damaged by faulting in 

the 1971 San Fernando, Calif., earthquake 10 

III 
8. l'viap of Gulf of Alaska area, showing tectonic uplift and subsidence during 

the 1964 Alaska earthquake 11 

9. Photograph showing former sea floor around Montague Island, exposed by 

tectonic uplift during the 1964 Alaska earthquake 12 

10. Photograph of flooding in Seldovia, Alaska, caused by tectonic subsidence 

during the 1964 earthquake 12 

11. Strong-motion recordings from the 1978 Santa Barbara, Calif., earthquake 13 

12. Aerial photographs of Seward, Alaska, waterfront before and after the 1964 

earthquake 15 

13. Seismic-reflection record showing slumps in foreset beds of the Copper 

River delta, Alaska 16 

14. Seismic-reflection records showing sediment failure on the Continental 

Shelf induced by the 1980 Eureka, Calif., earthquake 17 

15. Map of submarine slump and turbidity current triggered by the 1929 Grand 

Banks earthquake off eastern Canada 18 

16. Photograph showing waterfront damage in Seward, Alaska, from slides, 

waves, and fire caused by the 1964 earthquake 2 0 

17. Plots of sea-level fluctuations on Kodiak Island from the principal 

tsunami generated by the 1964 Alaska earthquake 21 

18. Maps showing locations of epicenters off central California, determined by 

different procedures 2 3 

19. Maps showing seismicity offshore of southern British Columbia, Canada 25 

20. Seismic-reflection profile across the Nootka fault zone off Vancouver 

Island, British Columbia, Canada 27 

21. Photograph and map showing offset stream channels along the San Andreas 

fault, central California 2 8 

22. Shallow geologic section across a fault that slipped in the 1971 San 

Fernando, Calif., earthquake 2 9 

23. Plot of earthquake-magnitude-recurrence relations for two earthquakesource zones off British Columbia, Canada 31 

24. Map of the Gulf of Alaska area, showing the Yakataga and Shumagin seismic 

gaps 32 

25. Map showing earthquake epicenters and basement faults on the Continental 

Shelf south of Newfoundland 33 

26. Bar graph of earthquake activity and reservoir pressures at Rangely, 

Colo., oil field 34 

27. Plot of fault-creep displacement measured on the San Andreas fault, 

central California 35 

28. Plots of fault displacement along the Imperial fault, southern California, 

as a function of time after the 1979 Imperial Valley earthquake 36 

29. Map of the Imperial Valley area, southern California, showing surface 

fault ruptures accompanying the 1979 Imperial Valley earthquake 37 

30. Plots of maximum surface fault displacement as a function of earthquake 

magnitude 38 

31. Plots of peak horizontal ground acceleration and velocity versus distance 

from the fault as a function of earthquake magnitude 42 

32. Velocity-response spectra for seismic surface waves on typical onshore and 

offshore geologic structures 4 3 

33. Theoretical sea-floor seismograms for a representative Continental Shelf 

structure 44 

34. Schematic illustration of steps in method for the probabilistic estimation 

of ground motion 45 

35. Probabilistic maps of peak horizontal acceleration and velocity in western 

British Columbia, southeastern Alaska, and adjacent offshore region 47 

36. Probabilistic maps of peak horizontal acceleration for the Gulf of Alaska 47 

37. Maps showing earthquake-risk zones for the design of fixed offshore 

platforms in U.S. coastal waters 48 

38. Seismic profile of submarine slide in the Kayak Trough, northern Gulf of 

Alaska 49 

39. Records of stress, strain, and pore-water pressure from laboratory shear 

tests on sand 50 

40. Relative-slope-stability map of part of the Continental Shelf south of 

Kodiak Island, Alaska 53 

41. Probabilistic map of tsunami elevation for the contiguous United States 

and Alaska 56 

IV 
TABLES 

1. Submarine slope failures attributed at least in part to seismic shaking 52 

2. Classification of relative slope stability, Kodiak shelf area, Alaska 54 

3. Magnitudes and moments for earthquakes discussed in this report 69 

v 

Earthqulake Hazards in the Offshore Environment 

By Robert A. Page and Peter W. Basham1 

Abstract 

This report discusses earthquake effects and 

potential hazards in the marine environment, describes 

and illustrates methods for the evaluation of earthquake hazards, and briefly reviews strategies for 

mitigating hazards. The report is broadly directed 

toward engineers, scientists, and others engaged in 

developing offshore resources. 

The continental shelves have become a major 

frontier in the search for new petroleum resources. 

Much of the current exploration is in areas of moderate to high earthquake activity. If the resources in 

these areas are to be developed economically and 

safely, potential earthquake hazards must be identified 

and mitigated both in planning and regulating activities and in designing, constructing, and operating 

facilities. 

Geologic earthquake effects that can be hazardous to marine facilities and operations include 

surface faulting, tectonic uplift and subsidence, 

seismic shaking, sea-floor failures, turbidity currents, 

and tsunamis. Seismic shaking typically contributes 

most to earthquake losses, both directly through 

vibratory damage and indirectly through triggering of 

landslides and loss-of-strength failures in sedimentary 

deposits, which, in turn, may cause local tsunamis and 

turbidity currents. Examples of damage to oceanbottom and coastal facilities from sea-floor failures, 

slide-generated tsunamis, and turbidity currents are 

numerous. Uplift, subsidence, and tsunamis are likely 

to be serious hazards only in coastal areas, where large 

water waves can wreck shoreline facilities and uplift 

or subsidence can destroy or limit the utility of such 

facilities. Surface faulting affects only a limited area; 

it can rupture pipelines and offset or distort foundations, however, and thus cause structures to fail. 

Strategies for mitigating earthquake hazards rely 

on our ability to predict the location and size of future 

potentially damaging earthquakes, their imminence 

and frequency of occurrence, and the type, severity, 

extent, and likelihood of resulting geologic effects. 

1Earth Physics Branch, Canadian Department of 

Energy, Mines and Resources, Ottawa, Ontario, 

Canada. 

Rather precise estimates of the location and size of 

future earthquakes can be provided for many seismic 

zones, principally those with high rates of activity 

where the causes of earthquakes are relatively well 

understood, such as along the Pacific margin of the 

United States and Canada. Although, in such areas, 

satisfactory estimates of the frequency of earthquake 

occurrence can commonly be derived, routine prediction of the precise time of future shocks is currently 

beyond the state of the art. In less seismically active 

regions, such as the Atlantic and Arctic margins of the 

United States and Canada, estimates of the location, 

size, and frequency of future earthquakes are difficult 

and less precise and reliable because the causes of 

earthquakes in areas distant from plate boundaries are 

still poorly understood. 

The general absence of detailed documentation 

and measurement of earthquake effects in the offshore 

environment seriously constrains our ability to identify 

and mitigate potential hazards. The hazards posed by 

some earthquake effects are obvious, and in some 

places the methods and techniques for evaluating 

earthquake hazards on land can be applied with little 

or no modification in the marine environment. For 

example, experience with surface faulting on land is 

directly applicable to assessing the potential for and 

style of surface faulting in offshore earthquakes. For 

other earthquake effects, differences between the 

offshore and onshore environment are important and 

limit the applicability of experience with earthquake 

effects on land. For strong seismic shaking and seafloor failure, ignorance of the onsite geotechnical 

properties of saturated marine sediment and its 

response to seismic disturbances is a serious impediment. Finally, there are earthquake phenomena 

peculiar to the marine environment whose or1gms are 

poorly understood and whose significance as hazards 

are neither known nor appreciated. 

Mitigation strategies for earthquake hazards 

include selectively siting facilities and limiting activities in hazardous areas to reduce exposure to hazards, 

designing and constructing facilities to withstand or 

accommodate expected earthquake effects, and instituting response plans for earthquake emergencies and 

for advance warnings of delayed earthquake effects, 

such as tsunamis. The effectiveness of any particular 

strategy depends on both the hazard being addressed 

and the type of facility or activity at risk. 

Introduction 1 
INTRODUCTION 

In the last decade offshore petroleum exploration 

and development have greatly increased. Much of the 

recent and planned economic activity is in areas of 

moderate to high earthquake potential. Safe exploration and eventual development of petroleum resources 

in such areas require that offshore and coastal facilities be designed to resist earthquakes and their 

geologic and hydrologic effects. 

The purpose of this report is to promote an 

understanding among scientists, engineers, and others 

engaged in developing offshore resources of earthquakes and their effects, as related to the marine 

environment, and to provide an introduction to the 

methods for assessing earthquake potential and for 

evaluating and mitigating earthquake hazards. 

Although the techniques for evaluating earthquake 

potential and associated hazards are relevant to any 

offshore or coastal site worldwide, the discussion here 

is illustrated primarily with examples from the United 

States and Canada. Readers particularly interested in 

earthquake hazards in the North Sea area are referred 

to Ritsema and Gl1rpinar (1983). Because the offshore 

region is a new frontier, few well-documented 

examples of earthquake effects in the marine environment are available, and so frequent reference to 

experience with earthquakes on land is necessary. 

The potential for petroleum exists at some scale 

in essentially all unmetamorphosed sedimentary 

basins. The geologic history of the Phanerozoic Eon 

(the past 500-600 m.y.), which has controlled basin 

formation, is influenced by the major horizontal 

crustal movements associated with global tectonics. 

Earthquakes are known to be occurring today in most 

regions that have undergone orogenic events or reactivation of older geologic structures during the Phanerozoic Eon. Much earthquake activity is occurring at or 

near continental margins and thus in close proximity to 

areas of petroleum development. 

Acknowledgments.-ln addressing the broad 

subject of earthquake hazards in the offshore environment, we have benefited from informative discussions 

with numerous colleagues and from their comments on 

and criticism of the manuscript. In particular, we 

recognize the contributions of D.M. Boore, M.G. 

Bonilla, M.A. Hampton, H.J. Lee, D.S. McCulloch, T.L. 

Youd, and J .1. Ziony of the U.S. Geological Survey, and 

M.J. Berry and R.D. Hyndman of the Earth Physics 

Branch of the Canadian Department of Energy, Mines 

and Resoures. 

OFFSHORE EARTHQUAKES OF THE UNITED 

STATES AND CANADA 

Epicenters of the larger historical earthquakes of 

the United States and Canada are shown in figure 1. 

West of long 110° W., 2

the earthquakes shown are 

restricted to magnitudes (M) of at least 6; however, 

2Magnitude is a measure of the size of an earthquake based on the amplitude of seismic waves. This 

measure is logarithmic rather than linear; a difference 

2 Earthquake Hazards in the Offshore Environment 

the epicenter patterns are schematic because all 

significant events cannot be depicted at this scale. 

East of long 110° W., M=5 earthquakes after 1930 are 

included to more clearly illustrate the patterns of the 

lower level of seismicity in the eastern part of the 

continent. A large number of significant historical 

earthquakes have occurred on the continental margin 

(fig. 1), or close enough to that margin to have caused 

significant effects offshore. The 2,000-m bathymetric 

contour marks the approximate position of the edge of 

the continental margin and delimits the potential 

offshore petroleum-resource areas that have been 

identified in coastal and offshore unmetamorphosed 

sedimentary basins (for example, McCrossan and 

Porter, 1973.) 

Pacific margin 

The greatest concentration of earthquakes in 

North America is along the west coast, where several 

earthquakes of M ~ 7. 7 have occurred during historical 

time. These earthquakes reflect brittle deformation 

occurring in response to contemporary movement 

between two large lithospheric plates, the Pacific and 

the North American (fig. 2). The local sense of 

relative movement between these two plates varies 

along the length of the boundary. Along the coasts of 

California, northern British Columbia, and southeastern Alaska, the plates are slipping horizontally 

past each other; the Pacific plate is moving northnorthwestward relative to the North American plate. 

In contrast, the relative plate motion is convergent 

along the Alaska Peninsula and the eastern Aleutian 

Islands. 

The style of faulting along the plate boundary 

changes in response to the local direction of relative 

plate motion. Right-lateral strike-slip faulting is 

dominant along the coasts of California, northern 

British Columbia, and southeastern Alaska. Major 

strike-slip faults that have generated great earthquakes include the San Andreas fault in California 

(1857, M=7.9; 1906, M=7.7), the Queen Charlotte fault 

off British Columbia (1949, M=8.1), and the 

of one unit of magnitude corresponds to approximately 

a factor of 30 in the energy radiated from the source 

as seismic waves. Seismologists use several methods, 

or scales, for calculating magnitudes from 

instrumental records and for estimating the 

magnitudes of earthquakes not instrumentally recorded 

from the effects of ground shaking on structures, 

people, and natural features. Throughout this report, 

we use the simple term "magnitude" without indication 

of the method used for its computation. This usage is 

adequate for a general discussion of earthquake 

hazards; however, in applying the methods of hazards 

assessment and mitigation discussed in this report, it is 

essential that the practitioner be aware of the differences among various magnitude scales and their implications concerning the source characteristics of earthquakes (see supplementary section below entitled 

"Intensity, Magnitude, and Seismic Moment"). The 

magnitude values for specific earthquakes cited in this 

report refer to the moment-magnitude scale. 
60' 70° 80' 80° 10° 60' 

EXPLANATION 

EARTHQUAKE 

EPICENTERS 

0 M2:8 

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--L_._ l------------+--------1- -- ---- : 

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c:::d.~~ll"""' 

120° 110° 10[]4l goa 

1000 2000 KILOMETERS 

Figure 1. Epicenters of large (M ~6) historical earthquakes in and adjacent to the continental United States and 

Canada. Larger magnitude symbols obscure numerous historical events of smaller magnitude, particularly along 

western margin. Historical (M=S) earthquakes after 1930 are also included for region east of long 110° W. Z,OOOm bathymetric contour marks approximate location of edge of continental margin. 

Fairweather fault in southeastern Alaska (1958, 

M=7. 7). Along the Alaska Peninsula and the eastern 

Aleutian Islands, where the Pacific plate is being 

subducted beneath the North American plate, dip-slip 

faulting predominates. The largest earthquakes in 

North America have occurred along this subduction 

zone. Most of the Aleutian subduction zone has been 

ruptured during the past 50 years in a series of four 

great earthquakes: the Alaska Peninsula (1938, 

M=8.2), the Andreanof Islands (1957, M=9.1), the 

Prince William Sound (1964, M=9.2), and the Rat 

Islands (1965, M=8.7). -

There are many significant complexities in this 

gross view of the relation between plate tectonics and 

seismicity along the west coast of North America. 

Not all the transcurrent motion between the Pacific 

and North American plates in California is accommodated on the San Andreas fault (fig. 3); part of this 

motion is distributed among several subparallel faults 

that define a broad zone of active strike-slip faulting, 

approximately 100 km wide, which includes the San 

Andreas (Jennings, 1975). South of San Francisco, 

where the San Andreas fault is onshore, some of these 

subparallel faults lie offshore and pose the principal 

Offshore Earthquakes of the United States and Canada 3 
CIFIC 

LATE 

80' 90" 80' 7[)' 

UNITED STATES 

130' 120° 110' 100' 90' 80° 70° 

1000 2000 KILOMETERS 

Figure 2. Index map of the United States and Canada, showing tectonic plates, and places and features referred 

to in text. 

earthquake hazard at marine sites. A complexity in 

southern California is the bend in the San Andreas 

fault in the Transverse Ranges, north of Los Angeles 

(fig. 3). This bend is a westward deflection of the 

northwesterly striking fault and is flanked by a broad 

east-west zone of compressional tectonics that 

extends offshore and includes the Santa Barbara 

Channel (Yerkes and Lee, 1979a). East-west-trending 

active folds and reverse faults characterize this zone 

and accommodate north-south shortening of crustal 

rocks. Active faults off southern California have 

generated historical earthquakes as large as M=7 .3. 

Between Cape Mendocino in northern California 

and the north end of Vancouver Island, the Pacific and 

North American plates do not meet along a common 

boundary; instead, they are separated by smaller plates 

of oceanic lithosphere-the Juan de Fuca and Explorer 

4 Earthquake Hazards in the Offshore Environment 

plates (fig. 2)-that are being subducted beneath the 

North American continent (Silver, 1971; Keen and 

Hyndman, 1979; Smith and Knapp, 1980). The tectonic 

complexities associated with the intervening plates are 

only now being resolved. Along the continental margin 

at these latitudes, earthquake mechanisms indicate 

brittle deformation between plates, within the oceanic 

and underthrust segments of the subducting plates, and 

within the overlying continental lithosphere. The 

largest historic earthquake along the continental 

margin is M=7 .3; however, the possibility of a major 

(possibly aSiarge as M=8.3) subduction shock along the 

Juan de Fuca-North American plate boundary has been 

suggested (Weaver and Smith, 1983; Heaton and 

Kanamori, 1984). Seaward of the continental margin 

to distances of about 400 km, earthquakes as large as 

M=6.5 delineate the boundary of the Pacific plate, 
42° 

40° 

38° 

36° 

34° 

32° 

124' 122° 120° 

OREGON 

118° 

I 

I 

116° 114° 

--------- j _________ l Crescent City ----~------------

Klamath River 1 \ 

Humboldt BayGa~~ 

MendocinoPoint Arenai I 

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Santa Barbara 

SANTA BARBARA CHANNEL-

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200 

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400 KILOMETERS 

Figure 3. Index map of California, showing places and features referred to in text. 

Offshore Earthquakes of the United States and Canada 5 
which is composed of a sequence of spreading oceanic 

ridges and transform fracture zones. 

The continental margin along the northeastern 

Gulf of Alaska between Cross Sound and Prince 

William Sound (fig. 4) is also characterized by complex 

tectonics. This is a region of transitional tectonics 

from transcurrent faulting along the coast of southeastern Alaska and British Columbia to subduction 

along the Aleutian Island arc. Smaller plates, whose 

geologic histories and current tectonic significance are 

only partly understood at this time, separate the 

Pacific and North American plates and introduce 

additional boundaries on which tectonic adjustments 

may occur to generate significant earthquakes (Lahr 

and Plafker, 1980; Perez and Jacob, 1980). The capability of this segment of the continental margin to 

generate significant earthquakes was demonstrated by 

a pair of great (M=8.1) onshore earthquakes in 1899 

that ruptured the coast between Yakutat Bay and 

Kayak Island. 

Atlantic and Arctic margins 

Although the continental margins of eastern and 

Arctic North America (fig. 2) are far less seismically 

active than the Pacific margin, they have been shaken 

by several significant earthquakes in historical time. 

The Atlantic and Arctic margins were formed by 

rifting and transform (shear) motion between 

continental masses that separated to form new ocean 

basins. The evolution of rifted margins includes an 

early rifting stage, during which the crust is uplifted 

because of heating and thermal expansion in the 

mantle, followed by a spreading stage, in which the 

continents move apart as sea floor is created between 

them. During the spreading stage, the margin subsides 

as the lithosphere cools and sediment is deposited on 

the subsiding basement. The age of the Atlantic 

margin, defined by its time of the first opening, ranges 

from Jurassic for the Eastern United States and 

Maritime Canada, where North America separated 

from Africa, Iberia, and the British Isles, to Late 

Cretaceous and Tertiary off Labrador and Baffin 

Island, where Greenland separated from North 

America. The Arctic Ocean continental margin was 

formed during Cretaceous time by the separation of 

Alaska from the Canadian Arctic islands (Keen and 

Hyndman, 1979; Sweeney and others, 1978). 

The eastern and Arctic margins are passive; the 

nearest active tectonism currently occurs far to the 

east at the Mid-Atlantic Ridge or, in the Arctic 

Ocean, on the Nansen-Gakkel Ridge, which extends 

from near northeastern Greenland to near the 

Novasibirskye Islands. Nevertheless, earthquake 

activity is occurring, albeit at lower rates and with 

generally smaller magnitudes than on the active 

margin of western North America. Among the largest 

historical earthquakes are the following: the Gulf of 

Maine off Cape Anne, Mass. (1775, estimated M=5.8-

6.1), the Charleston, S.C. (1886, estimated M=6.9-7 .3), 

the Beaufort Sea off Banks Island (1920, M=6.2), the 

Grand Banks (1929, M=6.5), and the BaffinBay (1933, 

M=6. 7). It is generally assumed that most earthquakes 

along the eastern margin are occurring in the old 

6 Earthquake Hazards in the Offshore Environment 

rifted continental lithosphere. However, too few 

earthquakes have been accurately located during the 

era of instrumental seismology to delineate any significant patterns. Furthermore, the geologic features of 

the margin are too poorly mapped, and the locations of 

the more significant historical shocks (fig. 1) too 

uncertain, to support definitive correlations of epicenters with geologic structure. 

The origins of these passive-margin earthquakes 

are not well understood, and several models for their 

origins have been proposed. One aspect common to 

many of these models is that the earthquakes are 

occurring in ancient zones of weakness within the 

crust that have been reactivated by the current stress 

regime (Sykes, 1978; Zoback and Zoback, 1981). The 

notion is that ancient zones of weakness can be the 

site of modern fault movement if they are suitably 

oriented in the current stress field and if the shear 

stress on the fault plane is sufficiently large. For 

example, Zoback and Zoback (1981) concluded that the 

Atlantic coast of the United States is under northwestward-oriented compression, on the basis of hydrofracture stress measurements, earthquake focal mechanisms, and late Cenozoic fault offsets. Wentworth 

and Mergner-Keefer (1981) argued that this 

compression causes reverse slip on ancient northeasttrending faults and that many of these faults 

preexisted as normal faults in early Mesozoic time, 

during the initial rifting to form the Atlantic Ocean. 

They suggested that reverse faulting, which continues 

at an average rate of a fraction of a meter per million 

years, can account for much of the historically 

observed seismicity along the eastern seaboard, including the 1886 Charleston, S.C., earthquake. Seeber and 

Armbruster (1981) offered an alternative model for the 

Charleston earthquake in which another preexisting 

weakness is reactivated. In their model, the proposed 

mechanism is backslip on the subhorizontal fault or 

detachment surface underlying the Appalachian 

Mountains; this surface was formed in the continentalplate collision that preceded the opening of the 

present Atlantic Ocean. Other possible mechanisms 

for the localization of seismicity in the Eastern United 

States were reviewed by Zoback and Zoback (1981). 

Numerous explanations have been offered for the 

stresses responsible for earthquakes that occur away 

from plate boundaries. Although numerical models of 

the driving mechanisms of lithospheric plates seem to 

explain some of the gross features of stress orientations observed in intraplate regions (Richardson and 

others, 1979), other geologic processes, such as loading 

and unloading of the crust, are also important. From a 

study of the focal mechanisms of earthquakes on the 

margin of eastern Canada, Stein and others (1979) 

argued that stresses due to the removal of Pleistocene 

glaciers are sufficient to reactivate old basement 

faults parallel to the continental margin. Hasegawa 

and others (1979) inferred the stress regime of the 

Canadian Arctic margin along the continental slope in 

the Beaufort Sea to be a combination of crustal 

loading from Quaternary sedimentation, remnant 

tectonic stress associated with the opening of the 

Arctic Ocean during the Early Cretaceous, and 

tectonic stress generated currently by ocean spreading 

at the Nansen-Gakkel Ridge and transmitted across 

the Arctic Ocean. 
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ALASKA I 

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BANK 

~odiak 

Womens 

Boy 

61° .__ __ _L_ ___ .:::::::;:h_ ______ _ 

200 400 KILOMETERS 

Figure 4. Index map of southern Alaska and the Gulf of Alaska, showing places and features referred to in text. Inset shows Port Valdez region. 
Thus, the passive margin of eastern and Arctic 

North America has undergone significant 

earthquakes. Unlike the situation for shocks along the 

Pacific margin, no single unifying tectonic model is 

recognized that explains the origin of these 

earthquakes. Different processes, in fact, are 

probably generating earthquakes in different regions. 

Uncertainties in regard to the causes of historical 

shocks limit our ability to predict the location and 

frequency of future similar-size events in eastern and 

Arctic North America, and thus present a particular 

difficulty to the evaluation of earthquake hazards 

along this margin. 

EARTHQUAKE EFFECTS 

The various earthquake effects that can pose 

hazards to offshore facilities range from primary 

effects, such as tectonic deformation and strong 

ground shaking, to secondary effects, such as failure of 

geologic materials, turbidity currents, and tsunamis. 

Primary effects result directly from the sudden 

displacement of one rock mass past another along a 

buried fault, the mechanism that generates an earthquake; secondary effects are those induced by shaking 

or tectonic deformation. 

The ranking of earthquake effects in terms of 

degree of potential hazard, or in terms of importance 

to the design engineer, is possible only in a general 

sense. The order will vary from one earthquake to 

another, depending on the size and location of the 

earthquake, and from site to site, depending on 

distance from the earthquake and on local geologic 

conditions. Seismic shaking and the resulting geologic 

effects generally constitute the greatest overall 

potential earthquake hazards to offshore facilities, 

including both platforms and pipelines. Tsunamis and 

tectonic changes in elevation are typically of little 

consequence to offshore facilities in comparison, for 

example, with the loads imposed by storm waves. For 

coastal facilities related to offshore petroleum 

production, however, tsunamis and uplift or subsidence 

of the coastline can constitute significant hazards. 

Measures to accommodate or mitigate 

earthquake hazards, if they are to be effective, must 

be based on recognition of the nature and range of 

earthquake effects and on an appreciation of the 

potential consequences of these effects. The purpose 

of this section is to describe and illustrate the various 

earthquake effects, beginning with the primary ones. 

Many earthquake effects have not been well 

documented in the marine environment; for these 

effects, examples are drawn from onshore earthquakes. 

Tectonic deformation 

When opposing rock masses suddenly slip along a 

buried fault to generate an earthquake, the Earth's 

surface is deformed over a broad region above the 

buried fault. For shallow earthquakes, the area of the 

deformed region is comparable to the area of slip on 

the buried fault. Such deformation accompanies all 

earthquakes; however, the amount is generally too 

small to be significant for earthquakes of Mt{,7 or with 

8 Earthquake Hazards in the Offshore Environment 

sources deeper than a few tens of kilometers below the 

Earth's surface. If the buried fault is shallow, 

displacement on the fault during an earthquake may 

extend to the surface and cause localized shear offsets 

and distortions in surficial geologic materials, as well 

as broadly distributed deformation. Accordingly, two 

types of tectonic deformation are distinguishedsurface faulting and regional deformation. 

Surface faulting 

Surface faulting is a potential hazard to 

structures built within or across active fault zones. In 

large shallow-focus earthquakes of M ~ 6, shear 

displacements at the surface may exceed 1 m over a 

zone with a width comparable to or less than the 

amount of the offset. Such displacements may be 

sufficient to rupture a buried or anchored pipeline or 

to distort the foundation of a structure to the degree 

that collapse occurs or the structure must be razed. 

