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3 Facing the Earthquake Threat Earthquakes rival all other natural disasters inthethreat they pose to human life and habitat. Unlike floods, hurricanes, and volcanic eruptions, specific earthquakes cannot be predicted with the short-term accuracy required for effective emergency management. The science is now capable of identifying where earthquakes will happen and how big they might be, but such forecasts are valid only for intervals measured in decades or even centuries. Once an event has occurred, there is very little time for warning and action; the fast-moving seismic ground waves do most of their damage in a macroseismic zone within the first minute or so after the rupture nucleates (1). Preparation and rapid emergency response are therefore the bulwarks of a good seismic defense. This chapter describes the context of current efforts to improve seismic safety and performance by summarizing what is known about the principal types of earthquake hazards, their distribution across the nation and the world, and the knowledge-based approaches to reducing earthquake risk. It concludes by addressing the issue of how scientists can help to implement the knowledge gained through research by stimulating civic actions that actually reduce risk. 3.1 TYPES OF SEISMIC HAZARDS Earthquakes pose several types of threats that often proceed as chain reactions. The primary hazards are the breaks in the ground surface caused when faults rupture, the seismic shaking radiated from the fault slip dur-
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ing rupture, and the permanent subsidence and uplift. Strong ground motion may, in turn, cause ground failure—slumps, landslides, liquefaction, and lateral spread—depending on shaking intensity (usually stronger nearer the source) and local site conditions. If it occurs offshore, fault displacement can generate tsunamis capable of inundating nearby and distant shorelines. Ground failure and tsunamis are examples of secondary hazards (2). Fault Rupture Tectonic earthquakes are spontaneous releases of tectonic stress that produce macroscopic, permanent displacements across fault surfaces (ruptures) and within the rock mass around faults (co-seismic deformations). Most fault ruptures are confined to buried regions of the crust where brittle behavior allows stick-slip instabilities to nucleate (e.g., between 2 and 20 kilometers deep in most continental deformation zones). Such ruptures propagate to the surface only in larger earthquakes. When this happens, however, almost any structure built across the rupture path will be deformed by the severe strains characteristic of primary ground failure (Figure 3.1). Predicting the magnitude and extent of fault rupture is therefore a major issue in seismic hazard analysis. Ruptures tend to occur along faults that have produced large earthquakes in the past, so a map of active faults is a first-order representation of the rupture hazard. The average amount of co-seismic slip increases systematically with earthquake magnitude (3), and the maximum displacement tends to occur toward the middle of the rupturing segment. These behaviors can be used to quantify the hazard along well-defined active faults. For example, where the Hollywood subway crosses the Hollywood fault in Los Angeles, California, the maximum expected slip is estimated to be 1 to 2 meters. In anticipation, the Metropolitan Transportation Authority overbored the subway tunnel to allow the tracks to be realigned after such an earthquake. Mapped faults are often categorized as active and inactive, but doing so is problematic because the maximum magnitude, frequency of rupture, and other measures of activity can be highly variable among faults in the same tectonic province. Even within a single zone, the distribution of recent faulting can be considerably more complex than the simple traces that represent active faults on small-scale geologic maps. Detailed mapping reveals a wide range of features, such as segmentation, stepovers, and faulting at conjugate angles, often with self-similar scaling (Figure 3.2). The faulting patterns observed in large earthquakes show similar complexity, which can vary rapidly along strike. In some places, the rupture may be a single, clean break, while elsewhere it may occupy a zone tens or hundreds
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FIGURE 3.1 Damage due to fault rupture during the September 20, 1999, Chi-Chi, Taiwan, earthquake (magnitude [M] 7.6) was extensive along the Chelungpu fault. Structures that were built to withstand strong ground motion nevertheless did not survive severe dislocations along the fault. The left abutment of this bridge across the Ta-An River was constructed through the fault plane (top). The thrust fault slipped about 10 meters during the earthquake, severely deforming the pillar and, thus, destroying the bridge. After the earthquake, the abutment was reconstructed in the same location. Most reinforced multistory concrete structures survived the shaking, but those on the fault trace collapsed or suffered severe tilts and other distortions, which rendered them uninhabitable (bottom). SOURCE: Photographs courtesy of Kerry Sieh, Caltech.
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FIGURE 3.2 Map of the surface trace of the 1968 Dasht-e-Bayez, Iran, earthquake rupture (M 7.3). SOURCE: J.S. Tchalenko and M. Berberian, Dasht-e-Bayez fault, Iran: Earthquake and earlier related structures, Geol. Soc. Am. Bull., 86, 703-709, 1975. of meters wide comprising en echelon offsets, anastomosing fractures, mole tracks, nonbrittle warping, and other types of co-seismic deformation (4). A more complete characterization of the rupture hazard will require a better understanding of how the distribution of surface breaks depends on the details of the fault slip at depth and how fault movements interact with a variety of structural factors, including topography, near-surface sedimentary layering, and fault-zone complexity. Ground Shaking Ground shaking is typically the primary cause of earthquake damage to the built environment. Shaking occurs during the passage of seismic waves as they propagate away from the rupturing fault. The most destructive shaking is usually the horizontal ground motion from S waves and surface waves, although the vertical component of motion can also excite a damaging structural response. The severity of the shaking is typically measured by the peak ground acceleration (PGA) or peak ground velocity (PGV), as recorded on strong-motion seismographs in the free field (i.e., on open ground away from buildings or other structures), or by the spectral response of a standard oscillator, either spectral acceleration Sa or spectral velocity Sv, calculated from the observed “time history” of
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the shaking (5). Measurements by strong-motion instruments near large earthquakes have shown that the time histories can be complex and can vary rapidly from place to place, especially at high frequency where interference effects are typically strong. Seismic waves are generated by fault slip during an earthquake. The distribution of the slip in space and time determines the radiation pattern (i.e., how the wave amplitudes vary with direction away from the fault). Small earthquakes can usually be approximated by the beachball-looking radiation patterns described in Section 2.3, but larger events show significant complications and asymmetries. For instance, the propagation of a rupture along a fault may produce a directivity pulse of coherent, high-amplitude shear motion at locations in the propagation direction. Rupture directivity effects, amplified by basin-edge effects, were the primary cause of the damage in the 1995 Hyogo-ken Nanbu earthquake. During the 1994 Northridge earthquake, the ground motion at frequencies below 2 hertz was observed to be highest at locations around the top edge of the fault to the north of the hypocenter, consistent with the directivity pulse expected from the Northridge rupture. At higher frequency, the radiation of waves from fault surfaces becomes less coherent, owing to small-scale fluctuations in fault slip and nearby material irregularities, causing the rupture directivity effect to become subdued and other effects such as
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proximity to the fault to dominate. In the Northridge earthquake, ground-motion amplitude greater than 2 hertz was observed to be highest above the hypocenter on the hanging wall side (6), not around the top edge of the fault, which experienced the longer-period directivity effects. Systematic differences in ground motion have been observed for different faulting types (7), but as yet no clear explanation of this exists. Elucidating how faults radiate seismic energy across the entire frequency band relevant to earthquake engineering (0.1-10 hertz) is a research challenge of major importance. The amplitude of seismic waves generally decreases as the waves propagate away from the source (as required to conserve energy), but the measurement from large earthquakes always exhibits a high degree of scatter. In seismic hazard analysis, the decay of ground-motion intensity with distance is represented by an attenuation relation, usually derived by fitting smooth functions to the scattered data (see Section 2.7). An objective of current research is to explain the variations in shaking intensity through a more fundamental understanding of the wave propagation process (8). Important physical effects include refraction by variation in the seismic velocity, reflection from surfaces of material discontinuity, and damping by the anelastic response of the rock and soil media. Some of the strongest variations are associated with horizontal layering of the crust and upper mantle. In the 1989 Loma Prieta earthquake, shear waves, critically refracted from the M discontinuity at the base of the crust (SmS waves), were partially responsible for the shaking that damaged parts of San Francisco nearly 90 kilometers from the epicenter (9). Data from aftershocks of the 1994 Northridge earthquake demonstrated that reflections from midcrustal interfaces can increase the shaking from shallow sources at certain shorter distances. Seismic waves can be amplified or attenuated by three-dimensional structures such as fault-bounded blocks and sedimentary basins. Earthquakes can excite resonance in the deep basins, shaking the soft sediments like jelly. A striking example was the massive destruction and loss of life during the 1985 Michoacan earthquake (moment magnitude [M] 8.0) in the parts of Mexico City underlain by soft, lake-bed clays. The source was in a subduction zone more than 350 kilometers away, which under normal circumstances would have caused little damage; however, sediment resonance was observed to amplify the spectral acceleration at low frequencies (about 0.5 hertz) by factors as large as 8 to 50 times relative to hard-rock sites (10). Other mechanisms for amplification include the focusing of waves by lens-like structures (11) and the generation of surface waves by the fault-bounded edges of sedimentary basins. Basin-edge effects of the latter type were partly responsible for the extreme damage to the Japanese city of Kobe in the 1995 earthquake (Box 2.4).