The largest surface fault displacements accompanying North American earthquakes have occurred 

along the Pacific margin. In the 1906 San Francisco 

M=7. 7 earthquake, spectacular strike-slip faulting 

ruptured the surface over a 430-km-long segment of 

the San Andreas fault along the north coast of 

California (Lawson, 1908). About half of this rupture 

lay offshore. Over the entire length of the rupture, 

the ground west of the fault was displaced northward 

relative to that on the east by as much as several 

meters (fig. 5). The maximum measured horizontal 

Figure 5. Offset and distortion of fence caused by 

right-lateral strike-slip fault displacement during the 

1906 California earthquake (from Lawson, 1908, pl. 

49A). Offset of 2.6 m on main fault trace broke fence 

(in middle of photograph). Beyond break, which is 

repaired with poles, fence is distorted by distributed 

right-lateral shear. Total displacement between 

straight fence segments on opposite sides of fault is 

3.4 m. 
offset was about 6 m over a zone 15 to 18 m wide in a 

road built on marshy ground; however, it may be that 

ground failure could have contributed to the observed 

displacemeht. On competent ground nearby, offsets of 

about 5 m were measured where the slip was localized 

in a narrow zone less than 1 or 2 m wide. Larger 

horizontal displacements have been inferred for earthquakes of comparable magnitude occurring on other 

segments of the San Andreas fault. Small 

streamcourses in the Carrizo Plain area in southern 

California have been offset repeatedly by as much as 9 

or 10 m (Wallace, 1968; see subsection below entitled 

"Geologic Record" and fig. 21). 

Dip-slip fault displacement of several meters 

was documented for the great (M=9.2) Alaska earthquake of 1964 (Plafker, 1967)-:- This earthquake 

resulted from northwestward underthrusting of the 

Pacific plate beneath southern Alaska along a gently 

dipping fault. Although the primary fault on which the 

principal slip occurred did not rupture the surface, two 

reverse faults, possibly branching upward off the 

primary fault, broke the ground on Montague Island 

and caused large vertical displacements. Dip slip on 

the Patton Bay fault caused 6 to 7 m of vertical 

offset, measured across the full width of the fault 

zone. A prominent scarp accounted for 2 to 3 m of 

this offset; the rest was distributed across the downwarped margin of the upthrown block, within 300 m of 

the scarp. The rupture along the Patton Bay fault was 

traced for at least 27 km offshore from Montague 

Island by hydrographic soundings and seismic profiling 

(Malloy and Merrill, 1969). Surface faulting generated 

a more spectacular scarp along the Hanning Bay 

fault. Localization of slip within a shear zone 0.5 m 

wide created an abrupt 4-m-high scarp in bedrock (fig. 

6). 

A graphic example of the damage that surface 

faulting can cause to pipelines and structures is the 

1971 San Fernando, Calif., earthquake (Youd and 

others, 1978b). About 3 km of surface faulting associated with this M=6.6 earthquake crossed an urban 

area. Maximum components of cumulative displacement across the principal zone of faulting were 1.9 m 

of left-lateral slip, 1.4 m of vertical slip, and 0.6 m of 

horizontal shortening in a direction normal to the 

trend of the rupture zone (U.S. Geological Survey 

Staff, 1971). Practically all the lateral slip and horizontal shortening and about half of the vertical slip 

were concentrated in a narrow band less than 30 m 

wide. This band sharply defined one boundary of the 

principal rupture zone and marked the leading edge of 

the overthrust fault block. Within the overthrust block 

Figure 6. Bedrock scarp, about 4 m high, formed by secondary faulting on Montague Island during the 1964 Alaska 

earthquake (from Plafker, 1967, fig. 31). Fault dips to left about 55°. Slumping and erosion of material from 

initially overhanging scarp face result in scarp surface that slopes to right. Entire area of photograph was 

uplifted about 10 m. 

Earthquake Effects 9 
away from the leading edge, the balance of the 

vertical displacement was distributed across numerous 

small normal (extensional) faults. The number of 

extensional faults and the displacement on them 

gradually decreased away from the edge of the overthrust block. Thus, the style of horizontal deformation 

divided the principal zone of surface faulting into two 

subzones-a narrow band of lateral shearing and horizontal compression, and a wider, more diffuse band of 

horizontal extension. 

Ruptures within the principal zone of surface 

faulting in the San Fernando earthquake extended 

beneath tens of homes and several one- and two-story 

industrial and commercial buildings. Many of these 

buildings were damaged beyond repair, although none 

totally collapsed; nearly all sustained significant 

structural damage (Youd and others, 1978b). Most 

pipelines crossing major fault breaks were ruptured, 

commonly at several nearby locations. Individual 

pipelines failed in response to both compressional and 

extensional deformation. A 16-in. welded-steel gastransmission line, for example, broke in seven places 

within the principal zone of faulting (compare 

Southern California Gas Co., 1973, fig. 1, and U.S. 

Geological Survey Staff, 1971, fig. 2). Near the tip of 

the overthrust block, one segment of the pipeline was 

shortened more than 0.1 m (fig. 7), whereas farther 

from the tip but still within the principal zone, the 

pipeline failed from horizontal extension of the ground 

coupled with vertical fault slip. Numerous ruptures in 

this and other gas-transmission lines occurred within 

the overthrust block at distances as large as a few 

kilometers from the principal zone of surface 

faultihg. Although some of these ruptures resulted 

from slope failures, others probably were related to 

subsidiary surface faulting. 

Subsidence and uplift 

Broad-scale vertical deformation of the Earth's 

surface accompanies large shallow-focus earthquakes 

generated by dip-slip faulting. Regional deformation, 

either uplift or subsidence, poses a potential hazard 

where it results in large permanent changes in water 

levels along developed shorelines. For offshore pia tforms, permanent changes in water depth that might 

accompany large local earthquakes are likely to be 

small in comparison with the expected heights of 

storm-generated waves and thus are not a major design 

consideration. Along a coast, however, permanent 

changes in water levels may make harbor and pier 

facilities useless, at least during some tidal stages, and 

may cause navigational hazards to shipping where the 

shallow sea floor is uplifted. 

Vertical deformation is a first-order effect of 

dip-slip faulting. Accordingly, tectonic uplift and 

subsidence are important earthquake effects to 

consider in coastal areas characterized by either 

normal or reverse faulting. Along the Pacific margin 

of North America, for example, potential uplift or 

subsidence should be a consideration in the siting and 

design of coastal facilities along the Santa Barbara 

Channel in southern California and along the south 

coast of Alaska west of about long 138° W. 

10 Earthquake Hazards in the Offshore Environment 

Figure 7. Shortened section of 16-in. steel gastransmission line taken from principal zone of surface 

faulting associated with the 1971 San Fernando, Calif., 

earthquake (from U.S. Geological Survey staff, 1971, 

fig. 5). Amount of shortening exceeds 0.1 m. 

The 1964 Alaska earthquake provides a graphic 

example of regional vertical deformation in a coastal 

environment. Tectonic uplift and subsidence occurred 

over two adjacent elongate zones (fig. 8), each approximately 800 km long and 150 to 200 km wide (Plafker, 

1969). The maximum measured uplift on land, about 

12 m, occurred within a narrow belt, about 3 km wide, 

between the two reverse faults on Montague Island 

discussed in the preceding section (fig. 9). Comparable 

uplift was measured in the adjacent offshore area 

(Malloy and Merrill, 1969). Most of the uplifted area 

was offshore, and so the extent of uplift is poorly 

known. From 1cat~ered coastal measurements, an area 

of at least 10 km , possibly much greater, is inferred 

to have risen 3 m. The region of subsidence was 

largely onshore and thus could be better determined. 

The maximum measured tectonic subsidence was about 

2.5 m. 

The extensive vertical deformation that accompanied the 1964 Alaska earthquake caused great 

damage to harbor and waterfront facilities, both 

directly through uplift and subsidence of coastlines and 

indirectly through the generation of destructive 

tsunamis (see subsection below entitled "Tsunamis") 

(Eckel, 1967; Arno and McKinney, 1973). Uplift or 

subsidence affected every coastal community in southcentral Alaska from Kodiak to Cordova (fig. 8), 

although damage from waves and shoreline slides 

typically far exceeded that directly due to vertical 

deformation. An exception was the Cordova area, 

where 2 m of uplift was the greatest cause of earthquake damage. Dock facilities became accessible only 

at very high tides, so that one cannery had to be 

abandoned and shallow waterways became 

unnavigable. No settlements were located in the 

region of maximum uplift. Subsidence affected more 

communities and left many waterfront and port facilities subject to flooding by high tides. For example, at 

Seldovia, where virtually all the damage was caused by 
62° 

60° 

58° 

56° 

I 

I 

I 

156° 154° 150° 

0 

148° 

~ 

\)'v 

146" 

100 200 KILOMETERS 

I' 

t r 

" 

Figure 8. Gulf of Alaska area, showing tectonic uplift and subsidence during the 1964 Alaska earthquake 

(modified from Plafker, 1965, fig. 2). Heavy contours show displacement (in meters); dashed where approximately 

located. 200-m isobath (dotted line) marks edge of the Continental Shelf. 

tectonic deformation, slightly more than 1 m of 

subsidence subjected the boardwalk on which the 

business section of the town was located to flooding at 

extreme high tides (fig. 10). 

Seismic shaking 

Seismic shaking typically contributes more to the 

overall earthquake damage than does any other earthquake effect. Strong shaking contributes to economic 

and human losses not only directly, through vi bra tory 

damage to structures, but also indirectly, through 

triggering of secondary effects, such as landsliding and 

loss of shear strength in water-saturated sediment. In 

contrast to surface faulting, seismic shaking is a 

pervasive effect that subjects structures throughout a 

broad area surrounding the earthquake source to 

significant loading. 

Little information is available on the response of 

bottom-supported offshore structures to earthquakes, 

especially large events (Hove, 1983). The 1978 Santa 

Barbara Channel earthquake off southern California 

provides an example of the shaking effects from a 

nearby moderate (M=6.0) shock (Miller and Felszeghy, 

1978). A total ofT2 offshore platforms were located 

within 25 km of the buried fault rupture, as inferred 

from the spatial pattern of aftershocks (Lee and 

Earthquake Effects 11 
Figure 9. Former sea floor at Cape Cleare, Montague Island, exposed by tectonic uplift during the 1964 Alaska 

earthquake (from Plafker, 1969, frontispiece). Wide surf-cut terrace was created at southwest end of the island 

by 8 m of uplift. Photograph taken at about zero tide stage. View northward. 

Figure 10. Flooding of boardwalk and hotel in Seldovia 

at high tide, caused by about 1.2 m of tectonic 

subsidence during the 1964 Alaska earthquake. 

Photograph by U.S. Army Corp of Engineers. 

12 Earthquake Hazards in the Offshore Environment 

others, 1978); the closest platform was at a distance of 

about 10 km. Five of these platforms were instrumented with strong-motion accelerographs at the time 

of the earthquake; however, no records of the earthquake were obtained because of improper setting of 

the triggering levels of the recorders and inadequate 

maintenance. Five onshore accelerographs at 

distances comparable to that of the nearest platform 

recorded peak horizontal accelerations at ground level 

of between 0.21 and 0.42 g_ (Porcella, 1979; Porter and 

others, 1979); the larger values were recorded on the 

University of California campus at Santa Barbara in 

the direction toward which the fault ruptured. The 

shaking was amplified substantially in the upper levels 

of multistory buildings. In a three-story reinforcedconcrete shear-wall building on the university campus, 

the peak horizontal acceleration at the roof was 1.04 

g:, in comparison with 0.42 .[at the ground floor (fig. 

11). Although significant diagonal cracking occurred 

in the shear walls of the building, serious structural 

damage did not occur because the duration of strong 

shaking was only 2 to 3 s. Mechanical equipment on or 

near the roofs of multistory buildings on the campus, 

however, was considerably damaged from building 

motion, and objects were thrown from shelves. No 
1.0 

ROOF LEVEL 

North 

0.5 

0.5 

South 

0L----L----~----~----L---~----~----~----~----L---~----~----~----~--~ 

C)) 0.5 

~ North GROUND LEVEL 

z 

0 

~ <{ 

a: 

LLJ 

...J 

~ 0 ~--'-""""""'-,......"""'"""""Mr...,_J\ 

u 

<{ 

South 

0 

·

2

J- ~Up- • • '-.A. A "' i\ ~I... A Ill • ~ A "' • .!\ • • .A • IV. GRR:OUUNNDD LEVEEVELL J 

25L---~----~----L---~ 

~ :~~~-uvv,·~vr, vv~~~VN1' ____ ...JL v\N· 

____ 

~V 

L_ __ 

v· 

_J 

,~~ 

____ ...JL ____ 

v-~' ~ __ _...JL_ ~ 

__ _L ____ 

~ 

~----L---~ 

0.5 

East GROUND LEVEL 

West 

2 3 4 5 6 7 8 9 10 11 12 13 14 

TIME, IN SECONDS 

Figure 11. Accelerograms recorded on ground floor and roof of three-story North Hall on the 

University of California, Santa Barbara, campus during 1978 Santa Barbara M=6.0 

earthquake. Amplitude of north-south motion on roof (top trace) was more than twice that 

at ground level (second trace). Vertical and east-west components of motion were not 

recorded on roof. Recorder was triggered by vertical E_-wave motion sensed at ground 

level. Time is measured from trigger instant. 

Earthquake Effects 13 
damage was reported to any offshore platforms, wells, 

or oil and gas lines. On the platform closest to where 

the highest ground motions were recorded on shore, 

noticeable rocking was reported to have lasted a few 

seconds. Vibrations were sufficient to automatically 

shut down compressors on 3 of the 12 platforms. 

The ground shaking from a nearby M > 7 earthquake would be much more significant To- offshore 

platforms than that from the Santa Barbara shock. A 

larger earthquake would generate more intense 

motion, especially at periods longer than about 1.5 or 2 

s, where the fundamental deformational modes for 

pile-supported steel and concrete gravity platforms 

typically lie (Watt and others, 1978; Bea and others, 

1979). Because of the greater rupture dimensions, not 

only would ground-motion spectra for a larger earthquake be richer in long-period energy (Johnson and 

Traubenik, 1978), but also the duration of motion 

would be much longer. 

Strong ground shaking may pose a hazard not 

only to the survival of an offshore platform but also to 

the safety and continuity of operations on the platform. As illustrated above, the motion imposed at the 

foundation of a structure is amplified at higher levels 

in the structure. At the top of a structure, it may be 

sufficient to cause substantial damage to equipment 

and appurtenances, some of which may be critical for 

safe operation or shutdown, even though little damage 

may be done to the structure itself (Kost and Sharpe, 

1977). Nonstructural earthquake damage has been 

documented in considerable detail for the 1971 San 

Fernando, Calif., and the 1964 Alaska shocks-a nearby 

moderate event and a distant large shock, respectively 

(Ayres and Sun, 1973; Ayres and others, 1973). 

The direct effects of seismic shaking also constitute a potential hazard to other types of facilities 

associated with petroleum production. Shaking has 

caused extensive damage to modern storage tanks and 

connecting piping (Rinne, 1967; Hanson, 1973; Kennedy 

and others, 1977; Miles, 1977). By contrast, modern 

petroleum-transmission pipelines have withstood the 

direct effects of ground shaking well. Rupture of a 

buried welded-steel transmission pipeline from shaking 

alone has not been documented, although such pipelines have failed from the effects of faulting and 

permanent ground deformation (Kennedy and others, 

1977). Modern communication and control systems 

have been found to be vulnerable to shaking, as was 

documented in the damage reports for the 1971 San 

Fernando earthquake (Benfer and Coffman, 1973). 

Seismic shaking is a concern not only for marine 

structures supported on the sea floor but also for 

floating structures (Hove and others, 1982). The 

32,500-ton Norwegian motor tanker Ida Knudsen was 

located above the source region of B1l M=7 .8 earthquake that occurred in 1969 about 450km west of 

Gibraltar. Compressional seismic waves from the 

earthquake violently shook the ship and caused damage 

so extensive that the ship was classified as a total 

loss. Hove and others (1982, p. 5) summarized the 

damage as follows: 

* * * all communication and navigation equipment 

was destroyed. Instruments mounted on the walls 

were torn off. Doors were torn off their hinges. 

Handrailing on stairways were shaken off. Loose 

14 Earthquake Hazards in the Offshore Environment 

equipment and furniture were thrown up into the 

air. Outriggers on mast were shaken off. Piping 

was broken. Some equipment was torn loose in 

way of anchorbolt failure or stretching. Leaden 

linings in way of machinery mountings were 

squeezed out. Welded connections and stiffeners 

had been broken. Bulkheads, hull frames and 

girders were buckled. Bulkheads were severely 

torn in one tank. All the wing tanks leaked, 

however the outer skin was tight except for one 

tank. In general the bottom plating and also the 

lower part of the side plating were torn away 

from the stringers/girders yielding gaps up to 50 

mm wide.*** 

After inspection in drydock at Lisnave the 

ship was condemned as total loss. The hull looked 

as if it had been subjected to a heavy mine explosion, the inspection report states, which indicates 

extreme dynamic pressure loadings. * * * 

Whether the damage to the Ida Knudsen is typical of what should be expected above the source region 

of other similar-size earthquakes is not known. Hove 

and others (1982) presented additional examples of 

seismic disturbances to ships; none involved comparable levels of damage, but none was near the focus of 

such a large earthquake. They argued that additional 

research is needed to determine the extent to which 

offshore structures, particularly floating structures, 

are vulnerable to compressional seismic waves transmitted through the water. 

Sea-fioor failures 

Ground failures generated by seismic shaking are 

a major cause of earthquake damage and casualties on 

land. Although earthquake effects in the marine environment are relatively poorly documented, there is 

ample historical and geologic evidence of offshore 

earthquakes triggering large-scale sea-floor failures. 

Furthermore, stability analyses indicate that seismic 

shaking is a likely triggering mechanism for failures on 

many submarine slopes in earthquake-prone regions 

(for example, Almagor and Wiseman, 1977; Hampton 

and others, 1978; Lee and others, 1981). Thus, the 

potential for seismically induced failure in sea-floor 

materials is an important consideration for the safe 

siting and design of offshore facilities in earthquakeprone regions. 

As used in this report, the terms "ground failure" 

and "sea-floor failure" refer to the temporary loss of 

bearing capacity in surficial geologic materials as well 

as to their loss of stability and resulting deformation. 

Accordingly, downslope movement of a slide and loss 

of foundation support in the liquefaction of granular 

water-saturated sediment are both classified as failures. 

Relatively few seismically induced sea-floor 

failures have been documented in relation to specific 

earthquakes. Nearly all these failures involved sliding 

of relatively steep submarine slopes associated with 

delta fronts or the continental slope, and have been 

recognized from the damage to coastal facilities or 

submarine cables resulting directly from either 

submarine slides or turbidity currents (see next subsection), or from waves generated by submarine ~lides 
(see subsection below entitled "Tsunamis"). The dearth 

of documented failures in continental-shelf areas, 

which are characterized by very gentle slopes 

(typically, approx 0.1 °), undoubtedly reflects inexperience with offshore earthquake phenomena rather than 

the absence of such failures. 

The 1964 Alaska earthquake generated 

NORTH 

B 

catastrophic submarine slides in the fiords of Prince 

William Sound and along the south coast of the Kenai 

Peninsula. In Seward, delta-front failure removed a 

strip of land about 1.2 km long and as much as 150 m 

wide from the city waterfront (Lemke, 1967). Harbor, 

dock, and railroad facilities were lost in the 

progressive sliding (fig. 12), which began some 30 to 45 

Figure 12.. Seward, Alaska, waterfront, before (A) and after (_m the 1964 earthquake (from Lemke, 1967, fig. 2). 

Note loss of dock, harbor, and railroad facilities from waterfront slides. White line in figure 12~ denotes area 

swept by water waves. 

Earthquake Effects 15 
s after violent shaking commenced. When the shaking 

ceased, the ground behind the landslide scarp was 

found to be fractured in an incipient-failure condition 

to distances of more than 250 m from the new shoreline. The submarine sliding generated massive waves 

that swept the remaining waterfront area, lifted 

railroad cars and vehicles, and carried burning oil from 

slide-toppled tanks and ruptured pipelines and valves. 

Submarine sliding and slide-generated waves also 

wreaked havoc in Valdez (Coulter and Migliaccio, 

1966) and Whittier (Kachadoorian, 1965). These calamitous failures, which were caused by liquefaction of 

deltaic sediment, occurred on moderately steeg slopes, 

ranging from a few degrees to greater than 20 . There 

is limited evidence that submarine failures also 

occurred on gentler slopes (Reimnitz, 1972). Multiple 

slumps in foreset beds of the Copper River delta are 

attributed to the 1964 earthquake ~ig. 13); the slope 

of the sea floor there is about 0.5 • These multiple 

failure surfaces suggest progressive failure of the 

slope during extended seismic shaking. 

The best documented case of earthquake-induced 

slope failure on the Continental Shelf known to us 

comes from the M=7 .4 earthquake of November 8, 

1980, off Eureka-;- northern California (Field and 

others, 1982). After this earthquake, local commercial 

fisherman reported changes of the sea floor in an area 

previously surveyed by deep-penetration and highresolution seismic-reflection techniques. Additional 

seismic profiles were obtained 5 weeks after the 

shock. Comparison of these postearthquake records 

with those obtained 1 to 3 years earlier (fig. 14A) 

reveals an extensive (approx 1 by 20 km) zone of 

shallow (5 m thick) failure in unconsolidated sediment 

at a water depth of 60 m on a nearly flat (less than 

0.25° slope) bottom. Geomorphic features diagnostic 

of sediment flow and lateral spreading are visible in 

side-scan sonar records obtained after the earthquake 

(fig. 14B). The failure zone is at least 30 km from the 

closest- point of seismic-energy release in the 

earthquake, as inferred from the distribution of aftershock epicenters. 

The potential hazard that submarine slides pose 

to offshore structures is exemplified by the loss of 

South Pass 70 Platform B in the Gulf of Mexico 

(Sterling and Strohbeck, 1973). This platform failed 

from large-scale downslope movement of sediment to 

a depth of at least 20 m. The platform was found 

resting on its side on the bottom, displaced about 25 m 

in the downslope direction. In this case, the movement 

was triggered not by seismic shaking but by 20-m-high 

waves generated by Hurricane Camille. 

Pipelines, because of their extended lengths, are 

particularly susceptible to failure from submarine 

slides, as is seen by analogy, using experience with 

communication cables lying on the sea floor. In the 

1929 Grand Banks M=6.5 earthquake off the south 

coast of Newfoundland, submarine sliding ruptureq 

seven cables within an area of about 35,000 km 

surrounding the epicenter (Heezen and Drake, 1964), 

and associated turbidity currents broke other cables to 

distances of 700 km from the epicenter (see next 

section and fig. 15). 

Loss of bearing strength without downslope 

movement is also a potential hazard to structures 

situated on or buried in liquefiable sediment in offshore or coastal environments. Earthquake-induced 

liquefaction can cause structures founded on sediment 

to settle and tilt, and buried structures to rise by 

buoyancy through liquefied sediment. Some of the 

most striking onshore examples of liquefaction effects 

were associated with the 1964 Niigata, Japan, M=7 .6 

earthquake (Seed and ldriss, 1967). -

Figure 13. Subbottom seismic-reflection record (vertical section), showing multiple slumps in foreset beds of the 

Copper River delta, Alaska (from Reimnitz, 197 2, fig. 4). Slumps are attributed to progressive slope failure 

during the 1964 earthquake. Vertical exaggeration, x27. 

16 Earthquake Hazards in the Offshore Environment 
Figure 14. Earthquake-induced sediment failure on the shallow Continental Shelf off mouth of the Klamath 

River, northern California. High-resolution seismic-reflection (Uniboom) records (vertical sections) along nearly 

coincident profiles were made before ~ Oct. 1979) and after @, Dec. 1980) the November 1980 earthquake (from 

Field and others, 1982, fig. 3). Vertical scale bars, 10 m; horizontal scale bars, 250m; triangles mark 60-m water 

depth. Terrace and toe-scarp morphology in figure 14~ results from sediment flow and lateral spreading after 

seismically induced liquefaction. Note smooth, undisturbed sea floor before the earthquake • .£, Side-scan-sonar 

record (plan view) along toe of failure, showing features diagnostic of sediment flow and lateral spreading (from 

Field and others, 1982, fig. 4). Lateral range is 100m to either side of ship's trace (centerline). Sinuous 1-m-high 

main scarp marks termination of flow. Pressure ridges are inferred to be caused by local compression of sediment 

seaward of advancing flow or spread. Circular features are inferred to be sand-boil vents and collapse craters 

resulting from ejection of trapped gas or liquefied sediment. 

Earthquake Effects 17 
Turbidity currents 

Sediment, after slumping, can exhibit a broad 

range of mobility states, from rigid block motion to 

turbulent flow. The type of motion after slumping 

depends on the slope profile, the shape and position of 

the slip surface, pore-water pressure within the sediment, and the strength and density of the sediment. 

However, the conditions that must be satisfied for the 

onset of turbulent flow, or turbidity currents, are not 

well understood (Plapp and Mitchell, 1960; 

Morgenstern, 1967; Hampton, 1972). 

A turbidity current is a density current in which 

water containing a large amount of suspended sediment flows downslope under the influence of gravity. 

After reaching a relatively level section of the 

bottom, the current may continue to flow for a long 

distance. The sediment can be put in suspension in 

several ways, for example, by sedimentation processes 

at a river mouth, by wave action, or by landslides into 

or within a body of water. The largest known turbidity 

currents have been caused by earthquake shaking on a 

continental slope or shelf that has triggered landslides 

and slumps. Evidence of the extent of many 

earthquake-induced turbidity currents has been 

provided by breakage of cables on the sea floor; for 

example, the Grand Banks south of Newfoundland 

(Heezen and Ewing, 1952; Heezen and others, 1954; 

Heezen and Drake, 1964), the Ionian Sea (Ryan and 

Heezen, 1965), the western New Britain Trench (Krause 

and others, 1970), and the Algerian coast (Heezen and 

Ewing, 1955). 

The Grand Banks turbidity current in 1929, 

caused by the M=6.5 earthquake on the continental 

slope south of Newfoundland, has received the most 

study. The earthquake caused a large gravity slump 

(fig. 15), the extent of which was inferred from 

interpretation of a seismic-reflection profile and the 

locations of cables that failed at the instant of the 

earthquake (Heezen and Drake, 1964). The slump 

generated a turbidity current that swept down the 

slope and onto the abyssal plain. On the basis of the 

sequence of cable breaks, the velocity of this current 

was estimated to be approximately 100 km/h near the 

base of the Continental Shelf. Piston cores, showing a 

graded layer of sediment, and the locations of the last 

cable breaks indicated the extent of the current on the 

abyssal plain (fig. 15). Kuenen (1952) estimated that 

the thickness of the initial slump was about 50 m and 

of the resulting turbidity current about 270 m. 

Turbidity currents on the continental slope are 

known to scour submarine canyons and to deposit large 

volumes of sediment on the sea floor, well out on an 

abyssal plain. Because of the gentle slopes that 

characterize the Continental Shelf, however, sea-floor 

failures on the shelf are unlikely to be transformed 

into turbidity currents unless they involve steep slopes 

in an area of rapid sedimentation. Such circumstances 

exist in Valdez Fiord in southern Alaska, where earthquakes repeatedly have triggered submarine slides that 

apparently were transformed into turbidity currents 

(Coulter and Migliaccio, 1966). On five separate 

occasions, submarine cables have been broken and 

buried. 

18 Earthquake Hazards in the Offshore Environment 

100 200 KILOMETERS 

EXPLANATION 

* EPICENTER 

X CABLE BREAK 

----FAULT 

Figure 15. Area of 1929 Grand Banks, Newfoundland, 

submarine slump and turbidity current, showing 

epicenter of M=6.5 earthquake and locations of cable 

breaks (from Doxsee, 1948) and faults (from King, 

1980). Bathymetric contours in meters. Approximate 

extent of slump and turbidity current from Heezen and 

others (1954) and Heezen and Drake (1964). 