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The response of soils at shallow depth to strong shaking is a complex phenomenon (12). Amplitude builds as the waves slow (another consequence of energy conservation), so seismic shaking is typically amplified in the soft soils and unconsolidated sediments near the ground surface, where the wave speed can be much lower than in hard rock. For this reason, the average shear velocity in the upper 30 meters or so has become the primary basis for the National Earthquake Hazard Reduction Program (NEHRP) site classification used in many building codes, including the 1997 Uniform Building Code (UBC), 2000 International Building Code (IBC), and 2000 American Society of Civil Engineers (ASCE) Standards 7-98 (13). The high amplitude predicted by the linear wave theory is thought to be reduced by the nonlinear response of the unsaturated near-surface layers. Laboratory tests clearly demonstrate nonlinear strain behavior in soils under dynamic loading, but the importance of nonlinearity during actual earthquakes continues to be debated. Using available ground-motion data to differentiate nonlinear strain behavior from other wave propagation effects has usually been difficult. For example, interpretations of the data collected in Mexico City from the 1985 Michoacan earthquake reached conflicting conclusions on the importance of the nonlinearity of the city’s soft clay deposits (14). On the other hand, direct evidence of significant nonlinear soil response was clearly observed in the motions recorded by surface and subsurface (borehole) instruments at saturated sandy sites that liquefied during the 1987 Superstition Hills, California, and the 1995 Hyogo-ken Nanbu earthquakes (15). Aside from these extreme cases where the soil failed, indirect evidence of nonlinear site response on soils that remained stable during strong shaking is becoming more apparent with the greater number of seismograms being recorded in strong-motion arrays throughout the world (16). However, more of these data are clearly needed to better understand and predict this phenomenon. Another interesting aspect of seismic shaking is that it can vary substantially from one tectonic setting to another. For example, the motion from similar-sized earthquakes is observed to be stronger in the central and eastern United States than west of the Rocky Mountains. Felt areas and areas of specific intensity (isoseismals) are also larger for earthquakes in the central and eastern United States compared to those of earthquakes with similar magnitudes in the western United States. Earthquakes in the older, stronger regions of the continent generally have greater stress drops and therefore radiate more high-frequency energy for a given amount of fault slip; moreover, their seismic waves propagate with less attenuation compared to earthquakes in plate boundary deformation zones. The attenuation difference is probably attributable to lower temperature, reduced scattering, and more continuous waveguide for crustal shear en-
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ergy (Lg waves) in the more stable crust of the central and eastern United States. Regional studies that deploy seismometers more densely will be needed to clarify these explanations and to understand how the vertical and lateral structure of the crust controls ground motions. Subsidence and Uplift Large thrust earthquakes in subduction zones can cause sudden, permanent elevation changes with damaging effects to coastal areas. Uplift and subsidence related to fault slippage on shallow thrusts have been documented in New Zealand, Japan, Chile, and southeast Alaska (17). During the great 1964 Alaska earthquake (Box 2.3), the shorelines of Prince William Sound rose in some places by several meters, draining small-craft harbors, while they dropped in others, causing the streets of coastal towns to flood at high tide. Submerged marshlands in several estuaries along the coasts of Washington, Oregon, and northern California indicate that similar episodes of sudden subsidence have resulted from large thrust events in the Cascadia subduction zone (described in Section 3.2). The pattern of uplift and subsidence during an earthquake can be predicted from elastic dislocation models if the area of the fault plane and the distribution of slip within that plane are known (18). Anticipating the damage from elevation changes in future events can thus be approached by combining theoretical studies with seismic, paleoseismic, and geodetic observations. Secondary Ground Failures The secondary hazards caused by seismic shaking include forms of mass wasting—such as landslides, rockfalls, and slumps—as well as soil failures associated with compaction, liquefaction, and lateral spreading (19). In some instances, these failures cause more damage than the ground shaking itself. An M 8.6 earthquake in China’s Gansu Province in 1920 triggered an extensive debris flow, which covered a region larger than 100 square kilometers and resulted in roughly 200,000 deaths. An immense rock and snow avalanche (60 million cubic meters) triggered by the 1970 Peru earthquake (M 8.0) buried the mountain towns of Yungay and Ranrahirca, killing 66,000 people (Figure 3.3). Many of those killed in the January 13, 2001, El Salvador earthquake were buried by a muddy landslide loosened from a slope in the capital’s suburbs. Liquefaction is the temporary conversion of water-saturated, unconsolidated soils into a medium that behaves like a fluid. It occurs when saturated sand or silty sand is shaken hard enough to mobilize individual grains. If the water cannot escape the granular soil matrix fast enough to
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FIGURE 3.3 Destruction of the mountain towns of Yungay and Ranrahirca, Peru, which were buried by an avalanche triggered by the 1970 earthquake (M 8). SOURCE: Photo by Servicio Aerofotografico Nacional de Peru; available from the U.S. Geological Survey, <http://landslides.usgs.gov/>. permit compaction, more of the overburden load becomes supported by the water, resulting in increased pore pressure. This process can progress relatively quickly to the point at which the pore-water pressure becomes equal to the overburden stress, creating quicksand-like conditions. The liquefaction potential of any particular saturated deposit depends primarily on the age and grain-size distribution of the deposit as well as the ampli-
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FIGURE 3.4 Tilting of apartment buildings at Kawagishi-Cho, Niigata, Japan, produced by liquefaction of loose, water-saturated sediments caused a loss of load-bearing capacity during the 1964 Niigata earthquake (M 7.5). The losses from this earthquake exceeded $1 billion in 1964 dollars. SOURCE: National Oceanic and Atmospheric Administration, National Geophysical Data Center. tude and duration of the ground shaking (20). The dangers of liquefaction are thus compounded in deep sedimentary basins, where the water table is often shallow and the shaking amplitude and duration tend to be increased by seismic-wave resonance within the basins. Liquefaction can severely damage foundations and other subsurface structures, causing large buildings to sink or tilt (Figure 3.4) and underground structures, such as pipelines and storage tanks, to float to the surface when they become buoyant in the liquefied soil. If the liquefied layer is close to the surface, it may break through dryer deposits overlying the water table, forming geysers that leave sandblows as postseismic evidence. In fact, the dating of such features has become an extremely useful tool for establishing prehistoric records of major earthquakes in the Charleston and New Madrid areas of the eastern and central United States (21). Lateral spreading is a form of landsliding caused when liquefaction occurs on a sloping surface or adjacent to an embankment or excavation, typically resulting in the opening of fissures perpendicular to the surface gradient. Embedded structures are dragged by the flow, and the variable