The potential hazards of turbidity currents to 

offshore facilities arise from the hydrodynamic forces 

imposed on obstacles in the path of the current and 

from sediment scour around the foundation of a structure or a buried pipeline. Although the hazard of a 

turbidity current may not match that of the initiating 

submarine slide in terms of the potential for concentrated damage within a localized area, the turbidity 

current will affect a much larger area, extending far 

downslope from its point of origin. 
Tsunamis 

A tsunami is a gravitational seawave caused by 

the sud?en displacement of a large volume of water. 

Tsunamis can be generated by seismic and volcanic 

activity or by submarine and shoreline landslides. 

Here, we confine our discussion to earthquake-related 

tsunamis caused by sudden vertical displacement of 

the sea floor and by the sliding of subaerial and submarine sediment induced by seismic shaking. 

Tsunamis on the open ocean fall under the 

general classification of long waves, with wavelengths 

of several hundred kilometers. Their amplitudes over 

the deeper part of the ocean do not exceed 1 m and t~us are difficult to detect from ships or from the 

air. The wave velocity is proportional to the square 

root of the water depth and in the deep ocean can be 

~ev~ral hundred kilometers per hour. As a tsunami 

Impmges on the Continental Shelf, its velocity 

decreases, and the wave height increases. A series of 

waves approaches the coast with periods ranging from 

about 5 minutes to more than 2 hours. The wave with 

the greatest height is generally not the first but commonly occurs among the first 10. A tsunami 

approaching a coastline is subjected to several modifications: energy may be reflected from the continental 

slope; the Continental Shelf may act as a waveguide 

tending to trap energy at wavelengths near that of the 

shelf width; and inlets, harbors, and embayments may 

cause significant amplification of the waves. 

The principal potential hazard from tsunamis is 

to coastal and shallow-water facilities. Tsunami wave 

height does not become comparable to that of storm 

waves until the wave reaches quite shallow water. 

Under the assumption that no dissipation or reflection 

of energy occurs as the wave approaches shore an 

appreciable proportion of the increase in wave h~ight ~ith shoaling w.ater depth can be accounted for by 

lmear theory (Wiegel, 1970). A 1-m-high wave on the 

open ocean would increase in height to about 3 m in a 

water depth of 50 m and to about 5 m in a water depth 

of 10 m. Therefore, tsunami waves would not subject 

platforms and other facilities in water depths of tens 

of meters to loads greater than those generated by 

storm waves. 

. Sudden vertical displacement of the nearly 

horizontal sea floor can displace large volumes of 

water, whereas horizontal displacement cannot. In the 

1964 Alaska earthquake, extensive vertical 

deformation of th~ sea floor (fig. 8), caused by lowangle thrust faultmg beneath the Continental Shelf 

generated several destructive tsunamis as discussed 

below in further detail. In contrast, no ~ajor tsunami 

was reported after the 1949 Queen Charlotte M=S.l 

earthquake, which had a strike-slip fault mechanism. 

A~l types of earthquakes, however, including strike 

s~Ip, ~an ~enerate local destructive tsunamis through 

VIbrahon-mduced submarine slides. 

Local tsunamis 

For the purposes of discussion, local tsunamis are 

seawaves impacting near their place of origin. Local 

tsunamis, which devastated coastal communities 

during or immediately after the 1964 Alaska earthquake, caused 103 deaths and more than $85 million in 

damage (Wilson and Torum, 1972a). These waves were 

generated by two types of mechanisms: the vertical 

displacement of a large area of the Continental Shelf 

(fig. 8), which generated trains of long-period waves 

that first struck the coast about 20 minutes after the 

earthquake (von Huene and Cox, 1972)· and the 

displacements of large volumes of water' by sudden 

failures of unconsolidated sediment at or below sea 

level, which generated short-period waves that were 

largely restricted to a single fiord or strait. Slidegenerated waves caused more destruction and loss of 

life than the tsunami generated by deformation of the 

Continental Shelf (Weller, 1972). 

At Seward, the wave generated by the slide that 

carried away a 1.2-km length of the waterfront (fig. 

12) reached a maximum height of about 10 m. It 

caused considerable damage to railroad facilities and 

spread burning oil along the waterfront from tanks 

that had ruptured (fig. 16). About half an hour later 

the first wave from the major tsunami source on th~ Continental Shelf struck Seward and caused additional 

extensive damage and spread burning oil farther over 

the waterfront (Spaeth and Berkman, 1972; Wilson and 

Torum, 1972a). 

Local waves were generated by slides in two 

separate areas of Port Valdez (fig. 4, inset). Waves 

caused by submarine slides off the terminal moraine at 

the mouth of Shoup Bay washed the Middle Rock 

navigation light in Valdez Narrows off its 10-m-high 

reinforced-concrete pedestal. The delta-front slide 

that carried the Valdez dock into the sea generated a 

violent surging wave that demolished most of the 

remaining waterfront facilities and completed the 

wrecking of the fishing fleet (Coulter and Migliaccio 

1966). ' 

Kodiak Island was the only place within the 

affected area of the Gulf of Alaska for which fairly 

detailed information on the wave sequence and wave 

heights is available, owing to a log kept by a naval 

officer at Womens Bay. A reconstructed marigram 

shows the water-level fluctuations in Womens Bay (fig. 

1 7). This record is resolvable into three major components: the astronomic tide; a train of modulated 

waves with periods of about 2-1/2 hours, related to the 

half-length of the wave generated on the Continental 

Shelf; and an oscillation of about SO-minute period . ' representmg the second mode of free oscillation of 

water on the Continental Shelf (Wilson and Torum, 

1972a). Waves continued to inundate the shoreline of 

Kodiak Island with progressively decreasing amplitude 

until about 12 hours after the earthquake. These 

~av.es caused extensive damage to shoreline buildings, 

fishmg vessels, docks, navigation equipment, bridges 

and highways. High-velocity currents associated with 

the repeated ebb and flood of waves resulted in 

extensive damage through erosion to artificial fills and 

unconsolidated deposits. Bottom changes offshore 

from Kodiak Island occurred in water depths of as 

much as 25 m (Plafker and Kachadoorian, 1966). 

Waves were also reported on the coast of Kodiak Island 

before the principal tsunami wave train illustrated in 

figure 17. 

On the Atlantic margin, the 1929 Grand Banks 

M=6.5 earthquake generated a tsunami with amplitudes 

Earthquake Effects 19 
Figure 16. Damage to railroad yard and petroleum-tank farm at Seward, Alaska, caused by waterfront slides, 

waves, and fire (from Lemke, 1967, fig. 6). 

of at least 12 m on the Burin Peninsula on the south 

coast of Newfoundland (fig. 15). The tsunami, which 

reached the shore at a time of abnormally high tide 

and during a heavy gale at sea, caused the loss of 27 

lives and extensive damage to homes and fishing 

equipment (Doxsee, 1948). The traveltime of the 

tsunami from the epicenter to the Burin Peninsula was 

approximately 2-1/2 hours (Murty, 1977). There was 

no warning because no warning system was, or is now, 

in place along the East Coast. This is the only documented case of tsunami damage on the east coast of 

North America, although the Caribbean Islands have 

undergone several tsunamis from local and eastern 

Atlantic sources (Murty, 1977). It is problematic 

whether the 1929 tsunami was generated by tectonic 

deformation of the sea floor related to faulting or by 

submarine slumping, although slumping seems more 

likely. 

Distant tsunamis 

The principal source of the major tsunami 

genera ted by the 1964 Alaska earthquake, which swept 

20 Earthquake Hazards in the Offshore Environment 

from the Gulf of Alaska across the length of the 

Pacific and lapped against Antarctica, was sudden 

uplift of the Continental Shelf (fig. 8). Extensive 

studies have been made of the source mechanism, 

oceanic properties, propagation, coastal modification, 

runup heights, and associated damage of this tsunami 

(Berg and others, 1972; Spaeth and Berkman, 1972; Van 

Dorn, 1972; Van Dorn and Cox, 1972; Wilson and 

Torum, 1972b). 

Along the coast of western Canada, the 

maximum heights of tsunami waves recorded at most 

tide stations ranged from 2 to 4 m. An exception was 

the Alberni Inlet on western Vancouver Island, where 

amplifying effects generated water levels as much as 9 

m above average and caused extensive damage near 

the head of the inlet (White, 1966). Some of the 

highest waves occurred along the Washington, Oregon, 

and northern California coastlines, owing to the coincidence of tsunami waves with high spring tide. 

Unusually high waves at Crescent City, Calif., have 

been attributed to dynamic amplification by resonance 

on the Continental Shelf, as determined by the shape 

of the coastline and the bathymetry of the shelf 

(Wilson and Torum, 1972a). Along the coast south of 
6~------------------------------------------------------------------------------------, 

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Breaking wave on southwest shore 

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17 18 19 20 21 22 23 0 1 2 3 4 5 6 7 8 9 12 

ALASKA STANDARD TIME, MARCH 27-28, 1964 

Figure 17. Reconstructed marigram for water-level fluctuations in Womens Bay, Kodiak Island, Alaska, caused by 

principal tsunami from the 1964 earthquake (from Wilson and Torum, 1972b, fig. 4). Widely fluctuating solid 

curve, observed water levels; smooth solid curve, astronomic tide; dashed curve, long-period tsunami waves. 

MHHW, mean higher high water; MLLW, mean lower low water; MSL, mean sea level. 

Crescent City, runup was generally much less than 

that farther north because the tsunami waves arrived 1 

to 2 hours after high tide. Elsewhere around the 

Pacific, runup above ordinary tidal level was about 4 m 

in Hawaii, 2 m on the coast of Chile, and 0.8 m on the 

coast of Japan. 

The 1929 Grand Banks tsunami did not cause 

extensive damage in areas other than the south coast 

of Newfoundland. The islands of St. Pierre and 

Miquelon, off the tip of the Burin Peninsula (fig. 15), 

escaped damage. The east coast of Nova Scotia was 

flooded in several places; fishing wharves and one ship 

were damaged at Canso. In Bermuda, a dredging 

vessel broke its mooring chains. The wave was large 

enough to be noticed in the Azores, but tide gages in 

the Gulf of St. Lawrence showed no unusual waves 

(Doxsee, 1948). 

The tsunamis from the 1964 Alaska and 1929 

Grand Banks earthquakes are two of the many hundred 

that have been documented in the Pacific, the eastern 

Atlantic, the Caribbean, and the Mediterranean 

(:viurty, 1977). 

ASSESSMENT OF EARTHQUAKE POTENTIAL 

There are three strategies for minimizing the 

potential hazards of earthquakes: avoidance, accommodation, and emergency planning. The success of all 

three strategies depends on the abilities of earth 

scientists, first, to assess the potential for earthquakes 

to occur and, second, to predict their effects. This 

section discusses the assessment of earthquake 

potential; prediction, evaluation, and mitigation of 

earthquake effects are discussed in the next section. 

Key considerations in assessing earthquake 

potential are the location and size of future earthquakes and the frequency and imminence of their 

occurrence. To assess this potential, the earth 

scientist must develop a conceptual model 

incorporating geologic and geophysical information 

available at various scales to relate the occurrence of 

earthquakes to ongoing tectonic processes. This model 

must pertain to a region considerably larger than the 

site or area of interest because it must encompass all 

earthquakes that could have a significant effect. 

Because of inadequacies and uncertainties in the 

underlying seismotectonic models, only rarely can 

earthquake potential be assessed with a high degree of 

certainty. For example, the historical record of 

seismicity within a given region typically is short in 

comparison with the repeat time for the maximum size 

of earthquake likely to occur in the region. Thus, the 

historical earthquake record alone does not usually 

afford reliable estimates of the size and recurrence 

rate of a maximum earthquake. Although the geologic 

record may preserve evidence directly related to the 

size and date of large prehistoric earthquakes, 

generally the earth scientist has to rely on less direct 

evidence and exercise considerable judgment in 

assessing the magnitude and frequency of the largest 

possible earthquake in a particular region. 

Assessments of earthquake potential are likely to 

be more reliable in seismically active regions than in 

those of low seismicity. In active areas, where the 

repeat time for large earthquakes is shorter, the 

Assessment of Earthquake Potential 21 
likelihood is greater that the historical record includes 

the maximum earthquake that can be generated by 

specific geologic structures, as well as the cyclic 

variations in seismicity related to the accumulation 

and release of strain energy in major earthquakes; 

moreover, the opportunities to investigate the relations between earthquakes and geologic structures and 

processes are more numerous. Also, in areas of high 

seismicity the tectonic processes that cause 

earthquakes are generally the dominant processes that 

sculpt the landscape, and so the earthquake-generating 

processes are more readily deciphered from the recent 

geologic history. In contrast, the time intervals 

between earthquakes in relatively inactive areas are 

longer, and so it is less probable that the maximum 

possible shock has occurred within historical time. 

Furthermore, erosional processes are more likely to 

obliterate or obscure evidence of current seismotectonic processes that might otherwise be clearly 

preserved in the surficial geology. 

Seismotectonic models are developed from 

information derived from two sources: the historical 

seismic record and the recent geologic record. The 

historical record contains relatively detailed and 

reliable information about earthquake occurrence, but 

typically it is short in comparison with the repeat time 

of large earthquakes. The geologic record supplements 

the historical record by greatly extending the time 

interval over which information on earthquake occurrence is available; however, the information gleaned 

is generally far less certain and detailed. 

Seismie reeord 

Historical seismicity 

Historical seismicity provides an initial estimate 

of earthquake potential, that is, the location, size, and 

frequency of future earthquakes in a given region. 

Seismic history is generally divided into two eras by 

the point in time when instrumental recording of 

earthquakes began. Not until about the turn of the 

20th century were even larger earthquakes routinely 

recorded by seismographs. The length of the 

preinstrumental period is determined by the history of 

human settlement and record keeping, and varies 

considerably among the active seismic regions of the 

globe. For example, useful records date back about 

150 years in western North America, 300 years in 

eastern North America, but about 2,000 to 3,000 years 

in China and the eastern Mediterranean region. 

Coastal settlements provide a record of offshore 

earthquakes large enough to have been felt on land, 

but the accuracy with which these events can be 

located, on the basis of the distribution of damage and 

other effects, is considerably less than for similar-size 

earthquakes onshore. For example, reports of felt 

earthquakes from fishing villages along the Labrador 

coast date from as early as 1809 (Smith, 1962), and 

epicenters for these earthquakes were assigned to the 

locations at which they were felt. No evidence is 

available, however, from recent instrumental data that 

significant earthquakes are occurring onshore in this 

area; the older events most likely occurred offshore in 

an active zone of seismicity defined by instrumentally 

22 Earthquake Hazards in the Offshore Environment 

determined epicenters of the past 50 years (Basham 

and Adams, 1982, fig. 2). 

The length and value of the instrumental record 

of seismicity depends on the history of seismograph 

operations, in particular, the number of stations, their 

sensitivity, and the methods used to compile data. For 

example, although it is unlikely that the occurrence of 

onshore earthquakes of M 2:.7 in southwestern British 

Columbia would have been missed after the beginning 

of regular publication of newspapers, or after about 

1860, a similar record for offshore shocks is 

unavailable before the routine operation of the 

Victoria seismograph, which began in 1898. Reporting 

of events as small as M=6 offshore of British Columbia 

is incomplete before T919, when systematic attempts 

were made by the British in the International Seismological Summary (ISS) to gather seismograph readings 

and publish a global summary of earthquake 

locations. There may, however, be significant errors 

in these results, in part because of the practice of the 

ISS to locate specifically only the first earthquake 

from a given region and then to assign subsequent 

earthquakes from the same region to the first 

epicenter (Dewey, 1979). At present, the onshore 

network of seismographs in western Canada is capable 

of routinely locating only earthquakes of M~4 in the 

offshore zones. However, within the southern Georgia 

Strait and Puget Sound region, monitored by the 

British Columbia and Washington State networks, the 

present threshold for complete detection and location 

is M"'2-1/2. 

- Studies of historical seismicity, in addition to 

revealing the locations and rates of significant earthquakes, provide information on the tectonic component 

in seismotectonic models. For example, comparison of 

the directions of fault slip inferred from focal 

mechanisms of individual earthquakes in the Gulf of 

Alaska region with those derived from analysis of 

global plate motions suggests that the Pacific plate is 

being subducted beneath the continental margin 

between Icy Bay and Cross Sound, but only at about 

one-sixth the rate of subduction occurring to the west 

off the Kenai Peninsula and Kodiak Island (Perez and 

Jacob, 1980). As a further example, from comparison 

of slip rates inferred from historical seismicity with 

those estimated from plate-motion models, in conjunction with sediment deformation along the continental 

margin and with geodetic measurements, Weichert and 

Hyndman (1983) have concluded that most of the 

convergent movement between the Pacific and North 

American plates in the Puget Sound region is accommodated by aseismic deformation associated with 

underthrusting and continental compression. 

Location of offshore earthquakes 

Three reasons for uncertainties or biases in the 

locations of earthquakes are: (1) poor data, a 

particularly serious problem for older earthquakes 

detected on insensitive seismographs with poor timing 

control; (2) poor distribution of stations about the 

epicenter; and (3) incorrect velocity models. Attempts 

(Gawthrop, 1978; Hanks, 1979b) to accurately locate 

the 1927 M=7.3 earthquake off Lompoc, Calif., 

illustrate the difficulties with both early seismograph 
recordings and azimuthally biased data from onshore 

networks. In attempting to locate this significant 

earthquake, these and other workers have used combinations of: the arrival times of seismic waves from 

the main shock; the time intervals between the P and S 

phases for the immediate aftershocks; the distribution 

of earthquakes, assumed to be aftershocks, that 

occurred for a few decades after 1934, when 

epicentral locations began to be routinely reported on 

the basis of data from an expanded California seismograph network; the distribution of strong shaking in the 

adjacent coastal region; geodetic-survey data; and the 

presence of major offshore faults that have been 

mapped in recent years by seismic-reflection 

profiling. Considerable uncertainty still exists, 

however, both in the exact location of the epicenter 

(uncertainty of tens of kilometers) and in identification of the causative fault. 

In addition to uncertainties, serious biases also 

may exist in the locations of offshore earthquakes 

determined from data recorded by onshore 

seismograph networks. Where seismograph stations 

are unevenly distributed in distance and azimuth 

relative to the earthquakes being located, the network 

can resolve hypocentral errors in some directions 

better than in others, and so, for example, errors in 

latitude are correlated with errors in longitude. 

Dewey (1979) presented an example (fig. 18) in which 

this correlation causes a spurious alinement of epicenters. This apparent alinement was initially 

interpreted (Vrana, 1971) as evidence for a nascent 

fracture zone striking northeast toward the California 

coast north of Point Arguello, in the vicinity of a 

proposed large nuclear powerplant. More thorough 

analysis of a larger data set, however, linked the 

earthquakes to a major northwest-trending fault 

mo 

35° 

mo J210 

• 

. "'· • • 

. / • • • Point • • • • Arguello ..0 

"10 • • /~ • /0 • 

34° 

• 00 

I:// 

""' A 

system (Gawthrop, 1975). The epicenters in figure 18A 

were determined individually by the International 

Seismological Centre with P-wave arrivals from 10 or 

more stations located to the north and east. The trend 

of these epicenters suggests an active northeasttrending offshore fault zone. Dewey employed a jointhypocenter procedure to relocate these earthquakes 

relative to one master event (fig. 18B). The better 

determined, relocated epicenters indicate a northwestward trend that is statistically significant at the 

90-percent-confidence level. 

The joint-hypocenter method minimizes the 

effect on computed hypocenters of unknown errors in 

the assumed traveltimes from a source region to 

individual stations. The difference between the 

observed and computed traveltime, or the traveltime 

residual, at a given station is postulated to be the 

same for all earthquakes occurring within a limited 

source region. To correct for the effect of traveltime 

errors, or anomalies, the observed traveltimes are 

adjusted by subtracting station corrections. In the 

joint-hypocenter method, both station corrections and 

hypocenters relative to a master event, whose location 

is assumed to be accurate, are determined simultaneously to minimize the observed traveltime residuals. 

This procedure yields precise relative locations for a 

group of earthquakes; however, if the master event is 

mislocated, the other events will be systematically 

mislocated in the same way. 

Routine location methods based on teleseismic 

arrivals (that is, data recorded at distances greater 

than about 2,000 km from the epicenter) assume a 

spherically or ellipsoidally symmetrical velocity structure for the Earth's interior. Where major lateral 

variations in structure occur, this assumption can bias 

the computed hypocenters by tens of kilometers, even 

1230 122° 121° 

35° 

..0 

"10 /~ 

/0 

340 

B 

50 KILOMETERS 

Figure 18. Earthquake epicenters off the coast of central California, determined by various location procedures 

(from Dewey, 1979, figs. 4, 5). Epicenters in figure 18A determined individually, and those in figure 18~ 

simultaneously by the method of joint hypocenter determination. Principal axes of a typical 90-percentconfidence ellipse for a computed epicenter are plotted to southeast of the group of epicenters in both figures. 

Dots in figure 18~ denote epicenters with semiaxes less than 15 km long, and circles epicenters with larger 

ellipses. Apparent northeast-southwest alinement of epicenters in figure 18A is spurious and is caused by a 

correlation of errors in latitude with errors in longitude. 

Assessment of Earthquake Potential 23 
when a relatively uniform azimuthal distribution of 

arrival-time data is available. Two of the strongest 

departures from assumed symmetry are the contrast in 

seismic-wave velocities across the continental margin 

and the high seismic velocities associated with the 

downgoing plate at subduction zones. For such departures, the accuracy of hypocenters can be improved if 

traveltimes are computed by tracing seismic rays 

through an appropriate Earth model, such as a model 

for a subduction zone with a dipping high-velocity slab 

corresponding to the subducted plate (for example, 

Engdahl and others, 1982). Ray tracing can also 

improve the location of hypocenters in relation to 

known near-surface faults if the velocity differences 

in the region are sufficiently well known (Engdahl and 

Lee, 1976). In locating regionally or locally recorded 

earthquakes, various approaches have been followed to 

account for lateral variations in velocity. For a 

discussion of this problem and of the theory, practice, 

and application of microearthquake studies, in general, 

see the review by Lee and Stewart (1981). 

Ocean-bottom seismographs 

The deployment of ocean-bottom seismographs 

(OBS's) can significantly improve both the detection 

threshold and the location accuracy of offshore seismicity. However, few OBS's have been deployed for 

long-term earthquake recording because of the costs 

and logistic and instrumental difficulties. Noteworthy 

long-term OBS recording efforts, all but one using long 

cables to transmit data, include: a single station 135 

km off Point Arena, northern California, operated 

from 1966 to 1972 (Sutton and others, 1965; Nowroozi, 

1973); a line of four stations deployed in 1978 and 

extending some 100 km seaward off the Pacific coast 

of central Honshu, Japan (Meteorological Research 

Institute, 1980); an array of five stations around the 

Dos Cuadras oilfield in the Santa Barbara Channel, 

deployed in 1978 (Henyey and others, 1979); and a 

single station in the North Sea between Norway and 

Scotland, deployed in 1980 and linked to the Beryl 

Alpha production platform via buoy-supported radio 

telemetry (Turbitt and others, 1983). To date, most 

OBS's have been deployed for short periods to investigate Earth structure or to accurately locate earthquakes occurring within special study areas during 

intervals of days to a few months. Prothero (1984) 

summarized the state-of-the-art in OBS technology. 

The potential of OBS's for providing data that 

help resolve the tectonic processes responsible for 

offshore seismicity was illustrated in a study of the 

seismicity of the Juan de Fuca plate off western 

Canada (Hyndman and Rogers, 1981). The historical 

seismicity (fig. 19A) shows a broad scatter, about 100 

km wide, from the south end of the Queen Charlotte 

fault to the north end of the Juan de Fuca Ridge. 

Isolated epicenters extend farther offshore and toward 

the British Columbia coast. This scatter could 

represent either location uncertainty or a real distribution reflecting breakup of the lithospheric plates in 

this area. Rogers (1980) revised the locations of many 

of these events from previously published values, some 

by more than 100 km, and reduced the scatter 

24 Earthquake Hazards in the Offshore Environment 

somewhat, but most epicenters have yet to be 

reassessed in a systematic way. 

For comparison, the results from studies utilizing 

temporary OBS arrays are superimposed on the 

principal offshore tectonic features (fig. 19B). The 

microearthquakes are located in a zone 20 to 30 km 

wide that closely follows the plate boundaries defined 

from other geophysical and geologic data. The 

estimated accuracy of the epicenters is better than 5 

km for most events within or near the OBS arrays and 

10 km or greater for many of the most distant 

events. The largest event (M=3.8) detected by the OBS 

arrays, which was located on the Revere-Dellwood 

Fracture Zone (fig. 19B), was also recorded by permanent land-based seismographs. The routine location of 

this earthquake by the Earth Physics Branch of the 

Canadian Department of Energy, Mines and Resources 

was about 20 km northwest of the epicenter based only 

on OBS data, and the preliminary routine location by 

the U.S. Geological Survey was about 50 km northeast 

(Hyndman and others, 1978). 

These results strongly suggest that the historical 

earthquakes actually occurred on or near the plate 

boundaries but that significant mislocation scatter and 

bias were introduced by the poor quality of the 

historical seismic data and the general difficulties 

associated with reliably locating earthquakes recorded 

by an azimuthally limited distribution of stations. The 

durations of OBS deployments-a maximum of 10 days 

in any one array configuration-were too short, 

however, to record a sample of earthquakes including 

the higher end of the magnitude range and thus to 

prove that all significant events in the region occur on 

or near these mapped sea-floor features. 

As an alternative to OBS's, buoyed hydrophones, 

or so no buoys, can be deployed to investigate offshore 

seismicity. For example, sonobuoys have been used to 

study microearthquakes occurring along the Blanco 

Fracture Zone and the Gorda Ridge off the coast of 

Oregon (Johnson and Jones, 1978; Jones and Johnson, 

1978). 

The use of the hypocenter-location (or 

relocation) techniques, described in the preceding 

subsection, and, where feasible, the deployment of 

OBS or sonobuoy arrays can significantly improve our 

understanding of historical seismicity patterns and the 

behavior of active fault zones for the purpose of 

seismotectonic modeling in regions of offshore 

petroleum development. 

Geologic record 

The geologic record, where well preserved and 

decipherable, may disclose evidence of previous 

earthquake activity and currently active tectonic 

processes that simply are unknown or unresolvable 

from historical information. The late Quaternary 

geologic record (approximately the past several 

hundred thousand years) and, especially, that of Holocene time (approximately the past 10,000 years) is a 

critical supplement to the historical seismic record. It 

affords a glimpse of earthquake activity over a time 

interval orders of magnitude longer than recorded 

history and thus provides a framework within which to 

interpret the long-term significance of historical 
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Figure 19. Seismicity off southern British Columbia, Canada, in the region of interaction between the Juan de Fuca, Explorer, and North American 

plates. ZOO-m bathymetric contour marks edge of the Continental Shelf. .!!J Epicenters of historical earthquakes (from Rogers, 1980); dashed 

polygons are seismic-source zones employed in regional probabilistic-ground-motion mapping (Basham and others, 198Z). Earthquake source zones: 

CSM, Coast Mountains; JFE, Juan de Fuca-Explorer; NBC, Northern British Columbia; NVI, Northern Vancouver Island; QCF, Queen Charlotte fault; 

SPT, Sandspit. _!!, Microearthquakes located with data recorded by three separate temporary arrays (squares, open triangles, solid triangles) of oceanbottom seismographs (from Hyndman and Rogers, 1981, fig. 5). Dots, epicenters with estimated location uncertainties of less than 10 km; circles, 

epicenters with larger uncertainties. Magnitudes range from O.Z to 3.8. Arrays operated for about 6 to 10 days. Spreading centers (double lines) and 

transform fault zones (single lines) separate Pacific plate from Explorer and Juan de Fuca plates. Thrust fault, with sawteeth on upper plate, 

separates North American plate. 
patterns of seismic activity and quiescence (for 

example, Allen, 1975). 