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FIGURE 3.5 Many apartment buildings on the fringe of the river delta at Golcuk, Turkey, slid into the sea during the Izmit earthquake of August 17, 1999. Land-use planning that accounts for the instability of young deltaic sediments could substantially reduce loss of life and property in future earthquakes. SOURCE: A. Barka, Istanbul Technical University. displacements can literally rip structures apart. Lateral spreading tends to allow material to fill topographic depressions, such as streams and rivers, causing the channels to narrow and the flow to become dense—a major source of damage to bridges during earthquakes (22). A more recent case of lateral spreading occurred during the August 17, 1999, Izmit, Turkey earthquake (M 7.4) when unconsolidated, water-saturated deltaic sediments collapsed into the sea (Figure 3.5), resulting in numerous deaths. Lateral spreading or landsliding can also be caused by the shaking-induced loss of shear strength in certain types of “quick” or “sensitive” layers of salt-leached, clay-rich marine sediments. The spectacular damage to the Turnagain Heights district of Anchorage during the great 1964 earthquake (M 9.2) (Box 2.3) has been attributed to large (150- to 180-meter) displacements within a relatively thin zone of the Bootlegger Cove clay, 25 meters below the surface (23). Empirical relations for predicting the extent and severity of liquefaction events have been developed through field studies, theoretical modeling, and laboratory experiments using geotechnical centrifuges. Less ex-
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ing major centers to promote interdisciplinary collaborations and deliver new products for seismic hazard analysis (110). The primary participants in these programs have been geoscientists, with engineers relegated to a relatively minor role. The Federal Emergency Management Agency leads NEHRP and holds the federal responsibility for seismic risk mitigation. Although FEMA is deeply involved in both risk assessment and emergency management, it is primarily a user, not a coordinator, of earthquake research. At present, no agency or organization is responsible for ensuring an integrated approach to earthquake science and engineering. This vacuum could be filled through structured collaborations between the science and engineering research centers, explicitly funded and frequently reviewed by NEHRP agencies. These collaborations should involve economists and social scientists with expertise in mitigation issues. Scientists also need better organizational and technological support for communicating with all levels of society about earthquake hazards, mitigation measures, and the appropriate use of earthquake information. The challenges are to select the right kinds of educational activities, target the appropriate audiences, and present them at the right places and times. Appropriately, NEHRP agencies are now placing more emphasis on efforts to interpret scientific research and reduce the results to understandable, usable products. Even if well-packaged, however, such products cannot be “simply thrown over the wall” for public consumption. Effective communication between researchers and end users requires a two-way, continuing dialogue with repeated opportunities for the exchange of ideas and plans. Likewise, effective public education requires interactive mechanisms that can engage an audience at an appropriate level. The new technologies of the Internet—interactive web pages backed by powerful, simple-to-use query languages and digital libraries with up-to-the-minute earthquake information—offer considerable promise. However, their utilization will depend on support structures with more financial and human resources than a typical research group. NOTES 1. As described in Section 1.2, the warning times for destructive tsunamis that cross wide ocean basins can be several hours or more. 2. Secondary hazards also include fires, dispersal of nuclear materials, and other threats indigenous to the built environment. 3. The displacement across the fault scales with the cube root of seismic moment for earthquakes of magnitude less than about 6.5 and with the square root of moment for larger events. 4. A complete discussion of these complexities is given in R.S. Yeats, K. Sieh, and C.R. Allen, The Geology of Earthquakes, Oxford, New York, 568 pp., 1997. 5. Throughout the engineering literature, the redundant term “time history” is used to describe ground motion as a function of time and is thus synonymous with seismogram,
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accelerogram, and seismic waveform. The spectral response method is described in Section 2.7. 6. N.A. Abrahamson and P.G. Somerville, Effects of the hanging wall and foot wall on ground motions recorded during the Northridge earthquake, Bull. Seis. Soc. Am., 86, S93-S99, 1996. 7. See summary of ground motion models described by N.A. Abrahamson and K.M. Shedlock (Overview of ground motion attenuation models, Seis. Res. Lett., 68, 9-23, 1997). 8. E.H. Field and the SCEC Phase III Working Group, Accounting for site effects in probabilistic seismic hazard analyses of Southern California: Overview of the SCEC Phase III report, Bull. Seis. Soc. Am., 90, S1-S31, 2000. This major study of regional attenuation relations and local site effects concluded that “any model that attempts to predict ground motion with only a few parameters will have substantial intrinsic variability. Our best hope for reducing such uncertainties is via waveform modeling based on the first principles of physics.” 9. P.G. Somerville and J. Yoshimura, The influence of critical Moho reflections on strong ground motions recorded in San Francisco and Oakland during the 1989 Loma Prieta earthquake, Geophys. Res. Lett.17, 1203-1206, 1990. The ground motions recorded in San Francisco and Oakland were actually stronger than those for some closer sites with similar geology. 10. S.K. Singh, E. Mena, and R. Castro, Some aspects of source characteristics of the 19 September 1985 Michoacan earthquake and ground motion amplification in and near Mexico City from strong motion data, Bull. Seis. Soc. Am., 78, 451-477, 1988. 11. See S. Gao, H. Liu, P.M. Davis, and L. Knopoff (Localized amplification of seismic waves and correlation with damage due to the Northridge earthquake: Evidence for focusing in Santa Monica, Bull. Seis. Soc. Am., 86, S209-S230, 1996) for an example of focusing during the 1994 Northridge earthquake. 12. H.B. Seed and I.M. Idriss, Analyses of ground motions at Union bay, Seattle during earthquakes and distant nuclear blasts, Bull. Seis. Soc. Am.,60, 125-136, 1970; M. Zeghal and A.-W. Elgamal, Analysis of site liquefaction using earthquake records, J. Geotech. Engr.,120, 996-1017, 1994; E.H. Field, P.A. Johnson, I.A. Beresnev, and Y.H. Zeng, Nonlinear ground-motion amplification by sediments during the 1994 Northridge earthquake, Nature,390, 599-602, 1997; J. Aguirre and K. Irikura, Nonlinearity, liquefaction and velocity variation, of soft soil layers in Port Island, Kobe, during the Hyogo-ken Nanbu earthquake, Bull. Seis. Soc. Am., 87, 1244-1258, 1997. 13. NEHRP site classifications are rated on a five-level scale ranging from A (hard rock with measured shear-wave velocity [vS] more than 5000 feet per second) to E (soft soil with vS less than 600 feet per second); see Building Seismic Safety Council, 1997 Edition NEHRP Recommended Provisions for Seismic Regulations for New Buildings and Other Structures, FEMA 302/303, Part 1 (Provisions) and Part 2 (Commentary), developed for the Federal Emergency Management Agency, Washington, D.C., 337 pp., 1998. 14. Some have argued (W.D.L. Finn, Geotechnical engineering aspects of microzonation, in Proceedings of the Fourth International Conference on Seismic Zonation, August 25-29, 1991, Stanford, California, Vol. I, pp. 199-259, 1991; K. Aki, Local site effects on weak and strong ground motion, Tectonophysics, 218, 93-111, 1993) that a very low but approximately constant shear modulus and site resonance at Mexico City can explain the ground motions without appeal to nonlinear effects. Others (S.K.E. Singh, E. Mena, and R. Castro, Some aspects of source characteristics of the 19 September 1985 Michoacan earthquake and ground motion amplification in and near Mexico City from strong motion data, Bull. Seis. Soc. Am., 78, 451-477, 1988; C. Lomnitz, Mexico 1985; The case for gravity waves, Geophys. J. Int., 102, 569-572, 1990) argue for a strong nonlinear shear modulus reduction during the strong shaking. Recent efforts to measure dynamic strains at depth in the Valley of Mexico from
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other earthquakes (S.K. Singh, M.A. Santoyo, P. Bodin, and J. Gomberg, Dynamic deformations of shallow sediments in the valley of Mexico; Part II, Single-station estimates, Bull. Seis. Soc. Am., 87, 540-550, 1997; P. Bodin, S.K. Singh, M. Santoyo, and J. Gomberg, Dynamic deformations and shallow sediments in the Valley of Mexico, I: Three-dimensional strains and rotations recorded on a seismic array, Bull. Seis. Soc. Am., 87, 540-550, 1997) suggest that the Mexico City clays exhibit only a mildly nonlinear response, even up to strains of 1 percent. This strain is one to two orders of magnitude beyond the strains at which more common, more granular sediments begin to respond inelastically. Even though this may not increase strong-motion amplitudes, it often leads to other ground failure problems, the most common and most serious of which is liquefaction. 15. See, for example, A.-W. Elgamal, M. Zehal, and E. Parra, Liquefaction of reclaimed island in Kobe, Japan, J. Geotech. Engr., 122, 39-49, 1996; M. Zeghal and A.-W. Elgamal, Analysis of site liquefaction using earthquake records, J. Geotech. Engr.,120, 996-1017, 1994. 16. For a concise summary of nonlinear site response, see E.H. Field and SCEC Phase III Working Group, Accounting for site effects in probabilistic seismic hazard analysis of southern California; An overview of the SCEC Phase III report, Bull. Seis. Soc. Am., 90, S1-S31, 2000, and references therein. 17. G. Plafker, Tectonics of the March 27, 1964, Alaska Earthquake, U.S. Geological Survey Professional Paper 543-I, U.S. Government Printing Office, Washington, D.C., 74 pp., 1969. 18. J.C. Savage and L.M. Hastie (Surface deformation associated with dip-slip faulting, J. Geophys. Res., 71, 4897-4904, 1966) showed how a dislocation model could be used to fit Plafker’s observations of uplift and subsidence following the 1964 Alaska earthquake; see also S.R. Holdahl and J. Sauber, Coseismic slip in the 1964 Prince William Sound earthquake: A new geodetic inversion, Pure Appl. Geophys., 142, 55-82, 1994. 19. National Research Council, Liquefaction of Soils During Earthquakes, National Academy Press, Washington, D.C., 240 pp., 1985. 20. H.B. Seed and I.M. Idriss, Ground Motions and Soil Liquefaction During Earthquakes. Earthquake Engineering Research Institute, Engineering Monograph on Earthquake Criteria, Structural Design, and Strong Motion Records 5, El Cerrito, Calif., 134 pp., 1982. 21. P. Talawani and W.T. Shaeffer, Recurrence rates of large earthquakes in South Carolina coastal plain based on paleoliquefaction data, J. Geophys. Res., 106, 6621-6642, 2001; M.P. Tuttle and E.S. Schweig, Recognizing and dating prehistoric liquefaction features; Lessons learned in the New Madrid seismic zone, central United States, J. Geophys. Res., 101, 6171-6178, 1996. 22. During the 1964 Alaska earthquake, compression resulting from such dense flows buckled or skewed spans and damaged abutments on more than 250 bridges. See National Research Council, The Great Alaska Earthquake of 1964, National Academy Press, Washington, D.C., 15 volumes, 1972-1973. 23. W.R. Hansen, Effects of the Earthquake of March 27, 1964, at Anchorage, Alaska, U.S. Geological Society Professional Paper 542-A, Washington, D.C., 68 pp. + 2 plates, 1966. 24. Tsunami propagation can be treated by the theory of shallow-water waves, which states that the propagation speed varies as the square root of water depth. An elementary discussion of the tsunami physics is given by T. Lay and T.C. Wallace, Modern Global Seismology, Academic Press, San Diego, pp. 147-153, 1995. 25. Early in their history, tsunami warning systems generated many false alarms because they relied on earthquake size determined from high-frequency magnitude scales, such as mb. See H. Kanamori, Mechanism of tsunami earthquakes, Phys. Earth Planet. Int., 6, 346-359, 1972. 26. Kanamori and Kikuchi (The 1992 Nicaragua earthquake; A slow tsunami earthquake associated with subducted sediments, Nature, 361, 714-716, 1993) suggested that there are two types of tsunami earthquakes: those that arise from slow rupture, such as the 1992
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Nicaragua earthquake, which caused a destructive 10-meter runup on the Nicaraguan coast, and those such as the 1896 Sanriku and 1946 Unimak Islands earthquakes, which may have involved large-scale slumping. Another type of tsunami source is exemplified by the 1883 Krakatau eruption in Indonesia, which inundated 165 coastal villages and killed more than 30,000. 27. T.Y. Wu, Long waves in ocean and coastal waters, J. Engr. Mech. Div., 107, 501-521, 1981; P.L.F. Liu and J. Earickson, A numerical model for tsunami generation and propagation, in Tsunamis: Their Science and Engineering, J. Iida and T. Iwasaki, eds., Proceedings of the International Tsunami Symposium, Sendai-Ofunato-Kamaishi, Japan, May 1981, Terra Scientific Publ., Tokyo, pp. 227-240, 1983; M. Shibata, One-dimensional dispersive deformation of tsunami with typical initial profiles on continental topographies, in Tsunamis: Their Science and Engineering, J. Iida and T. Iwasaki, eds., Proceedings of the International Tsunami Symposium, Sendai-Ofunato-Kamaishi, Japan, May 1981, Terra Scientific Publ., Tokyo, pp. 241-250, 1983. 28. M.J. Briggs, C.E. Synolakis, G.S. Harkins, and D.R. Green, Laboratory experiments of tsunami runups on a circular island, Pure Appl. Geophys., 144, 569-593, 1995; S. Tinti, and C. Vannini, Tsunami trapping near circular islands, Pure Appl. Geophys., 144, 595-619, 1995. 