Virtually all earthquakes occur as ruptures on 

preexisting faults. Accordingly, the identification of 

faults showing evidence of geologically recent movement is one approach in determining where future 

earthquakes are likely to occur. Recency and 

frequency of movement can be used as indicators of 

the degree of past fault activity and, by extrapolation, 

of the future behavior of a fault. For example, a fault 

that has ruptured repeatedly during Holocene time is 

more likely to generate a significant earthquake during 

the lifetime of an offshore structure than is one that 

has not slipped during this time. In many, but not all, 

circumstances, recency of movement on an individual 

fault can be determined from the age of the youngest 

displaced geologic deposit or judged from the geomorphic appearance of the fault, provided that such 

evidence is preserved. 

The identification of faults as potential 

earthquake sources has been particularly successful in 

California (for example, Jennings, 1975; Wesson and 

others, 1975) and in the adjacent offshore area (for 

example, Greene and others, 1973; Ziony and others, 

1974; Buchanan-Banks and others, 1978; Yerkes and 

Lee, 1979b). Contributing to this success are the 

following factors: the earthquakes are typically 

shallower than 15 or 20 km; surface faulting has been 

observed, when searched for, for most onshore earthquakes M>6.0 and many of M=5.0-5.9; and the rates of 

movement for many faults exceed the rates at which 

natural geomorphic processes obscure surficial 

evidence of faulting. Mapping of active offshore 

faults has also been successfully pursued elsewhere off 

the west coast of the United States and Canada (for 

example, Hyndman and others, 1979; Carlson and 

others, 1985). 

The methods and tools for fault mapping in the 

marine environment (Sieck and Self, 1977; Ploessel, 

1978) differ from those employed on land (Sherard and 

others, 1974; Slemmons, 1977). High-resolution 

seismic (acoustic)-reflection profiling is the most 

widely used technique for the identification of active 

offshore faults. The utility of the technique is 

illustrated by a seismic-reflection profile (fig. 20) 

across the Nootka fault zone, the seismically active 

strike-slip boundary between the Explorer and Juan de 

Fuca plates off Vancouver Island. This zone consists 

of a central core of active faults that offset the ocean 

floor, flanked by buried faults that are no longer 

active (fig. 2o_m. A recently developed tool of great 

potential is digitally rectified side-scan sonar, which 

provides undistorted, detailed plan-view images of the 

sea floor comparable in quality and utility to aerial 

photographs of the land {Clifford and others, 1979). 

Information pertaining to the frequency and 

maximum size of earthquakes within a particular 

region or on an individual fault can also be extracted 

from the recent geologic record in some 

circumstances. For example, detailed geologic 

investigations of active onshore faults in the Western 

United States have yielded approximate dates and, in 

some places, amounts of fault displacement for large 

prehistoric earthquakes on faults within several 

tectonic provinces (for example, Bonilla, 1973; 

Machette, 1978; Sieh, 1978a; Swan and others, 1980). 

26 Earthquake Hazards in the Offshore Environment 

Estimates of the size of prehistoric earthquakes have 

been derived from measured prehistoric fault displacements, using the empirical relations between 

magnitude and displacement derived for different 

types of faults (for example, Slemmons, 1977; Bonilla 

and others, 1984) or from comparison of the geologic 

effects preserved from a prehistoric event with those 

from a documented, historical shock (for example, 

Sieh, 1978a). 

An example illustrating the value of the geologic 

record is the San Andreas fault in central California 

(Wallace, 1968). On the Carrizo Plain, a series of four 

small stream channels (A-D, fig. 21) are offset right 

laterally 8 to 12 m along a 200-m length of the fault. 

These offsets are attributed to slip associated with the 

1857 earthquake, which had an estimated magnitude of 

7 .9. One of the channels (D, fig. 21) also exhibits two 

earlier comparable offsets (D', D"), features suggesting 

the recurrence of earthquakes of M rv8 on this segment 

of fault. The dates of these pre-1857 offsets are not 

known, however, an average recurrence interval of 240 

to 450 years has been determined for major earthquakes on this segment of the San Andreas fault from 

knowledge of the average slip rate on the fault and the 

amount of slip in prehistoric earthquakes (Sieh and 

Jahns, 1984). A second example, from the zone of 

surface rupture of the 1971 San Fernando, Calif., 

earthquake (Bonilla, 1973), illustrates geologic 

evidence of a prehistoric episode of dip-slip faulting 

(fig. 22). The earlier faulting was determined to have 

occurred about 200 years ago, by radiocarbon dating of 

a piece of wood buried in the debris that accumulated 

at the base of the scarp formed by prehistoric 

faulting. It is inferred that an earthquake larger than 

the M=6.6 shock in 1971 was associated with this 

earlier faulting because the prehistoric vertical 

displacement exceeds that observed in 1971. 

In the United States and Canada, where the 

written history practically nowhere encompasses the 

complete cycle of stress buildup and release associated 

with a large earthquake, the geologic record has 

provided most of what is known about the repeat time 

and size of large earthquakes on onshore faults. In 

contrast, marine geologic evidence pertaining to the 

date and size of offshore earthquakes has yet to be 

recovered and applied. In coastal regions of rapid 

plate convergence, episodes of vertical crustal deformation accompanying large earthquakes may be 

preserved in the coastal geology, for example, in a 

sequence of elevated marine terraces (Plafker, 1969). 

Although it might be hoped that the geologic record 

would permit inferences about the frequency and 

maximum size of large coastal shocks related to dipslip faulting, such inferences are difficult not only 

because an adequate Quaternary geologic record is 

rarely preserved but also because the patterns of 

deformation witnessed in historical earthquakes are 

complex and variable. 

The imprint in the geologic record from a single 

earthquake typically is relatively small and difficult to 

resolve. Thus, detailed fault investigations more 

commonly yield rates of cumulative fault displacement 

derived from a series of several shocks (for example, 

Sharp, 1981), rather than time intervals between 

successive large shocks. From cumulative displacement rates, the average rate of earthquake occurrence 
NORTHWEST SOUTHEAST 3 

Sea floor 

4 

(f) 

0 

z 

0 

u 

UJ 

(f) 

z 5 

A UJ 3 

~ 

f-

_J 

UJ Sea floor 

> 

<t 

a: 

f4 

5 

0 10 20 30 40 50 

B DISTANCE, IN KILOMETERS 

Figure ZO. Seismic-reflection profile (vertical section) across the Nootka fault zone off Vancouver Island (from 

Hyndman and others, 1979, figs. 7, 8). ~ Reflection profile showing basement uplift of several hundred meters, 

active faults offsetting the sea floor, and formerly active faults offsetting buried, older geologic horizons. .!!, 

Interpretation of reflection profile. Vertical scale is two-way traveltime for reflected seismic waves. In the 

sedimentary column, 1 s of traveltime corresponds to a distance of about 1 to 2 km. Vertical exaggeration, x16. 

can be estimated by relating earthquake magnitude to 

fault offset through seismic moment (Anderson, 

1979). Like magnitude, seismic moment is a measure 

of earthquake size; but unlike magnitude, it can be 

determined from geologic data, namely, the slipped 

area and average displacement on the fault, as well as 

from seismograms (see supplementary section below 

entitled "Intensity, Magnitude, and Seismic 

Moment"). Estimates of occurrence rates require 

assumptions with regard to the maximum size of 

earthquakes and the fraction of displacement that 

occurs aseismically, that is, without association with 

earthquakes, as in creep movement (Wallace, 1970). 

Similarly, on a larger scale, estimates of occurrence 

Assessment of Earthquake Potential 27 
Figure 21. Stream channels offset by right-lateral strike-slip displacement on the San Andreas fault, 

Carrizo Plain area, Calif • .!!:_, Photograph by R.E. Wallace, U.S. Geological Survey. ~ Sketch map of 

foreground of figure 21~ showing amounts of offsets (in meters) for channels of various ages (from 

Wallace, 1968, fig. 8). Youngest, least incised channels record amounts of offset from individual 

prehistoric earthquakes. Offset channels A-D attributed to slip from 1857 earthquake; D' and D" 

indicate offsets from previous earthquakes. Pair of arrows indicate direction of relative movement 

along fault. 

28 Earthquake Hazards in the Offshore Environment 
NORTH 

METERS 

' ' ' ' ' ', 

Gray sand (probably ', 

artificial fill) 

SOUTH 

8 METERS 

Figure 22. Geologic section of wall of trench excavated across fault that slipped during the 1971 San Fernando, 

Calif., earthquake (from Bonilla, 1973, fig. 10). 1971 rupture offset sand bed (C) and overlying wedge of gravelly 

sand (D) along fault A. Earlier episode of reverse slip greater than 1 m along fault B is inferred from absence of 

sand bed north of fault B. Wedge of gravelly sand is interpreted as colluvium and collapsed bedrock deposited 

against scarp formed by earlier episode of slip on fault B. Age of materials near base of wedge is known to be 100 

to 300 years from radiocarbon dating of a buried piece of wood. 

rates for earthquakes occurring on faults constituting 

major plate boundaries can be derived from long-term 

rates of relative plate motion. 

There are various approaches to estimating the 

size of the maximum possible earthquake in a region, 

all of which have limitations and problems that lend 

uncertainty to the estimates. For recognized faults, 

the most common approach deduces maximum possible 

magnitude from the mapped fault length or some 

fraction (commonly one-half) of that length, using 

empirical relations between magnitude and length of 

faulting for various styles of faulting (for example, 

Slemmons, 1977; Bonilla and others, 1984). The use of 

fault area in place of fault length has also been advocated (Wyss, 1979; Singh and others, 1980). Empirical 

relations between dimensions of faulting and 

earthquake magnitude are founded both on field 

measurements of surface faulting and on seismologic 

and geodetic inferences concerning the extent of 

subsurface rupturing. In the absence of recognized 

faults, estimates of the maximum possible earthquake 

within a region are generally based on one or more of 

the following considerations: historical seismicity of 

the region, tectonic deformation of young geologic 

units within the region, geologic and tectonic relations 

of the region to the surrounding area, and analogs with 

regions having similar geologic and tectonic characteristics. By considering as many of the above factors as 

is possible, the most reliable estimate is sought. 

Finally, the geologic record provides a 

framework for assessing the significance of apparent 

seismic quiescence in a region with a short recorded 

history. The historical absence of seismicity need not 

imply future quiescence unless some other evidence 

indicates the absence of tectonic deformation during 

Holocene or late Quaternary time. For example, 

historical quiescence in the vicinity of a fault that 

clearly exhibits repeated Holocene displacement is 

evidently a transient feature; significant earthquakes 

are likely to occur in the future, possibly even the near 

future. 

Seismotectonic models 

A seismotectonic model describes the 

earthquake-generation processes and structures within 

a given region and projects the spatial, temporal, and 

magnitude distribution of future earthquakes in that 

region. The accuracy or reliability of the model 

depends directly on the degree of knowledge 

concerning processes and structures, as gained from 

investigations of the type described previously. A 

seismotectonic model may be quite accurate, at least 

to the degree required for the evaluation of 

earthquake hazards pertinent to engineering design. 

For example, a fault zone may have well-documented 

historical and current seismicity; its spatial extent 

may be clearly defined; and the average rate of significant earthquakes may be estimable from historical 

seismicity and geologic evidence. For such a zone, the 

model may also specify additional characteristics, such 

as the typical dimensions and attitudes of fault breaks, 

and various source parameters of expected 

earthquakes, such as maximum magnitude or seismic 

moment, which is directly proportional to the product 

of the slipped area on the fault and the average 

displacement over the slipped area, (see supplementary 

Assessment of Earthquake Potential 2 9 
section below entitled "Intensity, Magnitude, and 

Seismic Moment"). 

Conversely, a seismotectonic model may be 

uncertain or tentative. The earthquake-source zone 

may be delineated by only a diffuse pattern of 

historical and recent earthquakes, with little 

documented evidence for their association with 

geologic features or with little knowledge of the 

causative tectonic forces. In addition, the available 

data may be insufficient to estimate the average 

recurrence rate of significant earthquakes or the 

magnitude of the largest possible earthquake. Nevertheless, such earthquakes must be accounted for in 

assessments of earthquake potential and associated 

effects, and the seismologist may have little choice 

but to draw an approximate boundary around the 

source zone and to assume that similar earthquakes 

will occur in the future, either randomly throughout 

the zone or at the estimated epicenters of the more 

significant historical events. 

Seismotectonic models may be formulated in 

various ways, depending on the purpose of the hazard 

assessment in which they are to be used. Different 

levels of knowledge concerning geologic hazards are 

needed at different stages and by different parties in 

the development of offshore petroleum resources. In 

deciding whether to hold a lease sale, a governmental 

body may want a national or broad regional overview 

of the relative geologic hazards among various 

proposed lease areas. For selecting tracts to be 

offered in a lease sale and for bidding on tracts in the 

sale, more detailed knowledge is required to identify 

the potential hazards associated with individual 

tracts. Finally, for the siting and design of offshore 

production facilities and for the governmental regulation of operations, highly detailed knowledge of the 

specific hazards affecting a given site is needed. 

The seismotectonic model prepared for a large 

region generally does not have the same degree of 

information and specificity needed for detailed 

application to a specific site or small area. For 

example, evaluation of the relative ground-shaking 

hazard for a continental shelf requires ground-motion 

estimates to be derived on a gross regional basis, so as 

to provide information applicable to typical sites 

throughout the entire shelf region. For such a broadscale evaluation, not all potential earthquake sources 

that may affect any site throughout the region can be 

investigated in the same degree of detail that is 

required for assessing hazards to a critical facility at a 

specific site. Thus, for regional evaluations, seismotectonic models will be generalized, as a rule. 

In the following subsections, we illustrate these 

concepts with discussion of seismotectonic models for 

parts of the active west and passive east margins of 

North America. 

Active western margin 

The seismicity of western Canada has recently 

been modeled to derive new probabilistic-groundmotion maps of Canada (Basham and others, 1982). 

For this application, broad earthquake source zones 

were defined. The Queen Charlotte transform fault 

and the system of spreading-ridge and transform-fault 

30 Earthquake Hazards in the Offshore Environment 

segments constituting the boundary between the 

Pacific plate and the Juan de Fuca and Explorer plates 

are selected for illustration (fig. 19A). 

The Queen Charlotte fault -zone defines the 

continental boundary along the west coast of the 

Queen Charlotte Islands. This area has virtually no 

continental shelf; instead, an almost continuous, but 

irregular, slope extends from the 1-km-high mountains 

on the islands to the 3-km-deep sea floor. A 25-kmwide terrace, bounded by parallel fault scarps, interrupts the continental slope at a depth of 2 km and 

represents the fault zone. This zone is the probable 

origin of most historical earthquakes along the west 

coast of the Queen Charlotte Islands; low-magnitude 

events, occurring beneath the continental edge of the 

terrace, suggest that the landward fault is currently 

the main active element of the fault zone (Hyndman 

and Ellis, 1981). The Queen Charlotte fault is modeled 

as a long narrow source zone (QCF), the south end of 

which is shown in figure 19 A. 

The JFE source zone Uig. 19A) encloses the main 

cluster of historical earthquakes associated with the 

interaction of the Juan de Fuca and Explorer plates 

with the Pacific plate. Although most of these earthquakes probably occurred on the spreading ridges and 

fracture zones shown to be active by the OBS experiments discussed in the subsection above entitled 

"Ocean-Bottom Seismographs" (Hyndman and Rogers, 

1981), gross characterization of this zone is adequate 

for the purpose of constructing a national map 

depicting probabilistic ground motion, which will be 

used primarily for earthquake-resistant design of 

common structures, principally onshore. 

The rates of earthquakes in the QCF and JFE 

source zones are given by cumulative magnituderecurrence curves (fig. 23). The best-fit line was 

computed by using a maximum-likelihood method 

(Weichert, 1980) that has been developed to account 

for estimates of earthquake rates for unequal observation periods and for various magnitudes. Upper limits 

to possible magnitudes are assumed for the two zones; 

this assumption leads to the downward curvature to 

zero rate at large magnitudes. 

A maximum magnitude of 8.5 is assumed for the 

QCF source zone, which would be generated by rupture 

of most of the length of the fault zone. The largest known earthquake was of M=8.1, which occurred in 

1949 with an epicenter west of the Queen Charlotte 

Islands. In the JFE source zone, the seismicity data 

alone may provide some indication of the maximum 

magnitude. Although the estimated rate for M ~ 6.5 

earthquakes is about one every 10 years, in fact-;ihere 

have been no earthquakes of M >6.5 in the estimated 

period of complete reporting starting in 1917. The 

faults in the JFE source zone have a maximum length 

of about 100 km and could have a vertical extent of 

about 10 {m; thus, the maximum fault area is about 

1,000 km . Relations between magnitude and fault 

area (Kanamori and Anderson, 1975; Singh and others, 

1980) suggest that the maximum possible earthquake is 

of Mrv 7 .0, consistent with the historical data. - For estimating the probability of ground shaking 

on a regional scale (see next section), earthquakes 

within a seismic-source zone are typically assumed to 

occur randomly over time unless the historical or 

geologic record indicates nonrandom occurrence. A 
100 

a: 

<( 

L.U 10 >-

a: 

L.U 

0.. 

(/) 

L.U 

~ 1.0 <( 

::> 

0 

I 

...... 

a: 

<( 

L.U 0.1 

u.. 

0 

a: 

L.U 

CXl 

~ 0.01 ::> 

z 

3 4 5 6 7 8 

MAGNITUDE 

Figure Z3. Cumulative magnitude-recurrence relations 

for the Queen Charlotte fault (QCF, dots) and Juan de 

Fuca-Explorer (JFE, circles) earthquake-source zones 

of figure 19A (from Basham and others, 198Z, figs. 15, 

17). Data points are for observed number of 

earthquakes per year larger than a given magnitude. 

Curves are fitted to data by assuming upper-bound 

magnitudes of 8.5 and 7 .0, respectively. 

nonrandom component can be incorporated into the 

seismotectonic model for a seismic gap, that is, a 

region in which repeated large earthquakes have 

occurred in the past but not recently, and thus in 

which a large shock may be expected in the near 

future. An example is the Gulf of Alaska region (fig. 

24). 

Sykes (1971) identified a seismic gap between the 

aftershock zones of the 1964 Prince William Sound and 

the 1958 Fairweather fault earthquakes. No major 

earthquake has occurred in this gap since two M=8.1 

earthquakes ruptured the gap in 1899 (McCann and 

others, 1980). Lahr and others (1980) demonstrated 

that this gap was only partly filled by the 1979 St. 

Elias earthquake (M=7.6). Lahr and Plafker (1980) 

concluded that the-gap-filling rupture(s) in a future 

major earthquake(s) would most likely occur on the 

main north-dipping thrust faults between Kayak Island 

and Icy Bay. If the Pacific and North American plates 

have been converging at the rate of 5 cm/yr since the 

1899 sequence and if movement between these two 

plates has not been accommodated in plastic 

deformation, enough elastic strain has already 

accumulated to cause a potential slip of 4 m. If this 

amount of slip occurred in a single earthquake, it 

would generate an event as large as M=8 that would 

likely fill the remainder of the gap. -

Thus, a seismotectonic model of the Gulf of 

Alaska region must account for the fact that the next 

great earthquake is more likely to occur in this gap 

than at any other point on the border of the northern 

gulf. This possibility, however, should not be 

considered to the exclusion of the fact that significant 

earthquakes will continue to occur on other major 

tectonic features bordering the gulf. 

Passive eastern margin 

The continental margin south of Newfoundland 

provides a good example of the difficulties in 

constructing reliable seismotectonic models to 

describe future earthquake occurrence along the 

passive eastern and Arctic margins of North America 

(Basham and Adams, 1982). 

The seismicity of Newfoundland and the adjacent 

shelf (fig. 25) is probably incomplete at the M=5 level 

before the mid-1950's, and at the M=4 level before 

1965. With the present seismograph network, few 

earthquakes of M<4 can be located in the offshore 

areas of figure 25."" 

The most significant earthquake in this region 

was the 1919 Grand Banks M=6.5 event near the edge 

of the continental slope-at the mouth of the 

Laurentian Channel (see also fig. 15). This earthquake 

had aftershocks as large as M=6.0. M "'5 earthquakes 

occurred also at the mouth Of the Laurentian Channel 

in 1951, 1954 (two events), and 1975, and northeast of 

Newfoundland in 1922 (the location of this last event is 

uncertain by at least 100 km). The M=4 earthquakes, 

and the few smaller events that have been located, are 

scattered throughout the shelf area. 

Although the epicenters at the mouth of the 

Laurentian Channel (fig. 25) are uncertain by tens of 

kilometers, the observed scatter is not simply the 

result of random errors in locating shocks, originating 

at a common point, on the basis of arrival times of 

seismic waves recorded at distant seismographs. 

Instead, seismologic evidence indicates that these 

events are occurring within an extended source zone, 

which for purposes of probabilistic calculations of 

ground motion (see subsection below entitled "Areal 

Evaluation") is modeled as the quadrilateral shown in 

figure 25. Although this source-zone model is 

currently favored, several key questions still remain to 

be answered. Can further seismologic studies of these 

earthquakes and other geologic and geophysical studies 

of the region lead to an understanding of the tectonic 

processes and active faults that generated the 1929 

earthquake? Do similar geologic features exist in 

other areas of the shelf that could cause similar 

earthquakes? Do small earthquakes scattered 

throughout the shelf provide evidence for the 

existence of these features? Is there evidence in the 

unconsolidated sediment that could reveal the 

recurrence period of such events in Holocene time? A 

speculative source-zone model that assumes large 

earthquakes occur uniformly along the eastern 

Canadian margin over long periods of time has been 

studied by Basham and others (1983). 

Gravity, aeromagnetic, and seismic-reflection 

data have been used to delimit the faulting mapped in 

figure 25. King and MacLean (1970) and King (1980) 

associated the 1929 and other earthquakes with the 

fault system that intersects the Laurentian Channel 

and extends eastward, possibly linking with the 

Newfoundland Fracture Zone. Various studies (for 

example, Sykes, 1978; Stewart and Heimberger, 1981) 

Assessment of Earthquake Potential 31 
150° 140° 

ALASKA CANADA 

60° 

500 1000 KILOMETERS 

Figure 24. Aftershock zones of M >7.3 earthquakes since 1938 in the Gulf of Alaska, showing the Yakataga and 

Shumagin seismic gaps (modified from Sykes, 1971, fig. 4, and Lahr and others, 1980, fig. 1). 1979 St. Elias 

earthquake sequence, which occurred on east edge of the Yakataga gap, only partly filled gap between 1958 and 

1964 rupture zones. 

have suggested that the seismicity of the Atlantic 

continental margin is related to such linear features as 

fracture zones and seamount chains which are nearly 

normal to the margin. Stein and others (1979) 

suggested that the stresses due to glacial unloading are 

sufficient to reactivate old faults parallel to the 

margin. The available seismic profiling has not shown 

clear evidence of recent offsets in the sea-floor 

faulting within this region; high-resolution seismic 

profiling of the uppermost 20 m of sediment would be 

required to investigate youthful faulting at these sites. 

The recurrence rate of 1929-size earthquakes in 

the Laurentian Channel source zone (fig. 25) is problematic. On the one hand, the historical frequency of 

shocks suggests the occurrence of an earthquake as 

large as that in 1929 about once in 300 years (Basham 

and others, 1982, 1983). On the other hand, a recurrence interval of at least 100,000 years has been 

inferred from marine geologic investigations (Piper 

and Normark, 1982). This inference is based on the 

assumption that an earlier shock of comparable 

magnitude would have caused widespread sea-floor 

failure, similar to that which occurred in 1929 (fig. 

15). High-resolution seismic-reflection profiling and 

bottom coring reveal no such disturbance in sediment 

at least 100,000 years old. 

Similar problems exist in constructing accurate 

seismotectonic models for elsewhere along the passive 

eastern and Arctic margins of North America. These 

32 Earthquake Hazards in the Offshore Environment 

problems may be most tractable in regions that have 

undergone a large historical earthquake, such as 

Charleston, S.C. (1886, estimated M=6.9-7 .3), Baffin 

Bay (1933, M=6.7), and the Beaufort Sea (1920, M=6.2); 

however, only a single large shock has occurred in each 

of these regions. In contrast to the western margin, 

where numerous large historical events delineate 

active tectonic features, a single large earthquake in 

each of these eastern regions represents a very limited 

data base on which to found a seismotectonic model. 

Although detailed geologic and geophysical investigations may possibly reveal the geologic features 

responsible for such individual earthquakes, similar 

seismogenic features elsewhere along the margin may 

go unnoticed because they have not genera ted a large 

earthquake during historical time and thus have not 

become a candidate for detailed investigations. The 

origin of the 1886 Charleston earthquake has been 

studied intensively (Gohn, 1983; Hays, 1983, but 

neither the tectonic structure on which this shock 

occurred nor the implication of this shock for similar 

events elsewhere along the Atlantic margin has yet 

been clearly resolved. 

Induced seismicity 

An additional consideration in assessing earthquake potential is the possibility that the activities 
50° 

48° 

46° 

200 400 KILOMETERS 

EXPLANATION 

EARTHQUAKE EPICENTERS 

e M~6.0 

• M 5.0 to 5.9 

• M 4 .0 to 4.9 

• M 3.0 to 3 .9 

M<3 .0 

Figure ZS. Continental Shelf south of Newfoundland, showing earthquake epicenters from 1929 to 1980, ZOO- and 

2,000-m isobaths, seismograph stations (triangles), and offshore basement faults (dashed where uncertain) (from 

Basham and Adams, 1982, fig. 3). Short-dashed fault traces west of long 55.5° W. were observed on seismicreflection profiles (King and MacLean, 1970); long-dashed traces east of long 55.5° W. were inferred from 

_magnetic and gravity studies (Lefort and Haworth, 1978). Quadrilateral at mouth of the Laurentian Channel is 

earthquake-source zone used in regional probabilistic ground-motion mapping. 

associated with petroleum recovery may significantly 

alter the natural rate of earthquake activity in the 

vicinity of a critical offshore facility. A common 

technique for stimulating secondary recovery of oil 

involves injection of fluids into the reservoir under 

high pressure. Evidence from various geologic and 

tectonic settings indicates that if the injection wells 

bottom in the proximity of existing faults, significant 

seismic activity can be triggered, even if earthquakes 

were previously rare or even unknown. 