29. P.L.F. Liu, C. Synolakis, and H.H. Yeh, Impressions from the first international workshop on long wave runup, J. Fluid Mech., 229, 675-688, 1991; H.H. Yeh, Tsunami bore runup, Natural Hazards, 4, 209-220, 1991; S. Tadepalli and C.E. Synolakis, Model for the leading waves of tsunamis, Phys. Rev. Lett., 77, 2141-2154, 1996. 30. E.P. Myers and A.M. Baptista, Finite element modeling of the July 12, 1993 Hokkaido Nansei-Oki tsunami, Pure Appl. Geophys., 144, 769-802, 1995; P.L.F. Liu, Y.S. Cho, S.B. Yoon, and S.N. Seo, Numerical simulations of the 1960 Chilean tsunami propagation and inundation at Hilo, Hawaii, in Tsunami: Progress in Prediction, Disaster Prevention and Warning, Y. Tsuchiya and N. Shuto, eds., Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 99-115, 1995. 31. G.F. Carrier, On-shelf tsunami generation and coastal propagation, in Tsunami: Progress in Prediction, Disaster Prevention and Warning, Y. Tsuchiya and N. Shuto, eds., Kluwer Academic Publishers, Dordrecht, The Netherlands, pp. 1-20, 1995. 32. F.I. Gonzalez, K. Satake, E.F. Boss, and H.O. Mofjeld, Edge wave and non-trapped modes of the 25 April 1992 Cape Mendocino tsunami, Pure Appl. Geophys., 144, 409-426, 1995. 33. A. Frankel, C. Mueller, T. Barnhard, D. Perkins, E. Leyendecker, N. Dickman, S. Hanson, and M. Hopper, National Seismic-Hazard Maps: Documentation June 1996, U.S. Geological Survey Open-File Report 96-532, USGS Federal Center, Denver, Colo., 111 pp, 1996; A. Frankel, C. Mueller, T. Barnhard, D. Perkins, E. Leyendecker, N. Dickman, S. Hanson, and M. Hopper, Seismic-Hazard Maps for the Conterminous United States, U.S. Geological Survey Open File Report 97-131, USGS Federal Center, Denver, Colo., 12 maps, 1997; A. Frankel, C. Mueller, T. Barnhard, D. Perkins, E. Leyendecker, N. Dickman, S. Hanson, and M. Hopper, Seismic-Hazard Maps for California, Nevada, Western Arizona/Utah, U.S. Geological Survey Open-File Report 97-130, USGS Federal Center, Denver, Colo., 12 maps, 1997. The maps and their documentation can be downloaded from <http://geohazards.cr.usgs.gov/eq/>. 34. The probability of exceedance in 50 years PE50 is related to the mean return period TR by the equation PE50 = 1 – (1 – 1/TR)50, so that the probabilities of PE50 = 2, 5, and 10 percent used in the national seismic hazard maps correspond to TR ˜ 2475, 975, and 475 years. The annual probabilities of exceedance are 1/TR ˜ 0.04 percent, 0.1 percent, and 0.2 percent, respectively. 35. In Guidelines for the Seismic Rehabilitation of Buildings (Building Seismic Safety Council and Applied Technology Council, FEMA Report 273, Washington, D.C., 400 pp., Octo-
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ber, 1997), the level of shaking defined by PE50 = 10 percent is called Basic Safety Earthquake 1 (BSE-1) ground motion, and the level of shaking defined by PE50 = 2 percent is called BSE-2 ground motion. In the NEHRP Recommended Provisions for Seismic Regulations for New Buildings and Other Structures (Building Seismic Safety Council, FEMA Report 368, Washington, D.C., 374 pp., March 2001), the PE50 = 2 percent shaking level is called the Maximum Considered Earthquake. Other hazard levels are routinely employed; for example, the California Seismic Safety Commission defines a “likely earthquake” by PE50 = 40 percent (about a 100-year return period) to represent the intensity likely to be experienced one or more times during a facility’s lifetime, and an “upper-bound earthquake” by PE50 = 5 percent (975-year return period) to represent the most severe shaking that could ever occur (EQE International, Earthquake Risk Management: A Toolkit for Decision-Makers, California Seismic Safety Commission Report 99-04, Sacramento, Calif., 185 pp., 1999). 36. R.K. McGuire, Probabilistic seismic hazard analysis and design earthquakes: Closing the loop, Bull. Seis. Soc. Am., 85, 1275-1284, 1995; S. Harmsen, D. Perkins, and A. Frankel, Deaggregated magnitudes and distances for probabilistic ground motions in the central and eastern U.S., Bull. Seis. Soc. Am., 89, 1-13, 1999; P. Bazzurro and C.A. Cornell, Disaggregation of seismic hazard, Bull. Seis. Soc. Am., 89, 501-520, 1999. Disaggregation tables for 100 U.S. cities can be downloaded from <http://geohazards.cr.usgs.gov/eq.> 37. Working Group on California Earthquake Probabilities, Earthquake Probabilities in the San Francisco Bay Region, U.S. Geological Survey Open-File Report 99-517, Reston, Va., 46 pp., 1999. 38. R. Bürgmann, D. Schmidt, R.M. Nadeau, M. d’Alessio, E. Fielding, D. Manaker, T.V. McEvilly, and M.H. Murray, Earthquake potential along the northern Hayward fault, California, Science, 289, 1178-1182, 2000. 39. T.L. Davis and J.S. Namson, A balanced cross-section of the 1994 Northridge earthquake, southern California, Nature, 372, 167-169, 1994. 40. Working Group on California Earthquake Probabilities, Seismic hazards in southern California: Probable earthquake 1994 to 2024, Bull. Seis. Soc. Am., 85, 379-439, 1995 (SCEC Phase II report). 41. T.H. Dixon, M. Miller, F. Farina, H. Wang, and D. Johnson, Present-day motion of the Sierra Nevada block and some tectonic implications for the Basin and Range Province, North American Cordillera, Tectonics, 19, 1-24, 2000; G. Peltzer, F. Crampé, S. Hensley, and P. Rosen, Transient strain accumulation and fault interaction in the eastern California shear zone, Geol. Soc. Am., 29, 975-978, 2001. 42. The case for great Chilean-type earthquakes in the Cascadian subduction zone was made by T.H. Heaton and H. Kanamori (Seismic potential associated with subduction in the northwestern United States, Bull. Seis. Soc. Am., 74, 933-941, 1984) based in part on previous geodetic observations by J.C. Savage, M. Lisowski, and W.H. Prescott (Geodetic strain measurements in Washington, J. Geophys. Res., 86, 4929-4940, 1981). Paleoseismic data supporting this hypothesis were first presented by B.F. Atwater (Evidence for great Holocene earthquakes along the outer coast of Washington State, Science, 236, 942-944, 1987). 43. B.F. Atwater, Geologic evidence for earthquakes during the past 2000 years along the Copalis River, southern coastal Washington State, J. Geophys. Res.,97, 1901-1919, 1992; D.K. Yamaguchi, B.F. Atwater, D.E. Bunker, B.E. Benson, and M.S. Reid, Tree-ring dating the 1700 Cascadia earthquake, Nature, 389, 922-923, 1997, correction in Nature, 390, 352, 1997. 44. K. Satake, K. Shimazaki, Y. Tsuji, and K. Ueda, Time and size of a giant earthquake in Cascadia inferred from Japanese tsunami records of January 1700, Nature,379, 246-249, 1996. 45. R.C. Bucknam, E. Hemphill-Haley, and E.B. Leopold, Abrupt uplift within the past 1,700 years at southern Puget Sound, Washington, Science, 258, 1611-1614, 1992; B.F. Atwater