A graphic example of seismicity induced by fluid 

injection was the earthquake sequence near Denver, 

Assessment of Earthquake Potential 33 
Colo., during the 1960's (Healy and others, 1968; Hsieh 

and Bredehoeft, 1981). Injection of waste fluids into a 

3. 7-km deep well bottoming in fractured Precambrian 

crystalline rocks triggered earthquakes as large as 

M=4.8. This sequence of earthquakes appears to have 

no parallels in the local seismic history. The earthquake hypocenters migrated away from the well over 

time and extended horizontally to a distance of about 

7 km by 1967, 5 years after injection started. Focal 

mechanisms indicate right-lateral slip within a buried 

fault zone parallel to the trend of the earthquake 

hypocenters. The mechanism offered to explain the 

triggering is the reduction of frictional resistance on 

preexisting faults that allows seismic slip to occur in 

response to ambient tectonic stresses. The injected 

fluid partially offsets the normal stress locking the 

fault, by an amount equal to the pressure of the fluid, 

and thus lowers the frictional resistance to sliding. 

After the Denver study, a classic experiment to 

examine more closely the role of fluid pressure in the 

triggering of earthquakes was undertaken in the 

Rangely oil field in Colorado, where waterflooding 

operations in the reservoir were triggering earthquakes 

(Raleigh and others, 1976). By the injection and backflow of fluid in four 2-km-deep injection wells 

straddling a subsurface fault in the reservoir rock, 

pore pressures in the vicinity of the fault were 

controlled, and the earthquake activity sequentially 

was turned off, on, and off again. During the first 

year of this experiment, fluid was injected, and earthquakes occurring within 1 km of the bottom of wells 

were recorded at the average rate of 28 per month 

(fig. 26). In November 1970, injection was stopped, 

and withdrawal of fluid caused the earthquake activity 

to decrease to about one event per month. Injection 

was resumed in May 1971; however, the probabilistic 

bottom-hole pressure did not substantially exceed the 

critical pressure predicted for earthquake triggering 

until November 1972, when the rate increased to about 

20 events per month. In May 1973, injection 

terminated, the wells were backflowed, and 

earthquakes within 1 km of the wells abruptly 

terminated. The largest earthquakes recorded during 

the experiment were of M rv 3.0. The earthquakes 

occurred on a minor fault, not expressed in the surface 

geology-one that would have gone undetected were it 

not within an oil field. Focal-mechanism solutions for 

the earthquakes indicate strike slip. 

Other places where earthquake activity may 

have been related to hydrocarbon production include 

Alberta (Milne, 1970; Wetmiller, 1985) and Oklahoma 

and Texas (Rogers and Malkiel, 1979). Induced 

seismicity has also been linked to high-pressure fluid 

injection in the vicinity of a fault in association with 

the hydraulic mining of salt in western New York 

(Fletcher and Sykes, 1977). Thus, evidence to date 

documents that significant earthquakes, at least 

approaching M=5, can be triggered on existing faults 

by fluid injection. There is no reason to believe that 

an earthquake of M>5 could not be triggered inadvertently on a sufficiently large fault favorably oriented to 

the regional tectonic stress field. 

200r-------------------------------------------------------------~5000 

150 4000 L.L.i (/) 0:: I UJ ::> u ~ (/) 

~ (/) ~ 

::> UJ UJ 

d 0:: 0:: a_ I ~ 

I- 0:: ::> 0:: -a ~ 100 3000 ~(/) UJ O::o:: u.. UJ UJ 

0 (/)a_ 

0:: UJ(/) 

UJ o::Cl 

co >- z ~ ....J ::> ::> Io z ~a_ 

50 2000 oz 

~-

0 1000 

0 N 0 J FM AM J J AS 0 N D J F M AM J J AS 0 N D J F M AM J J AS 0 N D J F M AM J J AS 0 N D J F M AM 

1969 1970 1971 1972 1973 1974 

- Fluid -1-- Fluid --1 Fluid 

injection withdrawal injection 

Figure Z6. Monthly number of earthquakes (bars) and monthly reservoir pressure (curve) at the 

Rangely, Colo., oil field during earthquake-control experiment (from Raleigh and others, 1976, 

fig. 7). Solid bars indicate earthquake activity within 1 km of bottom of experimental wells; 

open bars show number of shocks occurring elsewhere in the field. Dashed line shows predicted 

critical fluid pressure required to trigger earthquakes. 

34 Earthquake Hazards in the Offshore Environment 
EVALUATION AND MITIGATION OF EARTHQUAKE 

HAZARDS 

The evaluation of earthquake hazards involves 

the prediction of effects from earthquakes whose 

potential for occurrence is defined in a seismotectonic 

model. Should these estimated effects appear to be 

potentially hazardous, mitigation measures can be 

invoked. Effective mitigation strategies include 

avoidance, accommodation, and emergency planning. 

First, structures can be sited to avoid or limit 

exposure to hazardous earthquake phenomena. Second, 

structures can be designed and built to accommodate 

the estimated effects of earthquakes. Third, response 

planning for an earthquake emergency can minimize 

the adverse consequences of failure in a structure or 

system and can promote rapid recovery. 

Which of these three strategies would be most 

effective in a particular situation depends on the 

specific type of hazard. For example, avoidance may 

be a feasible and effective measure for mitigating the 

hazards associated with surface faulting or seismically 

induced sea-floor failures, whereas accommodation 

would likely be chosen as the preferred strategy for 

minimizing potential damage from strong shaking. 

Emergency response planning is an effective mitigation measure when advance warning of a potential 

hazard can be provided, as is done now for tsunamis. 

With advance warning, people can vacate dangerous 

localities and structures, hazardous operations can be 

shut down, and facilities can be protected. Were 

reliable methods of predicting earthquakes available 

today, the importance of emergency planning as a 

mitigation measure would be greatly increased over 

what it is now. 

As discussed in the preceding sections, earthquake potential rarely can be assessed with a high 

degree of certainty. Likewise, earthquake effects are 

rarely predictable with the level of precision and 

certainty that engineers and planners desire. 

Accordingly, the mitigation of earthquake hazards 

requires judgments: scientific judgments in estimating 

effects, engineering judgments in adopting mitigative 

measures, and socioeconomic judgments in deciding on 

acceptable risk. 

Surface faulting 

There are two effective strategies for mitigating 

potential hazards from surface faulting. First, structures of limited dimension can be sited selectively to 

avoid active fault zones where surface rupture is 

likely. Second, in circumstances where a structure 

must cross a zone of potential surface rupture, either 

because of the extended size of the structure or 

because of some critical siting constraint, the 

structure can be designed to accommodate the effects 

of surface faulting. For example, where it crosses an 

active fault, a pipeline can be oriented to minimize 

the strains induced by fault movement, and the mode 

of pipeline burial or anchoring in the vicinity of the 

fault can be chosen to distribute the strains induced by 

fault movement over a length of the pipeline greater 

than the width of the shear zone in the ground 

(Kennedy and others, 1977). In addition, automatic 

shutoff valves can be installed to stop the flow and 

limit discharge from a ruptured pipeline. The strategy 

of avoidance requires the ability to identify active 

surface faults. The principles and methods for identifying active faults were discussed in the preceding 

subsection entitled "Geologic Record." The strategy 

of accommodation requires not only that active 

surface faults be identified but also that the style of 

future faulting be predicted. 

Surface faulting is a complex phenomenon, 

important aspects of which are poorly documented and 

physically not well understood. Numerous observations 

of the patterns and amounts of surface fault displacements have been made after historical earthquakes. 

These observations provide an empirical basis for 

estimating the amount and type of displacements in 

future earthquakes. However, few well-documented 

eyewitness accounts and no instrumental records of 

faulting have been obtained during an earthquake that 

reveal, for example, when surface ruptures develop in 

relation to the time of strong shaking or how the 

displacement at one point on a fault increases over 

time. Two parties witnessed normal faulting on a 

secondary fault during the 1983 Borah Peak, Idaho, 

M=6.9 earthquake. Both parties reported the scarp 

formed within an interval of a few seconds and after 

the onset of shaking; however, one stated the formation of the scarp coincided with the onset of the 

violent phase of shaking (Pelton, 1984), while the other 

claimed it followed (Wallace, 1984). Bonilla (1970) and 

Slemmons (1977) summarized much of what is 

currently known about the phenomenon of surface 

faulting and provided empirical relations between 

earthquake magnitude, type of fault, and various 

parameters of surface faulting, such as the length and 

width of surface rupture and maximum fault displacement. 

Not all the surface displacement on seismically 

active faults accompanies earthquakes. Some earthquake-generating faults exhibit aseismic slip or creep 

behavior, in which surface slip accumulates slowly, 

commonly not in obvious association with earthquakes 

(fig. 27). Fault creep may occur in discrete episodes, 

characterized by as much as 1 to 2 em of slip within an 

interval of hours or days, or as nearly uniform motion 

at rates as high as 2 to 3 cm/yr (Burford and others, 

1978). The most notable example of creep behavior is 

.....:~ 40 Zw 

Wl-

:::i!:W 30 w::::E 

u-

:5~ 20 a...::::E 

~~ 10 

1969 1970 1971 1972 1973 

Figure 27. Fault-creep displacement measured on 

central section of the San Andreas fault 14 km 

southeast of San Juan Bautista, central California 

(from Nason, 1973, fig. 5). Most creep occurs in 

discrete episodes, typically with 1 to 3 mm of slip. 

Although some episodes correlate with the occurrence 

of nearby earthquakes, many do not. 

Evaluation and Mitigation of Earthquake Hazards 35 
on the San Andreas fault in central California, where 

creep accounts for as much as 3.2 cm/yr of rightlateral slip and locally accommodates all the motion 

between the crustal blocks on either side of the fault 

(Burford and Harsh, 1980). Tectonic creep has also 

been documented on several other faults, including 

normal and reverse faults. 

Most measurements of the surface fault 

displacements associated with large shallow-focus 

earthquakes have been made hours, days, or even 

weeks after the earthquake. In interpreting such 

observations, it is commonly assumed that all the 

measured displacement occurred at the time of the 

earthquake. This assumption has been proved 

appropriate in some events, most conclusively for the 

1971 San Fernando, Calif., earthquake (Savage and 

others, 1975), but incorrect in several other events in 

which repetitive measurements conducted over several 

weeks or months, starting within a few days after the 

earthquake, have documented continuing fault slip. 

Such continuing movement is referred to as afterslip. 

Afterslip may occur in the form of both aseismic slip 

and coseismic slip associated with the aftershocks that 

normally follow large shallow-focus earthquakes. A 

spectacular example of afterslip was associated with 

the 1966 Parkfield earthquake on the San Andreas 

fault in central California. In the year after the M=6.2 

main shock, at least 20 em of afterslip occurredat a 

site near the midpoint of the rupture, far exceeding 

the surface displacement occurring at the time of the 

main shock (Smith and Wyss, 1968). The rate of 

displacement decreased approximately logarithmically 

SOUTHEAST 

over time from about 10 mm/d a few days after the 

earthquake to about 5 mm/mo at the end of a year. 

Substantial afterslip has also been measured after 

several other earthquakes (for example, fig. 28); many, 

though not all, of these have been associated with 

strike-slip faulting. There is also evidence, though 

much more limited, that surface slip may occasionally 

precede earthquakes. It is doubtful, however, that 

foreslip either occurs so frequently or contributes so 

much to the total displacement as does afterslip. 

Until the presence or absence of creep, foreslip, 

and afterslip has been documented in many more 

events incorporating a wide range of fault type, 

tectonic and geologic environment, and earthquake 

size, the contributions of these modes of slip relative 

to the coseismic displacement accompanying a large 

earthquake cannot be confidently predicted. 

The pattern of surface ruptures accompanying a 

large earthquake is typically complex. This pattern is 

generally simplest for earthquakes on steeply dipping 

strike-slip faults and most complex for gently dipping 

reverse faults. Even for strike-slip faults, however, 

the complexities can be numerous (fig. 29). For 

example, in the M=6.5 Imperial Valley, Calif., 

earthquake of 1979, displacement on the 30-km-long 

main fault was predominantly strike slip, with a maximum strike-slip component of 55 to 60 em measured 

during the first day after the earthquake and a 

maximum dip-slip component of about 10 em (Sharp 

and others, 1982). Significant afterslip was observed 

over most of the zone of surface rupture, and by 160 

days after the main shock the maximum amount of 

NORTHWEST 

80~----------~-------------,------------~------------~----------~------------~~ 

70 

~ 

Zen 

UJa: 

~ ~ 60 Q..UJ 

~~ 

au~ 50 ...JUJ 

<(U 

~~ 

0 ~- 40 ~z a:w 

0~ Iw 

wU ;::: :s 30 

~a.. <Cen 

5o 

~ u. 20 

:::>0 

u 

10 

5 10 15 20 25 30 

DISTANCE ALONG IMPERIAL FAULT. IN KILOMETERS 

Figure Z8. Horizontal component of fault displacement as a function of position along the Imperial fault, 

determined for times of 4, 10, 20, and 160 days after the 1979 Imperial Valley earthquake (from Sharp and others, 

1982, fig. 101). Dots, measured values; circles, interpolated or extrapolated values. 

36 Earthquake Hazards in the Offshore Environment 
33° 

Coachella 

0 

116° 

Cillexico 

UNITED STAT~--- - .Q. -

-----MExico- Epicenter* 

15 30 KILOMETERS 

Figure 29. Imperial Valley area, southern California, 

showing surface fault ruptures accompanying the 1979 

Imperial Valley earthquake. Primary displacement 

occurred on the Imperial fault, and subsidiary 

movement on branch faults, most notably those 

composing the Brawley fault zone. Minor displacements were triggered on the Superstition Hills and San 

Andreas faults, which lie well outside the zone of 

seismic-energy release (from Sharp and others, 1982, 

fig. 84). 

strike slip had grown to nearly 80 em (fig. 28). On a 

13-km-long subsidiary fault branching off the main 

trace, the type of slip was different, predominantly 

normal rather than strike slip. The maximum displacement on a single break along this subsidiary fault was 

12 em of oblique slip. Small surface displacements 

were also documented on distant faults, not directly 

connected to the segment of the Imperial fault on 

which the primary rupture occurred. As much as 1 em 

of strike slip was measured on the San Andreas fault 

(Sieh, 1982) and 2.2 em on the Superstition Hills fault 

(Fuis, 1982) at distances of about 60 and 25 km, 

respectively, from the north end of the Imperial 

fault. These displacements on the San Andreas and 

Superstition Hills faults appeared to be aseismic slip 

triggered by the earthquake on the Imperial fault. 

Bonilla (1970) summarized measurements of the 

surface displacement on branch and secondary faults 

as a function of distance from the main fault zone, and 

measurements of the width of the main fault zone and 

the maximum distance to branch and secondary faults 

from the main fault zone as a function of earthquake 

magnitude. 

Although the observational evidence is limited, 

important regularities seem to exist in the patterns of 

faulting between successive earthquakes on the same 

fault system. For example, surface offsets occur on 

preexisting traces of faulting with few exceptions, 

although in many places preexisting traces have been 

recognized only after a new episode of faulting. In 

addition, the sense of dip-slip and strike-slip components at individual points along a fault generally is the 

same from one large earthquake to another, even if 

the rupture zones for the two shocks only partly 

coincide. The ratio between dip-slip and strike-slip 

components, however, appears to vary more (Sharp, 

1982). Thus, despite the complexities in the spatial 

patterns of faulting, important characteristics of 

faulting seem to be predictable on the basis of past 

fault history. 

Empirical correlations of maximum surface fault 

displacement with earthquake magnitude provide a 

basis for predicting the maximum surface displacements for earthquakes of a postulated magnitude. 

Relations between maximum surface displacement and 

earthquake magnitude for strike-slip, reverse-slip, and 

normal-slip faults show that maximum displacement 

increases with magnitude (fig. 30). The considerable 

scatter in the data, however, indicates that considerable uncertainty is attached to such estimates of 

maximum displacement. 

In selecting a design value of fault displacement 

for a structure that must cross an active fault, we are 

generally interested in knowing the most probable 

surface fault displacement for a postulated earthquake 

in addition to the maximum displacement. The amount 

of fault displacement may vary considerably as a 

function of position along the fault trace. The 

maximum displacement may occur near the center of 

the rupture zone, near one of the ends where the 

displacement decreases to zero (fig. 28), or at some 

other intermediate point. Points where the initial slip 

is low may also be sites of relatively high rates of 

afterslip, as was observed after the 1979 Imperial 

Valley earthquake (fig. 28). 

Subsidence and uplift 

Dip-slip faulting in a large shallow-focus 

earthquake causes uplift and subsidence over a broad 

region with dimensions approximating those of the 

buried fault rupture. In spite of its extent, such deformation is of serious consequence to relatively few 

facilities, namely, those offshore and coastal facilities 

whose function or safety would be compromised by a 

permanent change in elevation of the sea floor or 

coastline referenced to sea level. The most effective 

strategy for mitigating these potential adverse consequences is accommodation of the effects of uplift and 

subsidence in the design of facilities. Rarely is it 

Evaluation and Mitigation of Earthquake Hazards 37 
10 

• 

6 

(/) • • a: 6 6 

LlJ 5 0 1-

LlJ • ~ 0 • ~ lh • r-: 

z 

LlJ 0 • 0 0 ~ 

LlJ • u 

<( 

_J • a... 

(/) 0 

0 

1- 1.0 _J 0 

:::> 6 0 <( 

u.. 

LlJ 

u 

<( 0 .5 u.. 

a: 

:::> 

(/) • ~ EXPLANATION 

:::> 

~ FAULT TYPE 

x • Strike slip <( 

~ 6 Normal 

0 Reverse 

0.1 

6.0 7.0 8.0 

MAGNITUDE 

Figure 30. Maximum surface fault displacement as a function of surface-wave magnitude for strike-slip, reverseslip, and normal-slip faults (modified from Bonilla and others, 1984). Curves are weighted least-squares 

regressions of logarithm of displacement upon magnitude. No curve is shown for reverse-fault data alone because 

of the large scatter in that data set. 

possible to avoid this problem by selective siting, 

because, unlike surface faulting, the phenomenon is 

not localized. Moreover, the location of an offshore or 

coastal facility generally is severely constrained by 

geographic, geologic, and hydrographic requirements, 

and so suitable alternative sites that would be free 

from uplift or subsidence commonly are not available. 

Broad-scale vertical deformation is likely to be a 

significant design consideration only for shallow-focus 

earthquakes of M ~ 7.0 with a major component of dipslip faulting. Because coastal and marine facilities 

must be designed to accommodate tides and storm 

waves, a permanent change in elevation of the 

shoreline or in depth of water offshore of less than a 

meter is probably of little consequence in most 

places. To a first approximation, the maximum uplift 

or subsidence may approach half the amount of dip-slip 

displacement on the buried fault. From observations 

of surface faulting and inferences from seismologic 

and geodetic investigations, 2 m is a representative 

value for fault displacement for an M=7 reverse- or 

normal-fault earthquake (fig. 30). -The potential 

consequences of vertical deformation increase 

markedly with earthquake size because both the 

amount of uplift and subsidence and the size of the 

affected area increase with earthquake magnitude. 

The seismotectonic model for a given region 

incorporates judgments on the size, focal depth, and 

sense of faulting of future earthquakes and thus 

38 Earthquake Hazards in the Offshore Environment 

provides a preliminary basis for determining whether 

vertical deformation is a potentially significant earthquake problem. In coastal regions, more direct 

evidence of the potential for vertical deformation may 

be preserved in the late Quaternary history of ancient 

shorelines. For example, along the Santa Barbara 

Channel in southern California, a series of emergent 

marine terraces indicate average tectonic-uplift rates 

as high as 10 mm/yr for the past few tens of thousands 

of years (Lajoie and others, 1982). This area is characterized by active folds and reverse faults and abundant 

seismicity. Since 1912, three M=6 shocks have 

occurred offshore in the channel, and larger reverseslip earthquakes are thought to be possible in the 

future (Lee and others, 1979). The uplift rates along 

the Santa Barbara Channel are an order of magnitude 

greater than those calculated from displacements of 

dated shoreline features farther south along the 

California coast between San Diego and Los Angeles, a 

region of comparable seismicity dominated by strikeslip faulting. Although there have been no large 

historical earthquakes accompanied by significant 

coastal uplift along the Santa Barbara Channel, it may 

be prudent to interpret much of the observed deformation as the cumulative effect of many episodes of 

uplift associated with shocks of M27.0. 

Elsewhere along convergent plate boundaries, 

such as in southern Alaska and along the Aleutian 

Islands, where vertical deformation is more rapid and 
thus more clearly expressed in the geologic record, the 

late Quaternary history of deformed shorelines can be 

especially useful in assessing uplift potential (for 

example, Plafker, 1969; Plafker and Rubin, 1978). 

The pattern of uplift and subsidence 

accompanying a postulated earthquake is predictable 

from elastic-dislocation theory, assuming the 

geometry of the causative fault and the distribution of 

slip on the fault (Savage and Hastie, 1966). The 

accuracy of this prediction is limited by the validity of 

assumptions about the fault and the distribution of 

slip. Experience with modeling the deformation in 

historical earthquakes (Savage and Hastie, 1966; 

Hastie and Savage, 1970; Jungels and Frazier, 1973) 

indicates that gross patterns of uplift and subsidence 

are satisfactorily predictable. Experience also 

suggests, however, that significant detailed features in 

the deformation associated with future earthquakes 

may not be predictable because of ignorance 

concerning the complexities in subsurface fault 

geometries and the distribution of slip on primary and 

subsidiary faults. The widespread occurrence of thick 

layers of soft sediment beneath the sea floor further 

complicates these predictions in the offshore environment. 

Seismic shaking 

Seismic shaking affects a large area around the 

earthquake source. Because of its pervasiveness, this 

hazard typically cannot be avoided or significantly 

reduced by shifting the site for a facility by a few 

kilometers. Although the location of offshore 

facilities seldom can be governed by considerations of 

seismic shaking, such considerations may preclude 

certain structural configurations in particular areas 

and may guide the selection of pipeline routes 

(Patwardhan, 1978). 

The primary strategy for mitigating shaking 

hazard is accommodation by designing and 

constructing facilities to withstand earthquake shaking 

with limited damage. Relative to other mitigation 

strategies, the practice of earthquake-resistant design 

and construction is quite advanced. Years of 

experience with structures on land are relevant to 

marine structures founded on the sea floor. Sophisticated modeling and analysis techniques have been 

developed and applied in the earthquake-resistant 

design of costly critical onshore structures, such as 

high-rise buildings, bridges, pipelines (Ariman and 

Muleski, 1981), and nuclear powerplants. With some 

modifications, these techniques are also applicable to 

the design of sea-floor-supported offshore platforms 

(Delflache and others, 1977; Nair, 1978; Watt and 

others, 1978; Bea and others, 1979) and to the 

associated piping, equipment, and appurtenances 

essential to safety and operations (Kost and Sharpe, 

1977). McGavin (1981) addressed the problem of 

designing essential building equipment to remain 

operational after a major earthquake; though directed 

toward buildings, much of his discussion is also 

relevant to the aseismic design of essential equipment 

on an offshore platform. 

For many facilities, the consequences of damage 

or failure can be limited if safety procedures are 

implemented promptly when a large earthquake 

strikes. Earthquake-alarm systems can be designed to 

initiate emergency procedures in the advent of strong 

shaking. For example, an alarm system has been 

installed on the Maui A Platform in the Tasman Sea 

off New Zealand (Tyler and Beck, 1983). This system 

indicates when horizontal motions of the platform 

reach half the seismic-design levels. To avoid false 

alarms related to the noisy vibrational environment on 

the platform, motion is monitored only in the 

frequency band of the six lowest vibrational modes of 

the platform, and ground motion is simultaneously 

monitored at an adjacent onshore site 37 km away, so 

that the seismic origin of an alarm signal can be 

confirmed. 

In regard to floating offshore structures, seismic 

shaking has received little attention, either as a potential hazard or as a design consideration (Hove and 

others, 1982). 

The seismic waves from an earthquake can be 

approximately grouped in two classes: body waves, 

which radiate in all directions from the earthquake 

source; and surface waves, which propagate along the 

boundary separating two mediums of contrasting 

seismic velocity, such as the sea floor or the Earth's 

surface. The energy in body waves propagates in all 

directions, whereas that in surface waves is trapped 

within a boundary layer. Thus, body waves attenuate 

more rapidly with distance than do surface waves, and 

at large distances surface waves may be the dominant 

type of seismic wave. 

Surface waves generally have much longer 

periods than do body waves and travel at lower 

velocities. The periods of surface waves are characteristically longer than 1 s, in the range containing the 

fundamental natural periods of common bottomsupported offshore platforms (Bea and others, 1979). 

In comparison, the dominant periods of body waves are 

characteristically less than 1 s, in the range within 

which the resonant periods of ships and floating 

platforms typically lie (Hove and others, 1982). A 

further difference between these two wave types is 

the dispersion of surface waves. Because seismic 

velocities vary with depth within the Earth, surface 

waves of different period, or wavelength, propagate at 

different speeds and thus cause an extended wave train 

to be observed at distances from the source. In 

contrast, body waves are nondispersive and appear as 

short relatively high-frequency pulses, apart from the 

complication of refractions and reflections and the 

extended source dimensions of large earthquakes. 

Thus, strong ground shaking initiates with highfrequency body waves, followed by an extended train 

of low-frequency surface waves. 

The two types of body waves are (1) 

compressional (P) waves, for which the motions of 

particles in the medium parallel the direction of 

propagation, as with sound waves in air, and (2) shear 

(S) waves, which travel more slowly than P waves and 

for which the motions are perpendicUlar to the 

direction of propagation, as with waves on a vibrating 

string. In nearly all geologic situations, f_- and .§_-wave 

velocities increase with depth beneath the surface. 

Evaluation and Mitigation of Earthquake Hazards 39 
This behavior causes body waves to refract upward and 

propagate nearly vertically as they approach the 

surface. Accordingly, at the ground surface, ~waves 

excite predominantly horizontal motion, and R. waves 

vertical motion. On land, structures are designed to 

withstand vertical loads greater than gravity, but 

horizontal design loads are typically much less than 

gravity. Shaking damage to onshore structures is 

largely attributable to nearly vertically propagating 

shear waves, which excite structures in their weaker, 

horizontal directions. Offshore structures founded on 

the sea floor are subject to both P and S motions; 

however, floating structures are isOlated from shear 

waves, which cannot propagate in a fluid. 

Nonetheless, floating structures are vulnerable to 

shaking damage; compressional waves alone can carry 

enough energy to cause serious damage, at least in the 

source region of a major earthquake (Hove and others, 

1982). The damage potential of seismic shaking is 

determined by several properties of seismic waves: 

amplitude, duration, frequency content, and pulse 

sequence. Damage potential increases with the amplitude of shaking; however, the relation is complex 

because of the nonlinear inelastic response of sedimentary deposits and structures to damaging levels of 

motion. Structures and, in some places, sedimentary 

deposits respond to shaking in a resonant manner, 

whereby relatively large deformations and stresses 

result if the shaking includes several cycles of motion 

with frequencies close to the resonant frequencies of 

the structure or deposit. Duration and pulse sequence 

are important factors because failure mechanisms in 

both structures and geologic materials depend on the 

number of stress-strain cycles greater than some 

amplitude as well as on the sequence in which cycles 

of different amplitudes occur. Thus, no single 

parameter can adequately portray the damage 

potential of seismic shaking. 

The principal source of quantitative information 

on ground shaking is the set of onshore instrumental 

records of strong ground motion collected from past 

earthquakes. Summaries of significant strong-motion 

records obtained in the United States since recording 

began in 1932 are contained in the annual publication 

"United States Earthquakes," published by the U.S. 