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and A.L. Moore, A tsunami about 1,000 years ago in Puget Sound, Washington, Science, 258, 1614-1617, 1992; R.E. Karlin and S.E.B. Abella, Paleoearthquakes in the Puget Sound region recorded in sediments from Lake Washington, Science, 258, 1617-1620, 1992; R.L. Schuster, R.L. Logan, and P.T. Pringle, Prehistoric rock avalanches in the Olympic Mountains, Washington, Science, 258, 1620-1621, 1992; G.C. Jacoby, P.L. Williams, and B.M. Buckley, Tree ring correlation between prehistoric landslides and abrupt tectonic events in Seattle, Washington, Science, 258, 1621-1623, 1992. 46. R.A. Bennett, B.P. Wernicke, and J.L. Davis, Continuous GPS measurements of contemporary deformation across the northern Basin and Range Province, Geophys. Res. Lett., 25, 563-566, 1998; W. Thatcher, G.R. Foulger, B.R. Julian, J. Svarc, E. Quilty, and G.W. Bawden, Present-day deformation across the Basin and Range Province, western United States, Science, 283, 1714-1718, 1999. 47. R.B. Smith and L. Siegel, Windows into the Earth: The Geologic Story of Yellowstone and Grand Teton National Parks, Oxford University Press, New York, 242 pp., 2000. 48. T.C. Hanks and A.C. Johnston, Common features of the excitation and propagation of strong ground motion for North American earthquakes, Bull. Seis. Soc. Am., 82, 1-23, 1992. 49. S.F. Obermeier, P.J. Munson, C.A. Munson, J.R. Martin, A.D. Frankel, T.L. Youd, and E.C. Pond, Liquefaction evidence for strong Holocene earthquake(s) in the Wabash Valley of Indiana-Illinois, Seis. Res. Lett., 63, 321-336, 1992. Although small, nondestructive earthquakes are relatively common in the Wabash River Valley of southeastern Illinois and southern Indiana, no large earthquakes have struck the region in 200 or so years of historical record. Nevertheless, paleoliquefaction features indicative of very large earthquakes have been identified, including clastic dikes ranging up to 2.5 meters in width, are widespread throughout a region of about 200 kilometers by 250 kilometers. The most widespread set of these formed during a large event about 6100 years ago. The fact that the largest dikes cluster within a region about 50 kilometers in diameter suggests that the source of the earthquake was there, near the Illinois-Indiana border with a size of Mw about 7.5. More restricted sets of dikes formed in this same region during an event about 12,000 years ago and in a smaller region within Indiana about 3000 years ago. 50. A.J. Crone and K.V. Luza, Style and timing of Holocene surface faulting on the Meers fault, southwestern Oklahoma, Geol. Soc. Am. Bull., 102, 1-17, 1990. 51. A.J. Crone, M. Machette, L. Bradley, and S. Mahan, Late Quaternary Surface Faulting on the Cheraw Fault, Southeastern Colorado, U.S. Geological Survey Map I-2591, Reston, Va., 1997. 52. D. Amick and R. Gelinas, The search for evidence of large prehistoric earthquakes along the Atlantic seaboard, Science, 251, 655-658. 1991. Near Charleston, prehistoric sand-blow craters, similar to those that formed in 1886, formed four times in the 5000 years before 1886. Small twigs and bark that fell into these ancient craters soon after they were formed yield radiocarbon ages of about 600, 1250, 3200, and 5150 years. Judging from the size and geographic extent of the craters formed 600 and 1250 years ago, the magnitude of the causative earthquakes was at least Mw 7.5. 53. R.T. Marple and P. Tawani, The Woodstock lineament; A possible surface expression of the seismogenic fault of the 1886 Charleston, South Carolina, earthquake, Seis. Res. Lett.,63, 153-160, 1992. 54. C. Powell, G. Bollinger, M. Chapman, M. Sibol, A. Johnston, and R. Wheeler, A seismotectonic model for the 300-kilometer-long eastern Tennessee seismic zone, Science, 264, 686-688, 1994. 55. A. Frankel, C. Mueller, T. Barnhard, D. Perkins, E. Leyendecker, N. Dickman, S. Hanson, and M. Hopper, National Seismic-Hazard Maps: Documentation June 1996, U.S. Geological Survey Open-File Report 96-532, Denver, Colo., 111 pp, 1996.
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56. D. Giardini, G. Grünthal, K.M. Shedlock, and P. Zhang, The GSHAP Global Seismic Hazard Map, Ann. Geofisica, 42, 1225-1230, 1999. 57. The volcanism in subduction zones is caused primarily by water that is carried down with the subducted slab to depths on the order of 100 kilometers. This water eventually fluxes into the mantle wedge above the slab, lowering the melting temperature of the rocks and producing a small fraction of melt that rises into shallow magma chambers, which erupt to form the andesitic volcanoes characteristic of the island arcs. 58. S. Uyeda and H. Kanamori, Back-arc opening and the mode of subduction, J. Geophys. Res., 84, 1049-1061, 1979. 59. L. Ruff and H. Kanamori, Seismicity and the subduction process, Phys. Earth Planet. Int., 23, 240-252, 1980. 60. W.R. McCann, S.P. Nishenko, L.R. Sykes, and J. Krause, Seismic gaps and plate tectonics: Seismic potential for major boundaries, Pure Appl. Geophys., 117, 1082-1147, 1979. Offshore of Java, the subduction zone has produced only major earthquakes (M = 7.5) in the 450-year-long historic record, whereas offshore of Sumatra, giant earthquakes (M 8.5 to 8.8) have occurred. Future geodetic measurements across subduction zones will enable better quantification of the degree of coupling and, hence, better estimates of seismic potential. 61. T.J. Fitch, Plate convergence, transcurrent faults and deformation in Asia and Pacific, J. Geophys. Res., 77, 4432-4460, 1972. 62. A.Y. Le Dain, B. Robineau, and P. Tapponnier, The tectonic effects of the seismic and volcanic event of November 1978 in the Asia-Ghubbet Rift, Soc. Gèol. France Bull., 22, 817-822, 1979; T. Forslund and A. Gudmundsson, Crustal spreading due to dikes and faults in southwest Iceland, J. Struct. Geol., 13, 443-457, 1991; R.S. Stein, P. Briole, J.-C. Ruegg, P. Tapponnier, and F. Gasse, Contemporary, Holocene, and Quaternary deformation of the Asal Rift, Djibouti: Implications for the mechanics of slow spreading ridges, J. Geophys. Res., 96, 21,789-21,806, 1991. 63. Several million years of normal faulting in the million-square-kilometer Basin and Range Province have led to a northwest-southeast extension of more than 100 kilometers and the creation of dozens of tilted, 10- to 30-kilometer-wide crustal blocks that form the alternating basins and ranges. Although late Cenozoic normal faults are distributed relatively uniformly across this region, historical and instrumental seismicity is concentrated in two zones: the central Nevada seismic zone, which extends along the western margin of the province in eastern California and western Nevada, and the intermountain seismic zone, along the eastern edge of the province from southern Nevada across central Utah to southwestern Montana. The large (>M 7) earthquakes of 1872, 1915, and 1954 occurred within the former zone and the large events of 1959 and 1983 within the latter. Paleoseismic studies of normal faults in the Basin and Range Province suggest that many of the faults produce such big earthquakes only every few thousand years and that the current level of activity is abnormally high in the central Nevada seismic zone and abnormally low in the intermountain seismic belt. This possibility is of particular importance to Carson City and Salt Lake City, the capitals of Nevada and Utah, respectively, which sit on the edges of the province. 64. Examples include the 1987 Edgecomb earthquake (M 6.6) caused by failure of several normal faults within the volcanic arc of North Island, New Zealand, as well as dozens of historically important earthquakes in Greece and western Turkey that have occurred in the broad extensional back-arc setting of the Aegean Sea, for example, in the Bay of Corinth in 1861 and 1981, on the Pelopponese near Kalamata in 1981 and 1998, and probably an earthquake that destroyed Sparta in 464 B.C. (R. Armijo, H. Lyon-Caen, and D. Papanastassiou, A possible normal-fault rupture for the 464 BC Sparta earthquake, Nature, 351, 137-139, 1991). 65. An alternative explanation for this type of normal faulting involves behind-the-arc divergence associated with changes in the curvatures of the plate boundary thrust faults.