Department of Commerce and U.S. Geological 

Survey. In addition, records collected since 1976 in 

California under the California Strong Motion Instrumentation Program are cataloged in the Special 

Report Series of the California Division of Mines and 

Geology. The small number of recordings available 

from Canada were compiled by Weichert and Milne 

(1980) and Weichert and others (1982). Information 

from nearly 1,000 digitized strong-motion records 

from 16 countries, including Japan and Italy, was 

compiled by Crouse and others (1980). 

Three factors limit the reliability of predictions 

of seismic shaking at offshore sites. First, no instrumental data are available against which to test predictions. Although ocean-bottom strong-motion seismographs have been deployed in recent experiments 

(Reece and others, 1981; Steinmetz and others, 1981), 

to date no sea-floor recordings have been obtained 

close to the source of earthquakes large enough to be 

40 Earthquake Hazards in the Offshore Environment 

of engineering importance. In regard to sea-surface 

recordings of compressional waves from a submarine 

earthquake, we are not aware of any records obtained 

close to the source region where damage has been 

reported. Second, strong shaking at offshore sites can 

be predicted by using theory to extrapolate from 

empirical data obtained on land, although onshore 

studies do not encompass the range of sedimentologic 

and hydrologic conditions that are commonly 

encountered in the marine environment. Few onshore 

data concern the dynamic response of soft, saturated 

sediment to intense seismic shaking. A further 

complication in the marine environment is the common 

presence of gas trapped in the sediment. Third, the 

onshore data base is deficient in recordings of 

damaging levels of shaking obtained close to the 

source of large earthquakes. For example, only few 

records have been obtained within 40 km of an M=7 

earthquake, and none exists from within 75 km ofan 

M=8 shock. Thus, prediction of seismic shaking offshore is very much an extrapolation beyond evidence 

obtained on land. 

The seismic motion recorded at a particular site 

from a given earthquake is the result of three 

processes: radiation of seismic energy at the earthquake source, propagation of the radiated seismic 

waves through heterogeneous Earth structure, and 

modulation of the seismic waves near the site by local 

geologic structure and surficial materials. Each of 

these processes is complex, and their effect on the 

recorded motion is difficult to predict because either 

the physics of the process is poorly understood or the 

shallow structure of the Earth is not known in 

sufficient detail, or for both reasons. Thus, reliable 

estimates of seismic shaking, with the precision that 

engineers desire, are largely impossible at this time. 

Recent advances, however, both in understanding of 

the component processes and in modeling of these 

processes by sophisticated computational methods 

enable us to limit the problem and to provide 

estimates of seismic shaking in terms of ranges of 

expected values. 

A comprehensive review of standard practices 

and methods for estimating strong ground motion on 

land was provided by Hays (1980). Aki (1982) briefly 

summarized, from a seismologic perspective, the state 

of the art in predicting strong ground motion from a 

basic understanding of fault mechanics and seismicwave propagation. However, a comprehensive discussion of the application of such state-of-the-art 

techniques to engineering situations has yet to be 

written. Boore (1983) has reviewed progress in strongmotion seismology from 1979 through 1982, with 

particular emphasis on the prediction of strong ground 

motion; an extensive bibliography is included. 

In the following discussion, we focus on the 

problem of predicting strong ground motion and 

address two aspects of this problem: prediction of 

motion from postulated earthquakes at specific sites, 

and probabilistic estimation of shaking on an areal 

basis. The problem of predicting compressional-wave 

disturbances at the sea surface has received little 

attention (Hove and others, 1982) and is not discussed 

further here; it is amenable to theoretical analysis and 

deserves investigation. 
Site-specific evaluation 

Detailed, quantitative estimation of site-specific 

ground motion requires sophisticated computational 

techniques for modeling the generation, propagation, 

and local modulation of seismic waves. Because of the 

large number of variables and the complexity of 

physical processes entering into the problem, seldom 

can these methods provide a reliable, precise estimate 

of the amplitude, frequency, duration, and pulsesequence characteristics of strong ground shaking. 

Many engineering applications, therefore, rely on 

simplified quantitative representations of ground 

motion, such as average response spectra, representing 

ensemble averages of spectra over a range of magnitude, distance, and site situations, or on representative 

time histories of ground motion derived from recorded 

or simulated strong motions. For a particular 

earthquake and site situation, these engineering representations of seismic ground motion are generally 

normalized on the basis of some predicted gross 

measure of ground motion, such as peak acceleration, 

velocity, or displacement and duration (Newmark and 

Hall, 1969; Newmark and others, 1973; Watt and 

others, 1978; Bea, 1979; Bea and others, 1979). 

Numerous statistical correlations of peakground-motion parameters with magnitude, distance 

from the earthquake source, and site geology have 

been derived from instrumental strong-motion records 

(Idriss, 1978; Boore and Joyner, 1982). Such correlations are useful in estimating ground motions for 

design purposes, though with some limitations. Attenuation curves (fig. 31) of Joyner and Boore (1981) show 

mean and 84th-percentile values of peak horizontal 

acceleration and velocity, based on earthquakes in 

western North America, primarily California. The 

acceleration curves are well constrained by data at 

distances beyond 3 km for the magnitude range 5.0-6.6 

and beyond 30 km for the magnitude range 7.4-7.7; the 

velocity curves are reasonably constrained only for the 

magnitude range 5.8-6.6. 

The data from which these curves are derived 

exhibit considerable scatter as can be seen in the 

residual data for peak acceleration (fig. 31B). The 

standard deviation of the log-normally diStributed 

acceleration data is a factor of 1.8, about equal to the 

change in acceleration expected for a unit change in 

magnitude (fig. 31A). The scatter in the velocity data 

is comparable to that in the acceleration data, but 

velocity is a more sensitive function of magnitude and 

varies by a factor of about 3 for a unit change in 

magnitude (fig. 31C). Peak velocities recorded on soils 

more than a few meters thick are generally amplified 

about 50 percent above those recorded on rock (fig. 

31 C), whereas peak accelerations appear to be 

relatively insensitive to site geology. Such empirical 

attenuation relations are utilized in the probabilistic 

mapping of ground shaking, as discussed in the next 

subsection. 

Caution must be exercised in estimating design 

ground motions from empirical attenuation relations. 

Within the ranges of distance and magnitude for which 

instrumental data are currently sparse or absent, 

estimates from empirical relations are modeldependent extrapolations, and estimates based on 

different models are commonly discordant (for 

example, compare Joyner and Boore, 1981, and 

Campbell, 1981). Furthermore, such estimates are 

subject to change upon the acquisition of new instrumental data that would better constrain the model. 

The data base is most deficient close to the sources of 

earthquakes of M;::: 7.0, where damage is most likely. 

On the other hand, for distances and magnitudes for 

which data are more numerous, attenuation relations 

based on different model assumptions, analysis 

techniques, and data sets commonly predict similar 

motions (Boore and Joyner, 1982). A further complication is the unexplained variation in the data, which 

undoubtedly reflects processes not accounted for in a 

simplified empirical model, such as azimuthal dependence in the radiation of seismic waves or localized 

amplification or attentuation of seismic waves due to 

geologic structure at or near the recording site. 

In the absence of empirical data, we must look to 

theoretical methods to estimate peak ground 

motions. A simple, two-parameter model of the 

earthquake source (Brune, 1970) provides a basis for 

estimating both long- and short-period levels of 

motion. For long periods, Hanks (1976) demonstrated 

that this model satisfactorily predicts the relative 

amplitudes of long-period motion excited by the 1952 

Kern County (M=7.5) and 1971 San Fernando (M=6.6) 

earthquakes, and has applied the model to estimating 

the level of long-period ground motion in the Los 

Angeles basin arising from a repeat of the 1857 Fort 

Tejon earthquake (M=7 .9) along the southern section of 

the San Andreas fault, considered to be the maximum 

earthquake likely to affect the Los Angeles area. For 

short periods, Hanks (1979a) showed that the rootmean-square ground acceleration measured over the 

duration of faulting can be estimated from the source 

model, and McGuire and Hanks (1980) documented that 

measured values of peak and root-mean-square 

acceleration correlate well. These theoretical models, 

based on a simplified source model, hold much 

promise, but additional studies are needed to verify 

their accuracy and their applicability to a large 

number and wide variety of earthquakes. 

Peak-ground-motion parameters provide only 

limited information on the properties of strong 

shaking-namely, the amplitude of motions within 

limited frequency bands-and they convey no information about the duration of shaking or the sequence of 

seismic-wave amplitudes of different periods. Over 

the past decade, however, methods have been 

developed and demonstrated for calculating synthetic 

seismograms that closely approximate many of the 

important characteristics of actual strong-motion 

recordings. This advance became possible with 

improved understanding of fault mechanics and 

seismic-wave propagation, and with powerful new 

computational techniques. Currently, representative 

strong-motion seismograms, limited to periods longer 

than about 1 s, can be readily calculated. For 

example, long-period seismograms in the Los Angeles 

area have been estimated for a major earthquake on 

the southern section of the San Andreas fault from 

both a theoretical (Bouchon and Aki, 1980) and a 

semiempirical (Kanamori, 1979) approach. 

Seismogram-modeling techniques are an 

important tool for investigating strong shaking on the 

Evaluation and Mitigation of Earthquake Hazards 41 
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100 1 

DISTANCE, IN KILOMETERS 

10 

100 

10 

100 

M=5.0 

100 

7.5 

7 .0 

6.5 

6.0 

5.5 

Figure 31. Peak horizontal ground acceleration and velocity versus shortest distance to surface 

projection of fault rupture, as a function of moment magnitude. .!::., Attenuation curves for 

50th- (mean) and 84th-percentile peak accelerations (from Joyner and Boore, 1981, fig. 4). 

Shape of curve is assumed to be independent of magnitude. _b Residuals of logarithm of peak 

acceleration with respect to mean acceleration curves plotted in figure 31A (left), as a function 

of distance and magnitude (from Joyner and Boore, 1981, fig. 5). ~ Attenuation curves for 

50th- (mean) and 84th-percentile peak velocities as a function of local geology at the site (from 

Joyner and Boore, 1981, fig. 9). 

42 Earthquake Hazards in the Offshore Environment 
sea floor. Swanger and Boore (1978a) used the modal 

superposition of surface waves to model long-period 

ground motion in sedimentary basins at several tens of 

kilometers from shallow earthquakes. They showed 

that flat-lying deep sediment forms a waveguide that 

enhances the surface-wave component of the ground 

motion. Swanger and Boore (1978b) extended this 

method to predict long-period motion on the 

Continental Shelf. For Love-type surface waves from 

an M=7.5 earthquake propagating 50 and 100 km across 

a typical shelf consisting of 100 m of soft soils overlying 8 km of consolidated sediment, the longer period 

motion is amplified by a factor of 3 or 4 relative to 

that for a typical onshore path in hard rock. The 

pseudorelative-velocity-response spectra at 5-percent 

damping for surface-wave motions are shown in figure 

32. Swanger and Boore (1978b) also demonstrated that 

a strong velocity contrast at depth can cause 

resonance within a narrow period band that may be 

important in the nonlinear response of offshore 

structures. 

Using a different computational method, Spudich 

and Orcutt (1982) calculated synthetic sea-floor seismograms for a velocity model typical of the 

Continental Shelf, except for the important absence of 

the water layer. They assumed a shallow obliquethrust earthquake and computed vertical and radial 

seismograms over a broad distance range. The 

synthetic seismograms (fig. 33) show an extended train 

of low-frequency surface waves following the highfrequency body waves. For this case, the horizontal 

motions (radial component) are about 3 times larger 

than the vertical motions and continue about twice as 

long. Spudich and Orcutt separately considered the 

effect of a 1.5-km-thick layer of water on vertical 

sea-floor motions and found that the water overburden 

decreases high-frequency motions significantly and 

alters the dispersion characteristics substantially, so 

that the surface waves persist much longer. The 

effect of water overburden would be less significant 

for water depths typical of continental-shelf areas, 

that is, for depths less than 200 m. 

Although the complete source-propagation-site 

problem cannot be modeled satisfactorily at shorter 

PERIOD, IN SECONDS 

Figure 3Z. Pseudorelative-velocity-response spectra 

at 5-percent damping, computed for surface-wave 

(Love) motions for representative onshore and 

continental-shelf sites at distances of 50 km (A) and 

100 km @ from an M=7 .5 earthquake (from Swanger 

and Boore, 1978b, fig. 3). 

periods, the problem of local modulation of seismic 

shaking at the site is tractable. The dynamic response 

of surficial sediment can be computed by using 

vertically propagating body waves incident at the base 

of the sediment. A strong-motion record from a rock 

site is selected that is typical of the earthquake 

magnitude and distance under consideration. This 

selected record is then used as input motion at the 

sediment-rock interface, and the dynamic response of 

the sediment is computed to provide an estimate of 

the surface motion. The accuracy of this computation 

is limited by the extent to which the dynamic stressstrain relations of the sediment are known. The 

response of surficial geologic materials to 

nondamaging levels of motion, for which the stressstrain relation remains linear, can be predicted with 

considerable accuracy (Joyner and others, 1976; 

Johnson and Silva, 1981). 

Estimates have been made of the response of 

surficial sediment to damaging levels of input motion 

in the underlying bedrock, that is, of the high-strain, 

nonlinear response that might occur in the nearfield of 

a large earthquake. However, no empirical data are 

available for testing and calibrating such estimates. 

Joyner and Chen (1975) computed the nonlinear 

response of a 200-m-thick section of water-saturated, 

firm alluvium excited at its base by horizontal shaking 

with a peak acceleration of 0.7 .[and a peak velocity 

of 67 cm/s. Relative to a site with exposed rock, 

motion at the surface of the alluvium is amplified for 

periods longer than 1.5 s by a factor of as much as 2 

and is somewhat reduced for short periods. Moriwaki 

and Doyle (1978) analyzed the response of three postulated offshore sites characterized at the surface by 

rock, by 65 m of stiff clay, and by 65 m of soft clay. 

They found that the stiff clay amplified moderate 

input motion (peak acceleration, 0.3 .[) and generally 

attenuated intense motion (peak acceleration, 0.6 .[); 

whereas the soft clay greatly attenuated short-period 

components and amplified long-period components of 

both moderate and intense input motions. For 

moderate motion, the soft clay attenuated (by a factor 

of at least 2 or 3) spectral components with periods 

shorter than about 1 s and amplified components with 

periods longer than about 3 s; the attentuation of 

short-period components was greater for intense 

motion. For weaker input motion (peak acceleration, 

0.05 g), soft cohesive soils can amplify motion at 

periods less than 1 s (Martin and others, 1979). 

Thus, methods are available, some involving 

state-of-the-art mathematical modeling, to estimate 

ground-shaking characteristics for site-specific situations. The rigor with which these methods should be 

pursued to determine the shaking potential should 

depend on the earthquake exposure of the site, the 

type of structures or facilities being designed, and the 

possible consequences of damage or failure. 

Areal evaluation 

The types of ground-shaking evaluation described 

in this and the preceding subsection cannot be 

completely separated. A site-specific study may be 

extended to a small area, such as a petroleum 

reservoir or lease area, to assess the implications of 

Evaluation and Mitigation of Earthquake Hazards 43 
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Figure 33. Theoretical sea-floor seismograms as a function of distance for a representative velocity model of the 

Continental Shelf (modified from Spudich and Orcutt, 1982, figs. 2, 3) • .!!:., Vertical motion. _!!, Radial motion. 

Amplitudes indicate relative ground velocities at the various distances. Traveltime is reduced by time required 

for a 5-km/s wave to propagate to the site, or approximately by the .!:-wave traveltime. At increasing distance, 

seismograms are dominated by extended trains of late-arriving low-frequency surface waves. 

varying thicknesses of sediment or of varying locations 

of sites relative to known active faults. 

Site-specific evaluations are generally undertaken for critical facilities in earthquake-prone 

regions, and commonly for less critical structures as 

well, when maps of predicted strong ground motion are 

not available, are too small in scale, are not 

sufficiently conservative (that is, display values of 

ground motion that are too likely to be exceeded), or 

do not depict ground motion in the frequency band of 

interest. 

Ground-shaking maps are important for offshore 

development. They provide information on the spatial 

variations in shaking severity that can be factored into 

decisions on priori ties of areas to be leased and on 

strategies for safe and economic exploitation of offshore resources. They also provide ground-motion 

parameters for design of less critical facilities. We 

contend, however, that the information portrayed in 

ground-shaking maps derived by uniform application of 

present techniques to a large area is generally insufficiently precise and reliable to serve as a basis for the 

design of a critical facility in a region of moderate or 

high earthquake potential. 

44 Earthquake Hazards in the Offshore Environment 

Construction of probabilistic 

ground-shaking maps 

Most recent ground-shaking maps provide 

probabilistic estimates of potential shaking, which are 

based not only on the recorded history of damaging 

earthquakes but also on geologic evidence of active 

tectonism that suggests the likelihood of future earthquakes in areas devoid of historical seismicity. Such 

maps can serve as a reference source for the levels of 

ground motion to be accommodated in the earthquakeresistant design of common or noncritical structures 

(for example, Applied Technology Council, 1978). 

Probabilistic shaking maps have recently been 

published for Canada and its continental margins 

(Basham and others, 1982, 1984), the contiguous United 

States (Algermissen and others, 1982), coastal 

California and the adjacent Continental Shelf 

(Thenhaus and others, 1980), the Pacific Northwest and 

the adjacent shelf (Perkins and others, 1980), the East 

Coast and the adjacent shelf (Perkins and others, 

1979), Alaska and the adjacent shelf (Thenhaus and 

others, 1985), and individual lease areas on the Alaskan 
Continental Shelf (Woodward-Clyde Consultants, 

1978). 

Probabilistic ground-shaking maps have four 

basic elements in their construction: a set of seismicsource zones developed from a seismotectonic model 

of the study region; magnitude-recurrence relations 

for each of these zones; functional relations for the 

magnitude and distance dependence of the strongground-motion parameters being mapped; and an 

analytical technique for computation of the groundmotion parameters at selected points throughout the 

region at desired probabilities of exceedence (fig. 34). 

The principal difference between the source 

zones delineated for areal analysis of shaking potential 

and those developed for a site-specific study is one of 

scale and detail of knowledge. The level of geologic 

and geophysical investigation that can be pursued to 

define an earthquake source near a specific site commonly is not feasible for broad-scale analysis of a 

large area. There is a limit to the rigor of earthquakesource investigations, imposed by the requirement to 

depict strong ground shaking with comparable 

accuracy at all points on the map. Thus, the degree of 

confidence in the ground-motion estimate at an 

arbitrary point is less than that for a site-specific 

evaluation. Examples of seismic-source zones 

developed for areal evaluation of shaking potential are 

illustrated in figures 19 and 25; further illustrations 

were presented by Patwardhan and others (1981), 

• 

A 

z 

0 

f= 

0 

~ 

0 

z 

::::) 

0 

a: 

(.9 

c 

• 

• 

• • 

• • • t 

Computation 

grid 

~ DISTANCE 

B 

D 

z 

Cl 

_Q 

MAGNITUDE 

GROUND MOTION 

Figure 34. Schematic illustration of steps in method 

for determining probabilistic estimates of seismic 

ground motion (from Basham and others, 1981, fig. 4). 

.!::., Earthquake-source zones and computation grid. ~ 

Magnitude-recurrence relations, terminating at an 

upper-bound magnitude. .5; Ground-motion 

attenuation. .Q, Calculated probability distribution of 

ground-motion parameters at a gridpoint. 

Thenhaus (1983), and Basham and others (1982). 

Magnitude-recurrence relations establishing the 

frequency of shocks as a function of magnitude are 

required for each of the source zones. In some zones 

the historical and recent seismicity provides a reliable 

estimate of the rates of significant earthquakes (see 

fig. 23). In other zones it does not, and judgment is 

required to assign magnitude-recurrence relations 

based on available information (for example, Perkins 

and others, 1980; Basham and others, 1982). 

Magnitude-recurrence relations are terminated 

by some selected upper-bound magnitude. Though 

straightforward in zones that have undergone great(~ 

7 .5) historical earthquakes, the selection of upperbound magnitudes is generally difficult in less active 

seismic zones. If the cause of the earthquakes is 

poorly understood, as, for example, on the eastern and 

Arctic margins of the United States and Canada, then 

the dimensions of potentially active faults, the magnitudes of associated upper-bound earthquakes, and the 

frequency of shocks are not known. AI though great 

earthquakes occur principally in regions of major plate 

interactions, earthquakes approaching M=7 have also 

occurred on the passive Atlantic marginof the United 

States and Canada during historical time. The 

influence of the choice of upper-bound magnitudes on 

mapped values of shaking can be determined as part of 

a sensitivity analysis. For example, if an upper-bound 

magnitude has an estimated average recurrence rate 

of 0.001 event/yr, it will not contribute heavily to 

ground motions computed at an annual exceedence 

probability of 0.01. 

Attenuation relations, which relate the strongmotion parameters being mapped to magnitude and 

distance, are also required, even though available data 

to define these relations for the region in question may 

be scarce or absent. Extrapolations are sometimes 

possible from a region for which data are available to 

a region of similar geologic and tectonic environment 

that lacks data, for example, from the western 

conterminous United States to western Canada 

(Hasegawa and others, 1981). Even though the regions 

are dissimilar, the almost complete absence of 

relevant strong-motion recordings in eastern North 

America has promoted extrapolations of data from the 

Western United States based on empirical comparisons 

of the attenuation of modified Mercalli intensities (for 

example, Nuttli, 1979; Hasegawa and others, 1981). 

Two important limitations of attenuation 

relations are the absence of controlling instrumental 

data in the nearfield of large earthquakes and the 

substantial scatter in the limited data that do exist. 

Most compilations of ground-motion-amplitude data 

show the standard deviation for an individual observation to be a factor of about 2 about the average 

regression value (for example fig. 31). This scatter 

can be accounted for in estimating ground motion by 

including a stochastic term in the attenuation 

relation. Although the effect of this term on 

estimated ground-motion values with a small probability of being exceeded is to increase these values 

relative to the case in which there is no scatter in the 

data (Weichert and Milne, 1979), the amount of this 

increase will depend on the assumed standard deviation 

and on the slope of the magnitude-recurrence relation 

(see Algermissen and others, 1982). 

Evaluation and Mitigation of Earthquake Hazards 45 
Most recent probabilistic ground-shaking maps 

are based on the computational method developed by 

Cornell (1968). This method treats earthquakes as 

occurring randomly in space and time throughout 

planar source zones at rates specified by the 

magnitude-recurrence relations. For each gridpoint on 

a map, a distribution function for the probability of 

exceedence is computed by numerically integrating 

contributions to the ground-motion parameter from all 

relevant source zones (fig. 34). The results for all 

gridpoints can be displayed as contours of the groundmotion parameter at a fixed probability of 

exceedence, or of the probability of exceedence at a 

fixed level of ground motion. Computer programs 

have been developed to treat earthquakes as either 

point or linear sources (McGuire, 1976, 1978), and to 

allow for the nonrandom occurrence of earthquakes 

(Patwardhan, 1978). 

Maps depicting shaking at very low probabilities 

of exceedence should be used with caution. The choice 

of a very low level of probability yields extreme values 

of ground motion in the nearfield of large 

earthquakes. In the commentary accompanying the 

recent national ground-shaking maps of Canada, the 

user is cautioned to undertake additional investigations 

in areas of high earthquake potential (Heidebrecht and 

others, 1983). 

Western Canada 

Probabilistic ground-shaking maps of western 

British Columbia, southeastern Alaska, and the adjacent offshore region are shown in figure 35. The 

seismotectonic model (fig. 19A) and magnitude-recurrence relations (fig. 23) for part of this region were 

described above. These maps display, for sites underlain by firm soil, values of peak horizontal acceleration and velocity with probabilities of exceedence of 

10 percent in 50 years. In other words, there is a 90-

percent probability that the mapped values will not be 

exceeded over a 50-year interval. 

The active Queen Charlotte and Fairweather 

faults, which are the dominant earthquake sources in 

western Canada and southeastern Alaska, respectively, 

account for the belt of high-amplitude motion along 

the coast (fig. 35). Because of the stronger dependence of peak velocity on magnitude and the slower 

attenuation of peak velocity with distance, the influence of the major coastal faults is larger and more 

pervasive for velocity than for acceleration. 

These maps depict probabilistic ground shaking 

within two dominant frequency bands-acceleration in 

a band around 5 Hz, and velocity in a band around 1 

Hz-and thus provide information on relative levels of 

expected shaking, respectively, for ordinary rigid 

structures and for larger, more flexible structures with 

lower resonant frequencies, such as offshore platforms. 

Gulf of Alaska 

Two probabilistic peak-acceleration maps for the 

northeastern Gulf of Alaska region (fig. 36) enable the 

46 Earthquake Hazards in the Offshore Environment 

user to compare estimates of acceleration derived by 

different investigators using somewhat different 

approaches. 

Thenhaus and others (1985) published a series of 

maps depicting probabilistic esti~ates of acceleration 

on rock at varying return periods for Alaska and the 

adjacent offshore region. These maps are revisions of 

maps prepared earlier for Federal agencies responsible 

for leasing Outer Continental Shelf oil and gas 

resources and supervising exploration and production4 The Gulf of Alaska part of the 1 00-year-return-period 

map is shown in figure 36A. For this study, the entire 

State and adjacent shelf area were partitioned into 24 

seismic-source zones. Some zones are well controlled 

by knowledge of historical or Holocene faulting in 

southern and southeastern Alaska and of subduction of 

the Pacific plate beneath southern Alaska; others 

simply encompass areas of apparently homogeneous 

seismicity and (or) distinctive geologic structure. 

Upper-bound magnitudes range from 7.3 for lesser 

seismic zones to 8.5 for zones representing major 

interactions between the Pacific and North American 

plates. The smaller earthquakes (M <6.4) were modeled 

as point sources, and, where single faults comprise 

seismic-source zones, the larger earthquakes were 

modeled as line sources, with length appropriate to 

magnitude. Within each zone, earthquakes were assumed to occur randomly over time. The Schnabel and 

Seed (1973) peak-acceleration-attenuation relations for 

rock sites were employed, and no term was included to 

allow for stochastic scatter in acceleration. 

Woodward-Clyde Consultants (1978) analyzed an 

area of the Gulf of Alaska from Yakutat Bay to west 

of Kayak Island (fig. 36_m in their Offshore Alaska 

Seismic Exposure Study prepared for the 21 oil 

company members of the Alaska Subarctic Offshore 

Committee (see summary by Patwardhan and others, 

1981). They derived more detailed sources for the 

subduction zone and shallow fault structures, and 

represented each earthquake as a rupture on a planar 

surface. Earthquakes of M <7 .5 were assumed to occur 

randomly over time. In contrast, larger shocks within 

the subduction zone were modeled by a nonrandom 

probability of occurrence, consistent with the identification of a seismic gap within the study region (fig. 

24). Probabilistic estimates corresponding to a 100-

year return period were computed for an exposure 

time of 40 years, beginning at the time of the study. 

Two peak-acceleration- attentuation relations were 

3The concept of return period applies to a 

process in which events occur randomly over time. 

The return period is the average time interval between 

events. Here, the event of interest is the occurrence 

of shaking exceeding some threshold value. Return 

period R is related to the probability .e_ that the threshold value will be exceeded over an exposure time T 

through the equation 1-.e_=exp(-.!/R), where.!. and Rare 

measured in terms of the same unit. If T is much 

smalle.[, than B:, then .e_rv.!/R. - The actual return period is 95 years, corresponding to a 10-percent probability of exceedence over 

10 years. 
400 KILOMETERS 

Figure 35. Probabilistic maps of peak horizontal 

acceleration and velocity in western British Columbia 

and adjacent offshore region at a 10-percent 

probability of exceedence in 50 years, or a return 

period of about 500 years (from Weichert and 

Hyndman, 1982, fig. 5) • ..!!., Acceleration (contours in 

percent of gravity). _!!, Velocity (contours in 

centimeters per second). 

employed, one for earthquakes shallower than 20 km, 

and another for deeper events; both relations are for 

sites with less than about 45 m of stiff clay, sand, or 

gravel overlying rock. The relation for the deeper 

events predicts peak accelerations at a given distance 

from the rupture surface as much as 2 to 3 times 

larger than those predicted by the relation for shallow 

events. The scatter about the mean relations was 

assumed to be log-normally distributed. The Schnabel 

and Seed (1973) relations used by Thenhaus and others 

(1985) predict peak accelerations approximately intermediate between those given by the two WoodwardClyde relations. 

The results from these two studies are displayed 

in figure 36 as contours of peak acceleration, with a 

common annual probability of exceedence of about 

0.01. Analysis of the several factors in the two studies 

that may contribute to differences in the results is 

beyond the scope of this report. Given the differences 

in the various components of the analyses, the results 

are notably similar in predicting a range in peak 

acceleration from about 0.15 to 0.35 or 0.40 g_ on rock 

or stiff-soil sites along the Continental Shelf from 

Yakutat Bay to Kayak Island. 

A 

B 1 DO KILOMETERS 

Figure 36. Probabilistic maps of peak horizontal 

acceleration (in percent of gravity) for northeastern 

Gulf of Alaska region at a return period of about 100 

years. Dashed line represents 200-m isobath, 

approximate edge of the Continental Shelf. ..!!., Rock 

sites, 10-percent exceedence probability over 10 years 

(modified from Thenhaus and others, 1985, fig. 5). _!!, 

Sites on shallow stiff soil, 33-percent exceedence 

probability over 40 years (modified from Patwardhan 

and others, 1981, fig. 7). 

American Petroleum Institute guidelines 

The American Petroleum Institute (1982), in its 

guidelines for the planning, design, and construction of 

fixed offshore platforms, presented another example 

of ground-shaking maps for the Continental Shelf of 

the United States. The shelf is divided into six zones 

(0-5) in terms of earthquake potential and related 

ground motion (fig. 37). Each zone is assigned an 

effective horizontal ground acceleration, ranging from 

0.0 .[in zone 0 to 0.4 .[in zone 5, to be used for design 

of major steel-frame structures in the absence of sitespecific data. The effective accelerations are used to 

Evaluation and Mitigation of Earthquake Hazards 4 7 
CONTERMINOUS 

UNITED STATES 

Figure 37. Earthquake-risk zones used in American Petroleum Institute's guidelines for the design of 

fixed offshore platforms in United States coastal waters (from American Petroleum Institute, 1982, 

fig. 2.3.6-1). Effective horizontal accelerations for scaling of design response spectra are 0, 0.05, 

0.10, 0.20, 0.25, and 0.40 _g for zones 0 through 5, respectively. 

scale normalized response spectra, which are referenced to three types of foundation conditions: rock, 

shallow strong alluvium, and deep strong alluvium. 

The guidelines recommend that platforms located on 

significant accumulations of soft clay, loose sand, and 

silt deserve special analytical studies to determine the 

appropriate input ground motion. 

The guidelines state that for areas of low seismicity (zones O, 1, and 2), environmental factors other 

48 Earthquake Hazards in the Offshore Environment 

than earthquakes normally control platform design. 

They further state that for zone 0 no earthquake analysis is required but that for zones 1 and 2 sufficient 

calculations are required to confirm that earthquake 

loading does not exceed half the strength requirements 

imposed by other environmental conditions. For more 

seismically active areas (zones 3, 4, and 5), the guidelines recommend that the design ground motions should 

be determined by a site-specific study, but allow the 
design motions to be taken from figure 3 7 in the absence of detailed seismic data and site-specific studies. 

In reference to these guidelines, we recommend 

that for zones 3, 4, and 5 the corresponding effective 

horizontal ground accelerations of 0.20, 0.25, and 0.40 

g_ be used only for preliminary design considerations 

and that final strength requirements for earthquake 

loading of offshore platforms be based on site-specific 

studies utilizing state-of-the-art techniques to model 

expectable ground motion. Such modeling can provide 

better information on the relative levels of input 

ground motion within the period range of platform resonances. 

Sea-fioor failures 

Avoidance and accommodation are both important strategies for mitigating the potential hazards 

associated with sea-floor failures (Garrison and Bea, 

1977). The strategy of avoidance, which involves the 

selective siting of a structure to avoid potentially 

hazardous areas, recognizes that failure of surficial 

geologic deposits is geographically limited. Although 

strong ground shaking, which is the triggering mechanism for earthquake-related failures, is a pervasive 

earthquake effect, the extent of sea-floor failure is 

limited by the distribution of materials that are susceptible to failure. The strategy of accommodation is 

founded on the assumption that a structure can be designed to withstand threats to its integrity and performance arising from sea-floor failures. 

Avoidance is generally the preferred strategy 

where seismic shaking might trigger large permanent 

displacements within geologic foundation materials, 

because such displacements can impose catastrophic 

loads on the structure. For example, the hurricaneinduced soil movement that toppled South Pass 70 

Platform B in the Gulf of Mexico (Sterling and Strohbeck, 1973) displaced the base of the platform about 

25 m in a downslope direction and caused buckling failure in the piles on which the platform was supported. 

SOUTH 

Comparison of shear strength-depth profiles for soil 

borings obtained before and after the hurricane suggests that significant movements occurred to depths of 

at least 20 m. Measurable changes in bottom tppography extended over an area greater than 15 km , and 

the maximum change in water depth exceeded 10 m. 

Slides of comparable or greater size, which at least in 

part are attributable to seismic shaking and which in 

some places are probably the result of several episodes 

of shaking, are visible in bottom and subbottom reflection profiles obtained in seismically active zones off 

California (Edwards and others, 1980; Lee and others, 

1981), British Columbia (Yorath, 1980), and Alaska 

(Hampton and Bouma, 1977; Molnia and others, 1977; 

Carlson, 1978). For example, seismic profiles in the 

Kayak Trough in the northern Gulf of Alaska (fig. 38) 

reveal a massive submarine slide above the inferred 

rupture zone of the great 1964 earthquake. 

The strategy of accommodation, however, may 

be practicable for sea-floor failures in which a temporary loss of strength in geologic materials occurs, but 

little permanent displacement. Without significant 

displacement of the geologic foundation materials, a 

structure is not subjected to extreme external loads, 

as occur in sliding. Thus, the detrimental effects of 

liquefaction in shallow sediment can be countered by 

anchoring the foundation of a structure below the 

depth to which liquefaction occurs, such as for a pilesupported structure, or by designing a buried structure 

to be neutrally buoyant with respect to the liquefied 

sediment, such as for a pipeline. Implementation of 

either mitigation strategy requires the ability to identify areas subject to seismically induced failures. The 

strategy of accommodation further demands the capability to predict the probable characteristics of potential failures, particularly the depth and areal extent of 

failure and the amount and pattern of resulting displacements. 

This section focuses on failures in sedimentary 

deposits. Although submarine rock slopes may fail in 

response to strong seismic shaking, failure of sedimentary deposits is the dominant problem confronting 

coastal and offshore structures because of the wideNORTH 

10 KILOMETERS 

Figure 38. High-resolution seismic profile (vertical section) approximately along axis of an 18-km-long submarine 

slide in the Kayak Trough, northern Gulf of Alaska (from Carlson and Molnia, 1977, fig. 12). Multiple reflection 

arises from energy reflected twice between sea surface and sea floor. Thickness of disturbed zone decreases 

from about 120 m south of scarp at head of slide to about 20 m near toe. 

Evaluation and Mitigation of Earthquake Hazards 49 
spread presence of sediment and its greater susceptibility to failure. 

Few incidents of seismically induced sea-floor 

failure have been documented, and none has been 

thoroughly investigated to determine the specific 

geotechnical and seismic conditions under which it 

occurred. Accordingly, evaluation of the physical 

mechanisms and the seismic, geologic, hydrologic, and 

geotechnical conditions that contribute to sea-floor 

failures relies heavily on experience with subaerial 

failures and laboratory investigations of failure 

mechanisms. A descriptive summary of subaerial 

failures associated with historical earthquakes in 

northern California was compiled by Youd and Hoose 

(1978). Detailed field investigations and engineering 

analyses of individual ground failures since 1964 have 

significantly increased our understanding of the 

mechanism of and conditions for failure. Particularly 

relevant to the problem of sea-floor failures are 

studies of failures in sedimentary deposits on gentle 

slopes and level ground (for example, Seed and Idriss, 

1967; Seed and Wilson, 1967; Youd, 1973). The importance of pore-fluid-pressure changes and the cyclicity 

of seismic loading in the failure of sedimentary 

deposits have been demonstrated in laboratory studies 

(Seed and Chan, 1966; Lee and Seed, 1967; Peacock 

and Seed, 1968). 

The occurrence of earthquake-induced failures in 

sea-floor sediment depends on the character of seismic 

shaking and the susceptibility of the sediment to 

failure. Two important characteristics of shaking are 

its amplitude and duration. The importance of duration is documented both in case studies of failures in 

sedimentary deposits on land (Lemke, 1967; Seed and 

Wilson, 1967) and in laboratory investigations of 

failure processes (Lee and Seed, 1967). For example, 

the extensive lateral sliding at Turnagain Heights in 

Anchorage during the great 1964 Alaska earthquake 

began some 1-1/2 to 2 minutes after the shaking 

began, and the area of sliding expanded as shaking 

continued (Seed and Wilson, 1967). Significant sliding 

occurred only after many cycles of shaking had 

weakened the underlying soils. 

Both amplitude and duration of shaking increase 

with earthquake magnitude. Thus, sea-floor failures 

must be expected to be more numerous and to occur at 

greater distances from the source of seismic shaking in 

larger earthquakes. Such behavior has been 

documented for liquefaction failures on land (Youd and 

Perkins, 1978). For earthquakes of M<5, significant 

liquefaction effects on land have notbeen observed; 

for shocks of M=5.0-5.9, potentially damaging liquefaction failures have been observed to distances of 

about 10 km; and for shocks in the range M=6.0-6.9, 

the maximum distance is about 50 km. In-the 1964 

Alaska M=9.2 earthquake, liquefaction-induced ground 

displacements capable of causing significant damagedifferential horizontal or vertical ground displacements of more than 10 em-occurred out to distances 

of 130 km (Youd and Perkins, 1978). Because liquefiable sediment is more likely to be present on the sea 

floor, this onshore experience should be regarded as an 

uncertain minimum guide to the liquefaction potential 

in the offshore environment. 

The susceptibility of marine sediment to failure 

during earthquakes is governed by the dynamic 

50 Earthquake Hazards in the Offshore Environment 

strength of the sediment and by the slope of the sea 

floor. In some places, the inertial forces caused by 

seismic shaking may alone be sufficient to overcome 

the frictional resistance to downslope movement. This 

failure mechanism is probably more important in the 

offshore environment than on land because of the low 

shear strength of much marine sediment in its natural 

condition of saturation. In other places, seismic 

shaking may sharply reduce the shear strength of 

sedimentary formations, and lead to loss of bearing 

capacity in geologic foundation materials or to lateral 

or downslope mass movements. An important mechanism for seismic weakening of saturated sediment is the 

increase of pore-fluid pressures accompanying cyclic 

deformation of the soil during the passage of seismic 

waves (fig. 39). Such an increase of pore pressure 

results from local densification of the sediment as 

shaking causes particles to pack more tightly. This 

increase in pressure reduces the shear strength of the 

sediment. In extreme cases, the strength can drop to 

zero, so that the sediment will deform as a liquid, that 

is, liquefy. 

ffi 

..... 

z~ -;::: 

...;a:; 

~u 

cnw 

en a: L.Uct 

a:::::> 

0....0 

a: en 

~ffi <a... 

3!=cn 

~:E o< 

0.... a: 

z· 

t!J 

g 

:;;;:: 

- ..... <z 

a:w 

1-u 

en a: 

~~ LU Z ::t:-

Cl) 

Cl) 

:Ea: 

<w 

a: .... t!J LU 

g~ 

- ..... ::..::z 

zw 

-u 

a:::::::> 

~~ 1-0 

CI)CI) 

a: a: CS:LU 

LU 0.... ::t: 

Cl) 

A 

20 r 

1-

10 1-

,.... I 0 ..... "ij 

10-

24 cycles----.! - 20-

B 

0.4 

0.2 

0 

0.2 

0.4 

c 

Figure 39. Records from undrained cyclic-shear test 

on loose sand (from Seed and Idriss, 1982, fig. 42). 

Gradual increase in pore-water pressure (A) and 

ensuing abrupt decrease in shear strength-seen as an 

increase in shear strain ~-is caused by repetitive 

cyclic shearing of sample (C). Significant strain 

occurred after 24 cycles of shear, only when increase 

in pore pressure practically balanced initial confining 

pressure acting on the sample. 
The seismic stability of marine sediment is 

determined by the environment in which it is deposited 

and by its postdepositional history. Submarine failures 

are most common in areas of rapid deposition, where 

sedimentation outpacing consolidation results in underconsolidated deposits with naturally high pore 

pressures. Thus, submarine slumps and slides are 

particularly likely in the vicinity of active deltas (see 

Coleman and Garrison, 1977, for a discussion of the 

geologic aspects of marine slope stability in a deltaic 

regimen). Grain size and degree of sorting are 

important factors influencing seismic stability. 

Loosely packed well-sorted fine sand and coarse silt 

are particularly prone to liquefaction. The strength of 

a particular sediment increases with both consolidation 

and cementation, and because both processes are time 

dependent, stability generally increases with the age 

of a sediment. Thus, failures are more common in 

Holocene sediment than in Pleistocene deposits, and 

more common in Pleistocene units than in earlier 

Cenozoic formations. 

Because of the low strength of much marine 

sediment, submarine slope failures generally are larger 

and occur on flatter slopes than do subaerial failures 

(Hampton and others, 1978). Dimensions, slope, and 

relative importance of various instability factors 

(table 1) are compiled for several submarine slope 

failures attributable at least in part to strong seismic 

shaking. It is evident that failures on slopes of 1° and 

less are not uncommon, may involve large areas, and 

can extend to sufficient depths to cause severe 

foundation problems. 

Areal evaluation 

Methods have been developed for evaluating the 

potential of earthquake-induced failures on both areal 

and site-specific bases. Areal evaluations typically 

express the relative likelihood of failure throughout a 

region, possibly with dimensions of tens or hundreds of 

kilometers. Though qualitative and broad in scope, 

areal assessments are particularly valuable early in the 

development of an offshore area, for example, in 

identifying areas where sea-floor failures are particularly likely and in establishing minimum technical 

stipulations for the siting and design of offshore facilities. Areal evaluations, however, are insufficiently 

detailed and precise to determine the final siting and 

design of an important offshore structure. For that 

purpose, a site-specific evaluation is needed that will 

provide more specific, quantitative information 

regarding the probability of sea-floor failure during 

the lifetime of the structure and the probability and 

extent of resulting sea-floor deformation. 

Areal assessments of slope stability have been 

prepared for several earthquake-prone regions, both 

onshore (for example, Nilsen and others, 1979) and 

offshore (for example, Self and Mahmood, 1977). The 

variation in relative slope stability for part of the 

Kodiak shelf off Alaska is shown in fig. 40, using a 

fivefold classification of relative stability (table 2; 

Self and Mahmood, 1977). Classification of stability is 

based on three considerations: soil type, steepness of 

slope, and evidence of past slope failure. Soil types 

are differentiated both by stratigraphy-that is, 

geologic age and sediment type-and by engineering 

properties that could lead to varying foundation 

behavior. Soil types were mapped in a program of 

geophysical profiling, sea-floor sampling, and standard 

soil testing. The map provides a qualitative assessment of the variation in slope stability and identifies 

areas where slope-failure problems are likely to be 

either most acute or least severe. No information is 

provided, however, on the probability that slope failure 

will occur within a given exposure time. To date, 

areal evaluations of slope failure in seismic zones have 

not focused specifically on seismically induced slope 

failures but have considered all slope failures 

regardless of their origin. Nonetheless, recent 

research into the occurrence and mechanisms of 

earthquake-triggered slope failures is leading to areal 

assessments that specifically address seismically 

induced failures (for example, Wieczorek and others, 

1985). 

A method for probabilistic mapping of 

liquefaction potential on land (Youd and Perkins, 1978; 

Youd and others, 1978a) has been developed that 

incorporates information on the likelihood that a 

triggering level of ground shaking will occur within 

some future time interval. A liquefaction-potential 

map is prepared by combining two constituent maps: a 

liquefaction-opportunity map that depicts how often a 

level of ground shaking sufficient to cause liquefaction 

is likely to occur, and a liquefaction-susceptibility map 

that shows the distribution of sedimentary units and 

the relative ease with which the various units can 

liquefy during strong shaking. The resulting liquefaction-potential map delineates areas where 

liquefaction is most likely to occur over a given time 

interval, and is a useful guide to where additional sitespecific investigations may be needed before specific 

siting and engineering decisions. 

Site-specific evaluation 

The degree of hazard that sea-floor failures can 

pose to offshore and coastal facilities justifies detailed 

multidisciplinary investigations to evaluate the 

stability of the sea floor at and surrounding the site 

(Garrison and Bea, 1977). An early step in site investigations should be a search for evidence of past failures 

that might indicate the possibility of future failures. 

The seismic-reflection techniques used in the identification of surface faults (see subsection above entitled 

"Geologic Record") are also applicable to the problem 

of detecting and mapping failures involving substantial 

deformation (figs. 13, 38). Other investigations needed 

to assess the seismic stability of the sea floor include 

detailed bathymetric surveys to determine slopes, 

geologic mapping and sampling to identify what 

materials are present and to recognize active geologic 

processes, seismologic studies to evaluate exposure to 

seismic shaking (see subsection above entitled "Seismic 

Shaking"), and geotechnical studies to determine the 

engineering properties of geologic materials. 

Several methods are available for assessing the 

seismic liquefaction potential of saturated sediment 

(see reviews by Seed, 1979b, and Seed and ldriss, 

1982). The existing methods may be categorized as 

Evaluation and Mitigation of Earthquake Hazards 51 
U1 

t..:> 

r 

f 

~ 

I Table 1. Submarine slope failures attributed at least in part to seismic shaking 

[Length of disturbed area is measured in the downslope direction. Individual failure dimensions: L, length; W, width. Importance of instability factors: 

shaking; W, wave loading; 0, oversteepening; S, rapid sedimentation~ G, free gas. Importance: H, high; M, moderate; L, low; N, negligible] 

E, earthquake 

$' 

i 

0 

..... 

~ 

~ ~ 

~ 

5· 

::s 

a 

l 

Region Location 

--·-----·-------------- Southern California Off San Mateo 

Borderland. Point. 

Northern California Off Humboldt 

Borcterl and. Bay. 

Do------------------ Off the Klamath 

River. 

Gulf of Alaska-------- Off Icy Bay------

Do------------------ Kayak Trough-----

Do------------------ Copper River 

prodelta. 

Do------------------ Off Portlock 

Bank. 

Do------------------ Off Middle 

Albatross 

Bank. 

Physiographic Water Slope 

setting depth ( 0) (m) 

------·------

Lower continental 600- 3-4 

slope. 750 

Lower slope between 460- 1 

continental shelf 580 

and marginal 

plateau. 

Continental shelf----- 60-65 .25 

do------------------ 48-200 .5 

do------------------ 42-227 1 

do------------------ 40-125 .5 

Upper continental 760- 1.2 

slope. 1,200 

do------------------ 690- 3.4 

1,300 

Disturbed area Individual Sedirnen- Importance of ----------·-·----- failure tat ion instability 

Length Width Thick- dimensions rate factors Reference 

{km) {km) ness {km) (m/103 yr) E W 0 S G ( km) 

----·--------------

4.5 3 50 --- 0.05-0.1 H N - N N Edwards and others 

3.5 3 50 {1980). 

6.5 15 30-85? .35-0.60{W) --- H N - - - Lee and others 

( 1981). 

.6-1.4 20 5 --- --- H - N - - Field and others 

( 1982) 

10-20 90 15-40 .5-1.0( L), 68-80 M M N H N Carlson (1978), 

.2{W) Hampton and 

others ( 1978). 

18 15 20-120 --- 7.5-15 H M L H M Molnia and others 

( 1977), Hampton 

and others (1978) • 

8 100 20-40 • 3-1.0( L) 10-15 M M N H M Hampton and others 

( 1978). 

6.5 --- 200 --- --- H N M N N Do. 

5.3 12 300 --- --- H N H N N Do. 
--······ 58 ···· -- ······---- --··-· --- -·-- ---------

Figure 40. Relative-slope-stability map of part of the 

Continental Shelf south of Kodiak Island, Alaska, at 

approximately lat 56 ° N., long 154 ° W. (modified from 

Self and Mahmood, 1977, fig. 9). Heavy lines define 

areas of varying slope stability; numbers refer to 

classification scheme in table 2.. Light lines are 

bathymetric contours (in feet). 

either empirical or analytic. Empirical methods are 

based on field observations of liquefaction during 

previous earthquakes. The potential is evaluated by 

comparing the characteristics of the sediment and the 

expected earthquake shaking at the study site with 

those at sites where the presence or absence of liquefaction during earlier shocks has already been 

documented. Standard penetration resistance is the 

sediment characteristic most commonly used as the 

index of liquefaction susceptibility. Analytic methods 

are based on the behavior of saturated sediment during 

cyclic shearing as determined in laboratory investigations. Analytic methods featuring nonlinear-effectivestress analyses of the dynamic response of saturated 

sediment permit direct evaluation of cyclic mobility or 

liquefaction potential (Finn and others, 1977; Liou and 

others, 1977; Martin and Seed, 1979). Though new and 

relatively unproven, such methods promise to be highly 

useful in offshore applications. The application of 

both empirical and analytic methods was illustrated by 

Nataraja and others (1978) in a case study for a 

proposed waste-water outfall pipe extending about 800 

m offshore from Puerto Rico. 

The reliability and accuracy of empirical 

methods are limited by the sparseness of available 

field data on liquefaction and by the difficulties in 

reliably measuring sediment properties that are sensitive indicators of liquefaction susceptibility. Thus, 

empirical methods are most useful for preliminary 

evaluations of liquefaction potential. Where empirical 

methods do not indicate a large factor of safety, the 

evaluation should be supplemented by analytic methods 

(Seed, 1979b). Both empirical and analytic methods 

have been incorporated into probabilistic schemes for 

the site evaluation of liquefaction potential (for 

example, Yegian and Whitman, 1978; Haldar and Tang, 

1979; Davis and Berrill, 1982). 

Available methods for evaluating slope stability 

fall into two categories based on the failure criteria 

used in the analyses (Morgenstern and Sangrey, 1978). 

The first category includes limit-equilibrium methods, 

in which the stresses necessary to cause displacement 

along some postulated failure surface are compared 

with the available shear strength of the materials 

along this surface. Failure is considered to occur when 

stresses tending to cause displacement exceed the 

mobilized shear resistance of the material. The 

second category includes deformation methods, in 

which failure criteria are based on the deformation of 

a slope calculated from stress-strain analysis. Because 

it is the amount of deformation rather than the 

internal stresses in the material that poses a problem 

to the engineer, deformation methods of analysis are 

potentially more useful. Deformation methods, 

however, require knowledge of the stress-strain 

characteristics of the material, which is not needed in 

limit-equilibrium analyses, and also more sophisticated 

computations. 

In limit-equilibrium methods, the inertial forces 

arising from seismic shaking are modeled as static 

forces that are resisted by the undrained shear 

strength mobilized along the potential failure surface 

at the time of failure (Morgenstern, 1967). Both 

horizontal and vertical components of seismic shaking 

can be accommodated (Hampton and others, 1978). 

Psuedostatic analyses of seismic slope stability by 

limit-equilibrium methods have been used to estimate 

the sensitivity of submarine slopes to seismic shaking 

(Morgenstern, 1967) and to evaluate the possible effect 

of shaking on slope stability on the continental slope 

off Israel (Almagor and Wiseman, 1977) and on the 

Continental Shelf and upper continental slope in the 

Gulf of Alaska (Hampton and others, 1978). There are 

difficulties with such pseudostatic methods, in regard 

both to the gross simplification of a complex dynamic 

process as an equivalent static process and to the 

selection of appropriate static-force coefficients to 

represent cyclic seismic forces (Seed, 1967). 

In deformation methods of slope-stability 

analysis, the cyclicity of seismic shaking is explicitly 

accounted for in calculating the cumulative deformation of a slope. Newmark (1965) proposed a procedure 

for evaluating potential slope deformations for rigidplastic stress-strain behavior, on the basis of a simple 

model of a block resting on an inclined plane that is 

subjected to seismic accelerations. In this model, 

when the seismic acceleration exceeds some critical 

threshold, the inertial forces acting on the block 

overcome the yield resistance, and the block slips. 

The history of downslope movement is calculated by 

twice integrating over time the effective acceleration 

Evaluation and Mitigation of Earthquake Hazards 53 
Table l. Classification of relative slope stability, Kodiak shelf area, Alaska 

[From Self and Mahmood (1977, table 1). Rock types: Ki '.indurated met~morphic Cretaceous rock; Td, 

very dense and hard Tertiary sediment; Ti, indurated Tert1ary ro~k. So1l type~ (all of Quaternary .. age): Qc, soft to stiff clays; Qg, glacial till; Qls, loose to f1rm sand and s1lt; Qsl, sl1de depos1ts, 

Qsw, migrating sand waves.] 

------·------------------------·-------------------

Map unit 

(fig. 37) 

Re l.at i ve 

slope 

stability 

Rock and soil 

types Steepness of slope Condition of slope 

Highest---- Ti, Ki Nearly flat to moderate----- No evidence of mass movement or large-scale erosion. 

High------- Td Nearly flat to moderately 

gentle. 

Do. 

Moderate--- Qy, Qc, Qls, 

Qsv1. 

do------------------------ No evidence of mass movement; some scour apparent. 

4A Low-------- Qg, Qc, Qls, 

Qsw, Td, Ti. 

Moderate to very steep------ Evidence of creep or presence of 

tension cracks. 

48 Low-------- Qc, Qls Nearly flat to steep-------- Lies at base of potentially un- stable slope; may be covered 

by debris from future slides. 

5A Lowest----- {)sl, Qc, Qls, 

Td, Ti. 

Moderate to precipitous----- Strong evidence of incipient or recent rotational, translatatory, or flow slides; probably 

unstable. 

58 Lowest----- Qc, Qls Nearly flat to steep-------- Lies at base of unstable (5A) 

slopes. 

causing the block to slip. Using actual recordings of 

ground motion, Wilson and Keefer (1983) evaluated the 

procedure for a subaerial slope failure triggered by a 

moderate-size earthquake. The Newmark method has 

been extended in subsequent work to account explicitly 

for the buildup of pore pressures during earthquake 

shaking (Sarma, 1975) and to accommodate elastoplastic stress-strain behavior (Makdisi and Seed, 

1978). These deformation methods assume that the 

shear strength in the material along the failure surface 

is constant once failure begins. H.B. Seed and 

coworkers, employing finite-element-modeling 

techniques and dynamic testing of sediment and 

engineered soils in the labor a tory, have developed 

more sophisticated methods of deformation analysis 

that allow for changes in the strengths of materials 

that occur during the course of an earthquake (Seed 

and others, 1975; Seed, 1979a). To date, these 

methods have been applied primarily to embankment 

dams, although they could also be extended to the 

evaluation of submarine slope stability. A limitation 

to the application of such sophisticated finite-element 

analyses is the difficulty in determining the dynamic 

strain-dependent strength properties of sediment. 

Although improvements in methods of soil testing are 

helping to reduce this obstacle, a serious unsolved 

problem is the recovery of samples of marine sediment 

that are sufficiently undisturbed to be suitable for 

labor a tory testing to deduce their onsi te engineering 

properties and dynamic behavior. 

Turbidity currents 

Because turbidity currents derive from 

subaqueous slope failures, evaluation of slope stability 

54 Earthquake Hazards in the Offshore Environment 

is an important aspect of assessing the likelihood and 

possible effects of turbidity currents. Additional 

factors to be investigated are the conditions under 

which a submarine slump or slide is transformed into 

turbulent flow, the offshore location at which these 

conditions exist, and the characteristics of turbidity 

currents that would pose hazards to offshore 

facilities. The strategy of hazard mitigation would 

then be one of avoidance or accommodation. If 

avoidance is impossible for facilities downslope from a 

potential slump area, design features will be required 

to mitigate the potential hazards from sediment scour 

and hydrodynamic-current loads, which may result 

from turbidity currents. 

The first question to be addressed in any hazards 

assessment is whether the phenomenon can occur. It 

may be most useful to look first for evidence of past 

occurrence; circumstantial evidence can be sought 

both upslope and downslope from the site or project 

area. Upslope, bathymetry and shallow geologic 

structure can be examined for evidence of slope 

failures. Downslope, the fabric, sorting and composition of sediment can be examined where a turbidity 

current would be expected to dissipate and deposit its 

load. If the evidence for past currents is inconclusive, 

we must rely simply on evaluating the geologic and 

bathymetric conditions upslope from the site or 

project area. The methods discussed in the preceding 

subsection can be used to identify potentially unstable 

sedimentary slopes. In addition, consideration must be 

given to whether the sea-floor grade downslope from 

an area of potential instability is sufficiently steep for 

the slump or slide to attain the critical velocity for 

transformation into a turbidity current. 

Although the conditions under which a submarine 

slump or slide is transformed into a turbidity current 
are not well understood, steep slopes are clearly an 

important factor in the generation of turbidity 

currents. Thus, the potential for turbidity currents is 

greater on the continental slope, where petroleum 

exploration has not yet been vigorously pursued, than 

on the shelf. In particular, turbidity currents are 

recognized as an important sediment-transport process 

in submarine canyons on the continental slope; 

however, this does not imply that sufficient conditions 

of unstable sediment and steep slopes are absent on 

the Con tin en tal Shelf. 

Tsunamis 

Although tsunami wave heights are not great 

enough to be of concern in the design of offshore 

facilities in water depths of tens of meters, they can 

pose serious hazards to ancillary coastal facilities, 

such as loading terminals, storage tanks, and pipelines. All three mitigation strategies--selective siting, 

appropriate design, and early warning and 

preparedness-are to some degree effective in 

minimizing the losses from tsunamis, depending on 

whether the tsunami is of local or distant origin. 

Evaluation of the potential occurrence and 

effects of locally generated tsunamis must follow an 

assessment of coastal and nearshore submarine slides 

(see subsection above entitled "Sea-Floor Failures"). 

Sites that are hazardous from the standpoint of coastal 

slumping are also subject to inundation from local 

slide-induced waves, as evidenced in Seward during the 

1964 Alaska earthquake (fig. 16). The 1964 shock, 

however, also demonstrated that wave runup on shores 

adjacent to the area of coastal sliding is commonly 

smaller than on opposite-facing shores of bays and 

fiords (von Huene and Cox, 1972). To reliably evaluate 

potential local tsunami waves, knowledge is required 

of the history of local tsunamis and their generation 

mechanisms, the earthquake potential of the region, 

the seismic stability of coastal geologic units, and the 

hydrodynamics of the local coast, basin, or estuary. 

Typically, not all possible sources of local tsunamis 

can be identified, and so a design strategy to accommodate wave heights and runup greater than the 

largest known occurrences is important in coastal 

regions with high earthquake potential. 

For tsunamis from distant sources, the hazardmitigation strategy may involve not only prediction of 

tsunami effects as the basis for selective siting and 

appropriate design, but also early warning. Any site on 

a coast exposed to the open sea, or within a bay or 

inlet that is so exposed at its mouth, has some degree 

of tsunami exposure. The frequency of tsunami occurrence, however, varies significantly among the coastal 

regions of the world. Because approximately 80 

percent of global earthquakes border the Pacific 

Ocean, the Pacific has a much higher frequency of 

tsunamis than other oceans; from 1900 to 1973, 200 

tsunamis were observed or recorded ·in the Pacific 

(Spaeth, 1975). Though less frequent, tsunamis have 

also occurred in the Atlantic and Caribbean Oceans 

and in the Mediterranean and Black Seas. Compilations of historical occurrences are available for most 

coastal regions (Murty, 1977). 

The characteristics of distant tsunami waves 

depend on the generation mechanism and other 

influences of the source region, modifications during 

oceanic propagation, and coastal modifications in the 

vicinity of the site in question (Murty, 1977). Because 

records are not commonly available on the open ocean, 

the oceanic characteristics can be estimated only 

qualitatively from what is known of the generation 

mechanism and from observed properties of the waves 

at coastal sites. As a tsunami approaches a coast, the 

tsunami is modified extensively. The continental slope 

tends to reflect energy back to sea; the con tin en tal 

shelf tends to trap energy at wavelengths on the order 

of the shelf width; the coastal zone acts as a filter 

with characteristics determined by the coastal 

configuration and bathymetry (Cox, 1972); and, finally, 

the topography of the coastline influences wave runup 

on shore. 

To mitigate potential coastal hazards, either 

through selective siting to minimize exposure or 

through design to accommodate effects, tsunami wave 

heights and runup distances must be predicted on the 

basis of the historical record and numerical modeling 

(Houston, 1979; Camfield, 1980). If the historical 

record is adequate to establish the frequency of 

tsunamis of various sizes, then the probability that a 

tsunami wave will exceed a given height over some 

future time interval can be calculated. A probabilistic 

tsunami map for the contiguous United States and 

Alaska (fig. 41) shows the tsunami hazard to be greater 

on the Pacific coast than on the Atlantic, Gulf, or 

Arctic coasts, and greatest along the Gulf of Alaska. 

Detailed hazard maps for California, Oregon, 

Washington, and Hawaii were presented by Houston 

(1979, figs. 3-11). 

For a site at which significant tsunamis are 

expected, the last line of defense is shutdown and 

evacuation after early warning of an impending 

tsunami. The Pacific region is served by the Pacific 

Tsunami Warning Center in Honolulu (Spaeth, 1975), 

which issues two basic types of bulletins. "Watch" 

bulletins are issued when an earthquake has been 

detected that is of sufficient magnitude and in such a 

location that generation of a tsunami is possible. 

"Warnings" are issued when the Center has received 

positive evidence, from tide-gage stations in the 

source region, that a tsunami has been generated; such 

warnings contain the estimated times of arrival of the 

principal wave at the participating stations. Major 

offshore petroleum operations in the coastal Pacific 

would be linked to this warning system, either directly 

or through a local participating agency. 

Regional tsunami-warning systems are also in 

operation in areas with a history of damaging 

tsunamis; the most highly developed such systems are 

in Alaska (Butler, 1971; Sokolowski and others, 1984) 

and Japan (Ichikawa and Watanabe, 1983). These 

systems have data from regional seismograph and tidegage stations telemetered to a central location. 

Nearby, potentially tsunamigenic earthquakes are 

located rapidly, generally within 15 minutes, and a 

warning is issued to the local area. Because these 

warnings are based on seismic data alone, they are 

commonly issued when tsunamis have not actually been 

generated. Confirmation of tsunami generation, 

however, is soon obtained, and so inconveniences due 

Evaluation and Mitigation of Earthquake Hazards 55 
BER ING SEA 

ZONE 2 

ARCTIC OCEAN 

ZONE 1 

(EXCEPT ALEUT IAN ISLANDS) 

c;:>c:D ALEUTIAN ISLANDS 

!SAME AS GULF OF ALASKA) GULF OF ALASKA 

ZONE 3 

(EXCEPT END OF INLETS AND 

F lOADS THAT AAE IN ZONE 

4 ANO POSSIBLY ZONE 5) 

ZONE 1 0 TO 5 FT 

ZONE 2 5 TO 15 FT 

ZONE 3 15 TO 30FT 

ZONE 4 30 TO 50FT 

ZONE 5 50FT OA GREATER 

ALASKAN PANHANDLE 

ZONE 2 

(EXCEPT ENDS OF INLETS AND 

F lOADS ON PAC IFIC SlOE OF 

ISLANDS THAT AAE IN ZONE 3) 

Figure 41. Probabilistic map of tsunami elevation for 

the contiguous United States and Alaska at a 10-

percent probability of exceedence in 50 years, or a 

return period of about 500 years (from Houston, 1979, 

fig. 2). 

to false warnings are minimal, given the high level of 

protection provided by the system. 

CONCLUSION 

This report has provided a general overview of 

the subject of earthquakes and their effects at offshore sites. For examples, we have drawn on our 

experience and some of the pertinent literature 

pertaining to earthquakes in Canada and the United 

States and their adjacent offshore areas. A glance at 

a global-seismicity map reveals that similar seismic 

conditions prevail in numerous areas worldwide where 

an offshore petroleum potential exists in geologic 

environments similar to those of North America. The 

techniques for assessment of earthquake potential and 

the evaluation and mitigation of earthquake hazards 

discussed here pertain equally to these areas. 

The subjects discussed here are an inexact 

science. The earthquake potential of a given region, 

56 Earthquake Hazards in the Offshore Environment 

such as the Gulf of Alaska, may be obvious to anyone 

working in the Earth sciences. However, the effects 

that these earthquakes will have in areas of 

exploitable petroleum resources are not so obvious, 

even to the specialist. The offshore environment is a 

frontier not only for petroleum development but also 

for the evaluation of earthquake effects and their 

attendant hazards to marine facilities. Mitigation of 

potential hazards requires state-of-the-art investigations and good measures of scientific, engineering, and 

socioeconomic judgment. 

Potential earthquake hazards are not restricted 

to tectonically active continental margins that have 

undergone numerous historical earthquakes. Significant parts of this report have dealt with the passive 

margins of eastern North America, where evaluation 

and mitigation are made more difficult by our inability 

to explain the earthquakes there in simple global 

tectonic terms. 

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INTENSITY, MAGNITUDE, AND SEISMIC MOMENT 

Before the existence of instruments to record 

seismic waves, earthquakes were ranked by size in 

terms of the severity of the effects they caused. 

Early seismologists introduced the parameter of intensity as a qualitative measure of the effects of an 

earthquake upon people, structures, and natural features. Several descriptive scales were adopted, 

defining levels of intensity in terms of various earthquake effects. The intensity scale most widely used 

today in the United States and Canada is the modified 

Mercalli scale (Richter, 1958, p. 137-138), which has 

12 levels of intensity ranging from imperceptible 

shaking to catastrophic damage. Although intensity is 

a qualitative measure, it is expressed numerically, 

generally by Roman numerals; the larger numbers 

correspond to the more severe effects. The severity 

of effects varies with location, primarily as a function 

of distance from the earthquake source; higher intensities generally occur closer to the source. Thus, 

maximum observed intensity can serve as an indicator 

of the relative strength of an earthquake, except in 

those events for which no observations have been 

obtained close to the source because the earthquake is 

either shallow but distant, or close but deep. To 

overcome these limitations and the subjectivity 

involved in assigning intensities, an instrumental scale 

of earthquake size was introduced by C.F. Richter in 

1935. 

The instrumental-magnitude scale is based on the 

principle that the amplitude of recorded seismic waves 

reflects the strength of the earthquake at its source. 

If two earthquakes of different size occur in the same 

place and are recorded by the same seismograph, the 

larger shock will cause larger amplitudes on the 

seismograms. Richter (1935) defined earthquake 

magnitude as the common logarithm of the maximum 

amplitude recorded by a standard seismograph, 

corrected to a fixed reference distance to compensate 

for the attenuation of ground motion with distance 

from the source. By definition, a shock of M=3 would 

generate a maximum trace amplitude of 1-mm on a 

standard horizontal seismograph (magnification, 2,800; 

free period, 0.8 s; damping coefficient, 0.8) at a 

distance of 100 km from the epicenter. Richter's 

initial magnitude scale was derived for shocks 

recorded to distances of 600 km in southern California, 

where earthquakes are less than 15 km deep. This 

scale is commonly referred to as the Richter or localmagnitude scale, and values derived from it are commonly designated M • Subsequently, other magnitude 

scales were defined ror application to more distant and 

deeper earthquakes, and to earthquakes in regions 

where the efficiency of seismic-wave transmission 

differs from that in southern California. 

The seismic waves generated by an earthquake 

can be grouped approximately into two types: body 

waves, which penetrate deep within the Earth, and 

surface waves, which propagate near the surface. 

Long-period surface waves, which attenuate more 

slowly with distance than do body waves, are a 

dominant feature of seismograms from distant 

shallow-focus earthquakes. Accordingly, a surfacewave-magnitude scale was derived on the basis of 

surface waves with periods of about 20 s recorded at 

distances beyond 2,000 km; these magnitudes are 

designated M • Earthquakes deeper than several tens 

of kilomete~ do not excite large 20-s surface waves, 

and so a magnitude scale based on the amplitude of 

body waves was derived for application to both deep 

and shallow earthquakes. In current practice, bodywave magnitudes, ~' are based on the amplitude of 

compressional (P) waves of about 1-s period, recorded 

to distances as great as several thousand kilometers. 

Short-period P-waves undergo abnormal attenuation in 

the upper mantle in the depth interval 75-200 km 

beneath the Western United States, and so the Wb 

values calculated for Western United States earthquakes are diminished relative to Central or Eastern 

United States shocks. Nuttli (1973) developed a shortperiod-magnitude scale based on the amplitude of 

higher mode surface waves. Because higher mode 

Intensity, Magnitude, and Seismic Moment 67 
surface waves do not penetrate into the upper mantle, 

their amplitudes are unaffected by variations in the 

attenuation properties of the upper mantle. 

Magnitudes derived from this scale are designated 

.!!!bL!l' where the subscript b indicates that the 

measured waves are of 1-s period and Lg designates 

the type of surface waves used. 

When magnitude scales were first introduced, 

seismologists hoped that all such scales would give the 

same number for a given earthquake. However, such 

agreement between scales is impossible because the 

spectral content of earthquakes varies with earthquake 

size and different scales measure the amplitudes, or 

energy, in different spectral bands. M , m , and M 

measure spectral amplitudes at periods o~ 0~1, 1, a~ 20 s, respectively. Because the ratio of long- to shortperiod amplitudes increases with earthquake size, the 

ratio of M to~' for example, increases with earthquake si~ Thus, for a given earthquake, different 

magnitude scales generally yield different values. In 

describing the size of an earthquake, not only the 

magnitude but also the scale to which it pertains must 

be specified. As a corollary, it must be recognized 

that no single magnitude can adequately describe the 

size of an earthquake, in the sense of defining the 

spectrum of seismic motion at the source. 

The magnitude scales discussed above are 

observed empirically to saturate at larger magnitudes; 

above some saturation threshold, earthquakes do not 

yield larger magnitude values as they increase in 

size. The m scale saturates near 7, and the M scale 

near 8.5. i?rthquakes substantially larger th~~ =7, 

for example, register about m =7 regardless of their 

actual size, because the amp-Ift'ude of 1-s waves fails 

to increase with the increasing size of the earthquake. This saturation arises from the fact that the 

wavelength for a 1-s-period P-wave is several 

kilometers, typically 5 to 8 km, whereas the length of 

fault rupture for an earthquake larger than .!!!b =7 may 

be a few to many times larger. Thus, waves of 1-s 

period cannot reflect the gross faulting characteristics 

of the entire earthquake but, instead, those of only a 

fraction of the entire rupture. In contrast, at wavelengths longer than the rupture length, the earthquake 

source appears to be a point; the rupture process 

appears to be relatively coherent, and the amplitudes 

of the long-period waves truly reflect the size of the 

earthquake. 

Seismic moment, .M.o' is another measure of 

earthquake size, whose values can be determined from 

the spectral amplitude of long-period body or surface 

waves-periods longer than the time required for the 

rupture to propagate to the end of the fault. Seismic 

moment can be viewed as a measure of the spectral 

amplitude of the earthquake source at zero 

frequency. Seismic moment is a more fundamental 

measure of the size of an earthquake than is magnitude because seismic moment relates size to the 

physical attributes of the faulting episode generating 

the earthquake. In terms of faulting parameters, 

seismic moment equals the product ~D, where l.1 is the 

modulus of rigidity of the rock containing the fault, A 

is the area of the fault rupture, and D is the average 

displacement over the rupture. Thus,seismic moment 

relates the static properties of fault rupture to the 

68 Earthquake Hazards in the Offshore Environment 

long-period level of radiated seismic waves. Seismic 

moment permits an estimation of the average slip on a 

buried fault from seismograms of an earthquake if the 

area of the rupture can be determined independently, 

for example, either from the spatial distribution of 

aftershocks or from the radiation properties of longperiod seismic waves. Conversely, for historical 

earthquakes that occurred before seismographs were 

operating, the long-period level of shaking can be 

inferred from observations of the length of surface 

faulting and the amount of fault slip, assuming a value 

for the width or depth of faulting. 

Kanamori (1977) and Hanks and Kanamori (1979) 

defined a magnitude scale based on seismic moment, 

where the moment magnitude M is defined by the 

relation M=(2/3)log .M.o-10. 7. The moment-magnitude 

scale has the desirable features that it does not 

saturate for large shocks and that it provides a means 

for determining earthquake magnitude from geologic 

observations of surface faulting or from geodetic 

measurements of crustal deformation. For these 

reasons and because of the fundamental physical 

significance of seismic moment, the moment 

magnitude scale is increasingly used as a simple index 

of earthquake size in preference to other scales. 

M and .!!4J' however, remain appropriate scales for 

esbmating strong ground motion close to the earthquake source for the design of structures, because 

these scales directly measure the amplitude of ground 

motion at periods near 1 s, close to the resonant 

periods of many important structures. 

Table 3 lists magnitudes and seismic moments 

for the earthquakes specifically discussed in this 

report. Where available, different magnitudes are 

included to illustrate the differences in values 

obtained with different scales. The moment 

magnitudes were computed from the listed seismic 

moments, except for four midplate earthquakes for 

which they were computed from m or M magnitudes 

using the empirical relations of~uttli11983). The 

magnitudes cited in the body of this report are 

moment magnitudes. 

For shocks that occurred before 1899, no instrumental recordings are available from which 

magnitudes can be calculated; however, J!!bL!l magnitudes for Central and Eastern United State'S events 

have been estimated on the basis of the observed 

patterns of intensities in relation to patterns for more 

recent shocks for which instrumental magnitudes are 

available; these estimates are listed in parentheses. 

The uncertainity in a given magnitude derived from 

the amplitudes of seismic waves recorded at several 

stations over a wide range of azimuths is generally less 

than 0.3. The saturation of magnitude scales is clearly 

seen in comparing the M and M values for the 1957 

and 1964 earthquakes annhe ~m , M , and M values for -s - the 1965 shock; all these even s were large Alaskan 

shocks associated with subduction along the Aleutian 

Island arc. 

The values of seismic moment were determined 

from many types of data. For many events, the listed 

magnitude is the average of several values based on 

different data sets. The moment for the 1811 shock 

was estimated from the area within which damage 

occurred (modified Mercalli intensity of VI or 
Table 3. Magnitudes and moments for earthquakes discussed in this report 

[Magnitudes in parenthesis are estimated from intensity patterns. Bases for determining moment: A, area of aftershock zone; C, crustal 

deformation; F, surface faulting; I, area subjected to modified Mercalli intensity of at least VI; S, seismograms. 1 N·m=107 dyne·cm. 

Numbers in brackets are references] 

Oate Location Moment 

( 1018 N·m) Basis 

Nov. 18, 1755 Off Cape Anne, Maine------------------- ( 5.8-6.0) [ 1] -------------------- 5.8-6.1 ------------------ Jan. LJ, 1H57 Fort Tejon, Cal if------------------ ___ ---- ______________________________________ 7.8-7.9 [ 2] 530-880 [ 3] 

Aug. 31' 1886 Charleston, S .C- ----------------------- ( 6.6-6.9) [4] -------------------- 6.9-7.3 ------------------ Sept. 4, 1899 Yakutat Bay, A las ka ------------------------------------------------- 8.5 [5] 8.1 1,500 [6] s Sept. 10' 1899 do---------------------------------------------------------------- 8.4 [ 5] 8.1 l ,000 [6] s Apr. 18, 1906 San Franc i sea, Cal if---------------------------------------------_-- 8-1/4 [7] 7. 7 [ 2] 350-430 [3] F 

Nov. 16' 1920 Off 8anks Island, Beaufort ------------------------------------- 6-1/2 [ 7] 6.2 ------------------ Sea. 

Nov. 4' 1 n7 Off Lompoc, Cal i f--------------------------------------------------- 7. 3 [7] 7. 3 [2] 100 [8] S, I 

Nov. 18, 1929 Grand Ranks off Newfound- ------------------------------------- 7.2 [ 7] 6.5 6 [9] s land. 

Nov. 20, 1933 8affi n Bay---------------------------------------------------------- 7.3 [7] 6. 7 ------------------ Nov. 10, 1938 Alaska Peninsula, Alaska-------------------------------------------- 8.3 [7] 8.2 [10] 2,800 [10] s, A 

Aug. 22, 1949 Queen Charlotte Islands, -------------------------------------- 8.1 [7] 8.1 [ 10] 1,500 [ 10] 

British Columbia, Canada. 

July 21' 1952 Kern County, Calif---------- 7.2 [ 11] ----------------------- 7. 7 [7] 7. 5 [2] 200 [8] S, 

t~a r. q' 1957 Andreanof Islands, Alaska ------------------------------------- 8.1 [ 12] 9.1 [10] 58,500 [10] A 

July 10, 1958 Southeastern Alaska, ------------------------------------- 7. 9 [ 12] 7. 7 [ 12] 400 [ 13] s Alaska. 

------------------------------------------------- 8.2 [10] 2,900 [10] A 

t~a r. 28' 1964 Pr i nee William Sound, ------------------------------------- 8.4 [ 12] 9.2 [101 R?. ,000 [14] s, 

Alaska. 

June 18' 1%4 Niigata, Japan------------------------------------------ 6. 2 [ 15] 7.4 [ 14] 7.6 290 [ 14] s Feb. 4, 1965 Rat Islands, Alaska------------------------------------- 6.1 [ 15] 8.2 [12] 8. 7 [ 10] 12,500 [ 14] s June 28' 1966 Parkfield, Calif---------·--- 5.6 [ 16] 5.0 [ 15] 6.4 [17] 6.2 2.6 [ 14] S, 

Aug. 9, 1Y67 Denver, Colo-------------------------------------------- 5 .o [151 4.4 [ 4] 4.8 .021 [ 18] s Fe h. 28' 1 CJ6lj 14est of Gibraltar, -------------------------- 6.5 [15] 7.8 [ 12] 7.8 [10] 600 [ 19] s Atlantic Ocean. 

Feb. 9' 1971 San Fernando, Calif--------- 6.4 [20] 6.2 [15] 6.6 [14] 6.6 [2] 12 [ 14] S, 

Aug. 13' 1LJ78 Santa Barbara, Calif-------- 5.1 [21] 5.4 [ 15] 5. 7 [15] 6.0 1.1 [22] 

Feb. 2H, 1979 Saint Eli as, A l as ka------------------------------------- 6.2 [15] 7.1 [ 15] 7.6 [23] 250 [24] 

Oct. 15' 1979 Imperial Valley, Calif------ 6.6 [25] 5.6 [15] 6.9 [ 15] 6.5 7 [26] 

Nov. 8, 1980 Eureka, Calif------------------------------------------- 6.2 [27] 7.2 [27] 7.4 130 [28] 

Oct. 28' 1983 Borah Peak, Idaho--------------------------_-------- ____ 6.2 [27] 7.3 [27] 6.9 25 [27] 

References: 1. Street and Lacroix (1979, p. 164). 16. Kanamori and Jennings (1978, p. 484). 

2. Hanks and Kanarnori (1979, p. ?.349). 

3. Sieh (1978b, p. 1447). 

4. Nuttl i and others ( 1979, p. 898, 907). 

5. Thatcher and Plafker (1977). 

fi. Wayne Thatcher (oral commun., 1983). 

7. Gutenberg and Richter (1954). 

fl. Hanks and others ( 1975, p. 1132). 

9. Street and Turcotte (1977, p. 601). 

10. Kanamori (1977, p. 2982). 

11. Bolt (1978, p. 514). 

12. Abe and Kanarnori ( 1980, p. 195). 

13. Ando (1977, p. 438). 

14. Kanarnori and Anderson ( 1975, p. 1078). 

15. International Seismological Centre, Monthly Bulletins. 

greater). For the 1857 and 1906 shocks on the San 

Andreas fault in California, moments were calculated 

from the length of surface faulting and the observed 

fault offset. For the 1949, 1957, and 1958 

earthquakes, moments were derived from an empirical 

correlation between moment and rupture area, in 

which the area was inferred from the spatial distribution of aftershock hypocenters. The moment based 

on the aftershock area for the 1958 shock is about 7 

times greater than that calculated directly from longperiod surface waves. This discrepancy indicates the 

17. Wu (1968, p. 691). 

18. Herrmann and others (1981, p. 736). 

19. Fukao ( 1973, IJ· 210). 

20. Allen and others (1975, p. 260). 

21. Whitcomb and Hutton (1978). 

22. Wallace and others (1981, p. 1710). 

23. Joyner and Boore ( 1981, p. 2017) • 

24. Hasegawa and others (1980, p. 1655). 

25. Chavez and others ( 1982, p. 53). 

26. Kanamori and Regan (1982, p. 57). 

27. U.S. Geological Survey, Monthly Listings, 

Preliminary Determination of Epicenters. 

28. Lay and others ( 1982, p. 439). 

difficulty of determining reliable moment estimates by 

means of indirect techniques. Experience with many 

well-studied earthquakes indicates that moment estimates from seismograms and from static fault parameters generally agree within a factor of 2. 

For further information about magnitude scales 

and the relations between scales and about earthquake-source mechanics and magnitude scales, the 

reader is referred to Chung and Bernreuter (1981) and 

Boore (1977), respectively. 

Intensity, Magnitude, and Seismic Moment 69 
GPO 587-044/45001