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66. R. Armijo, P. Tapponnier, J.L. Mercier, and T.-L. Han, Quaternary extension in southern Tibet: Field observations and tectonic implications, J. Geophys. Res., 91, 13,803-13,872, 1986. A combination of right-lateral and normal faulting in the eastern part of this extensional region resulted in the great Beng Co earthquake of 1951 and its large aftershock. 67. T. Camelbeeck and M. Meghraoui, Large earthquakes in northern Europe more likely than once thought, Eos, Trans. Am. Geophys. Union, 77, 405-409, 1996. 68. R.S. Yeats, K. Sieh, and C.R. Allen, The Geology of Earthquakes, Oxford University Press, Oxford, U.K., 568 pp., 1997. 69. These oceanographic surveys that have included high-resolution bathymetric mapping, seismicity studies using portable ocean-bottom seismographs, and visual investigations involving bottom photography and manned submersibles (e.g., P. Lonsdale, Structural geomorphology of the Eltanin fault system and adjacent transform faults of the Pacific-Antarctic plate boundary, Marine Geophys. Researches, 17, 105-143, 1994). 70. N.H. Woodcock, The role of strike-slip fault systems at plate boundaries, Phil. Trans. Roy. Soc. Lond., A317, 13-29, 1986. 71. The 56,000 year paleoseismic record of N. Porat, A.G. Wintle, R. Amit, and Y. Enzel (Late Quaternary earthquake chronology from luminescence dating of colluvial and alluvial deposits of the Arava Valley, Israel, Quaternary Res., 46, 107-117, 1996) is also discussed by G. Leonard, D.M. Steinberg, and N. Rabinowitz (An indication of time-dependent seismic behavior—An assessment of paleoseismic evidence from the Arava Fault, Israel, Bull. Seis. Soc. Am., 88, 767-776, 1998). Fragments of the seismic history of the Dead Sea transform are known from four millennia of recorded history and from archeological evidence (A. Ben-Menahem, Four thousand years of seismicity along the Dead Sea Rift, J. Geophys. Res., 96, 20,195-20,216, 1991). A major earthquake destroyed the ancient city of Jericho, on the northern edge of the Dead Sea, in the sixteenth century, B.C. This earthquake may have influenced the Old Testament writer who described the collapse of the walls of Jericho with the sounding of Joshua’s trumpet (Joshua 6:20). Careful analysis of historical accounts (N.N. Ambraseys and M. Barazangi, The 1759 earthquake in the Bekaa Valley; Implications for earthquake hazard assessment in the eastern Mediterranean region, J. Geophys. Res., 94, 4007-4013, 1989; N.N. Ambraseys and C.P. Melville, An analysis of the eastern Mediterranean earthquake of 20 May 1202, in Historical Seismograms and Earthquakes of the World, W.H.K. Lee, H. Meyers, and K. Shimazaki, eds., Academic Press, San Diego, Calif., pp. 181-200, 1988) suggest that the northern 350 kilometers of this system, in Lebanon and Syria, ruptured in a series of eight major destructive earthquakes during the past millennium. These occurred in three temporal clusters in the twelfth, fifteenth, and seventeenth centuries. 72. P. Tapponnier and P. Molnar, Active faulting and Cenozoic tectonics of the Tien Shan, Mongolia, and Baykal regions, J. Geophys. Res., 84, 3425-3456, 1979. 73. A.A. Barka, Slip distribution along the North Anatolian fault associated with the large earthquakes of 1939-1967, Bull. Seis. Soc. Am., 86, 1238-1254, 1996; R.S. Stein, A.A. Barka, and J.H. Dieterich, Progressive failure on the North Anatolian fault since 1939 by earthquake stress triggering, Geophys. J. Int., 128, 594-604, 1997. 74. K. Rajendran and C. Rajendran, Paleoseismological investigations in Runn of Kachch, India, the site of the large 1819 earthquake, in Summer School in Active Faulting and Paleoseismology, M. Meghraoui, ed., European Centre for Geodynamics and Seismology, Luxembourg, pp. 123-124, 1998. 75. L. Seeber, G. Ekström, S.K. Jain, C.V.R. Murty, N. Chandak, and J.G. Armbruster, The 1993 Killari earthquake in central India: A new fault in Mesozoic basalt flows? J. Geophys. Res., 101, 8543-8560, 1996. 76. A. Crone, M. Machette, and R. Bowman, Geologic Investigations of the 1988 Tennant Creek, Australia, Earthquakes: Implications for Paleoseismicity in Stable Continental Regions, U.S. Geological Survey Bulletin 2032A, Reston, Va., p. 51, 1992.
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77. M.N. Machette and J.R. Bowman, Episodic nature of earthquake activity in stable continental regions revealed by paleoseismicity studies of Australian and North American Quaternary faults, Austr. J. Earth Sci., 44, 203-214, 1997. 78. Proprietary software packages for loss estimation were first developed in the private sector. An example was the Early Post-Earthquake Damage Assessment Tool (EPEDAT) developed by EQE International, Inc., which applies intensity-damage relationships over specified zip codes. EPEDAT was developed for the California Office of Emergency Services and has been customized for five southern California counties (Los Angeles, Orange, Riverside, San Bernardino, and Ventura). HAZUS, developed by FEMA, uses PGA, PGV, Sa, and Sv inputs and building capacity-pushover analysis for census tracts. 79. HAZUS was developed by a consortium of university and private sector researchers and is maintained through NIBS. Its trial release in 1994 was followed by extensive efforts to collect input data for the software, to refine the algorithms for calculating damage, to educate state and local planning officials about the need for accurate information to support loss modeling, and to validate the methodology against well-characterized earthquakes. The final report on the methodology was delivered in fall 1997. The currently released version is HAZUS ®99, SR2, available at <http://www.fema.gov/hazus>, and further improvements are continuing. 80. See <www.hazus.org>. 81. The user should also have a good technical understanding of seismic hazards and the earthquake vulnerability of the modeled facilities to create realistic representations of the urban inventory and its fragilities, as well as the geologic hazards. The default data for some of these inputs, which a less skilled user might be tempted to employ, are crude and can lead to unrealistic loss estimates. 82. Western U.S. earthquakes resulting in loss of life were the 1987 Whittier Narrows (M 5.9, 8 deaths), 1989 Loma Prieta (M 7.0, 63 deaths), 1992 Landers (M 7.3, 1 death), and 1994 Northridge (M 6.7, 57 deaths). Other severe earthquakes with no loss of life were 1983 Coalinga (M 6.5), 1992 Petrolia (M 6.9), 1999 Hector Mine (M 7.1), and 2001 Nisqually (M 6.8). 83. A worldwide total of 160,000 is a lower bound from conservative USGS estimates for the 21 earthquakes resulting in more than 1000 deaths each that occurred from October 1983 to January 2001. See <http://neic.usgs.gov/neis/eqlists/eqsmajr.html>. 84. The BSSC (<http://www.bssconline.org>) is an independent, voluntary membership body representing a wide variety of building community interests. It was established in 1979 under the auspices of the NIBS to enhance public safety by providing a national forum to foster improved seismic safety provisions for use by the building community in the planning, design, construction, regulation, and utilization of buildings. The BSSC promotes the development and adoption of seismic safety provisions in building codes suitable for use throughout the United States. 85. The ATC (<http://www.atcouncil.org>) is a nonprofit, tax-exempt corporation established in 1971 through the efforts of the Structural Engineers Association of California. Its mission is to develop state-of-the-art, user-friendly engineering resources and applications for use in mitigating the effects of natural and other hazards on the built environment. ATC also identifies and encourages needed research and develops consensus opinions on structural engineering issues in a nonproprietary format. ATC is guided by a Board of Directors consisting of representatives appointed by the American Society of Civil Engineers, the Structural Engineers Association of California, the Western States Council of Structural Engineers Associations, and four at-large representatives concerned with the practice of structural engineering. Funding for ATC projects is obtained from government agencies and from the private sector in the form of tax-deductible contributions.
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86. The Earthquake Engineering Research Institute (<http://www.eeri.org>) was founded in 1949 as an outgrowth of the Advisory Committee on Engineering Seismology of the U.S. Coast and Geodetic Survey. A founding purpose of the institute was to encourage research in the field of earthquake engineering. It is a national, nonprofit, technical society of engineers, geoscientists, architects, planners, public officials, and social scientists with the mission of advancing the science and practice of earthquake engineering and the solution of national earthquake engineering problems to protect people and property from the effects of earthquakes. 87. The American Society of Civil Engineers (<http://www.asce.org/>) is a not-for-profit organization actively involved in developing seismic standards and codes. 88. J.E. Bevers, ed., Theme issue: Seismic isolation, Earthquake Spectra, 6, 161-430, 1999. 89. R.D. Hanson, ed., Theme issue: Passive energy dissipation, Earthquake Spectra, 9, 319-641, 1993; R.D. Hanson and T.T. Soong, Seismic Design with Supplemental Energy Dissipation Devices, Earthquake Engineering Research Institute, Monograph MNO-8, Oakland, Calif., 135 pp., 2001. 90. T.T. Soong and M.C. Constantinou, eds., Passive and Active Structural Vibration Control in Civil Engineering, Springer-Verlag, Vienna, 380 pp., 1994. 91. The quote is from Seismology Committee, Recommended Lateral Force Requirements— Commentary, Structural Engineers Association of California, Sacramento, 203 pp., 1990. 92. Guidelines for the Seismic Rehabilitation of Buildings (Building Seismic Safety Council and Applied Technology Council, Federal Emergency Management Agency Report FEMA-273, Washington, D.C., 400 pp., October, 1997). Important precursors were two reports published soon after the Northridge earthquake, Performance Based Seismic Engineering of Buildings (Vision 2000 Committee, J. Soulanges, ed., Structural Engineers Association of California, Sacramento, 2 vols., 115 pp., 1995) and Guidelines for the Seismic Rehabilitation of Buildings (Building Seismic Safety Council and Applied Technology Council, Federal Emergency Management Agency Report FEMA-273, Washington, D.C., 400 pp., October, 1997). 93. SAC Joint Venture, Recommended Seismic Design Guidelines for New Steel Moment-Frame Buildings, FEMA 350, U.S. Government Printing Office: Washington, D.C., 207 pp., 2000. 94. K. Olsen, Site amplification in the Los Angeles Basin from three-dimensional modeling of ground motion Bull. Seis. Soc. Am., 90, S77-S94, 2000. 95. A. Cornell and H. Krawinkler, Progress and challenges in seismic performance assessment, Pacific Earthquake Engineering Research Center, PEER Center News, 3, 1-3, 2000. 96. J. Milne, Earthquakes and Other Earth Movements, D. Appelton and Company, New York, p. 304, 1886. 97. National Research Council, Real-Time Earthquake Monitoring, National Academy Press, Washington, D.C., 52 pp., 1991. See also H. Kanamori, E. Hauksson, and T.H. Heaton, Real-time seismology and hazard mitigation, Nature, 390, 461-464, 1997. 98. J.M. Espinosa Aranada, A. Jiménez, G. Ibarrola, F. Alcantar, A. Aguilar, M. Inostroza, and S. Maldanado, Mexico City seismic alert system, Seis. Res. Lett., 66, 42-53, 1995. 99. W.H. Bakun, F.G. Fischer, E.G. Jensen, and J. VanSchaack, Early warning system for foreshocks, Bull. Seis. Soc. Am., 84, 359-365, 1994. 100. In California, warning systems were first used to alert rescue workers about aftershocks following the Loma Prieta earthquake. 101. The centerpiece of this upgrade is the new installation of the TriNet network of 170 broadband sensors and 700 strong-motion instruments in southern California. The operational goal is that the system will detect all events to M > 1.8 and that all of the data will be available in real time for hazard mitigation purposes.
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102. D.J. Wald, V. Quitoriano, T.H. Heaton, H. Kanamori, C.W. Scrivener, and B. Worden, Trinet “shakemaps”: Rapid generation of peak ground motion and intensity maps for earthquakes in Southern California, Earthquake Spectra, 15, 537-555, 1998. 103. Such capabilities could be of great value following an earthquake when communication and transportation are difficult. Examples include improved coordination of the response of firefighting and medical efforts, as well as routing and prioritizing the overload of telephone calls in the critical hours after an earthquake. 104. Office of Technology Assessment, Reducing Earthquake Losses, OTA-ETI-623, U.S. Government Printing Office, Washington, D.C., 162 pp., 1995. 105. One example of this holistic approach to hazard mitigation is the Natural Hazards Center of the University of Colorado, which is a national clearinghouse for information on natural hazards mitigation, with emphasis on social and political aspects (<www.colorado.edu/hazards>). 106. See <http://peer.berkeley.edu/lifelines/>. 107. See <http://www.cityofseattle.net/projectimpact/>. 108. The NSF Directorate of Engineering funds the Pacific Earthquake Engineering Research Center (<http://peer.berkeley.edu/>), the Multidisciplinary Center for Earthquake Engineering Research (<http://mceer.buffalo.edu/>), and the Mid-America Earthquake Center (<http://mae.ce.uiuc.edu/>). 109. NEES will provide real-time remote access to a complete set of testing and experimental facilities, making them widely available to earthquake engineers. The on-line network, or “collaboratory,” will furnish researchers across the country with shared-use access to advanced equipment, databases, and computer modeling and simulation tools (<http://www.eng.nsf.gov/nees/>). 110. The NSF and USGS sponsor the Southern California Earthquake Center, which involves 40 universities, government laboratories, and other public and private research organizations. The USGS also maintains major centers for earthquake research in Menlo Park, California, and Golden, Colorado.
Representative terms from entire chapter: