4
Observing the Active Earth: Current Technologies and the Role of the Disciplines

Not long ago, seismologists worked in rooms filled with drum recorders and big tables for hand-measuring seismograms. They now use digital monitoring systems that integrate high-performance seismometers, real-time communications, and automatic processing to produce high-quality information on seismic activity in near real time. Geodesists have replaced the theodolite and spirit level with space-based positioning and deformation imaging that can map crustal movements precisely and continuously, and they can hunt for slow, silent earthquakes with arrays of sensitive, stable strainmeters. Geologists have learned to decipher the subtle features of the rock record that mark prehistoric earthquakes, and they can date these events precisely enough to reconstruct the space-time behavior of entire fault systems. Laboratory and field scientists who study the microscopic processes of rock deformation are now formulating and calibrating the scaling laws that relate their reductionistic approach to the nonlinear dynamics of macroscopic faulting in the real Earth.

In each of these four domains—seismology, geodesy, geology, and rock mechanics—key technological innovations and conceptual breakthroughs were made within the last decade. The Global Seismic Network (GSN), initiated with the founding of the National Science Foundation (NSF)-sponsored Incorporated Research Institutions for Seismology (IRIS) in 1984, is reaching its design goal of 128 broadband, high-dynamic-range stations (as of December 2001, 126 stations had been installed and 122 were operational). The first continuously recording network of Global Positioning System (GPS) stations for measuring tectonic deformation



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4 Observing the Active Earth: Current Technologies and the Role of the Disciplines Not long ago, seismologists worked in rooms filled with drum recorders and big tables for hand-measuring seismograms. They now use digital monitoring systems that integrate high-performance seismometers, real-time communications, and automatic processing to produce high-quality information on seismic activity in near real time. Geodesists have replaced the theodolite and spirit level with space-based positioning and deformation imaging that can map crustal movements precisely and continuously, and they can hunt for slow, silent earthquakes with arrays of sensitive, stable strainmeters. Geologists have learned to decipher the subtle features of the rock record that mark prehistoric earthquakes, and they can date these events precisely enough to reconstruct the space-time behavior of entire fault systems. Laboratory and field scientists who study the microscopic processes of rock deformation are now formulating and calibrating the scaling laws that relate their reductionistic approach to the nonlinear dynamics of macroscopic faulting in the real Earth. In each of these four domains—seismology, geodesy, geology, and rock mechanics—key technological innovations and conceptual breakthroughs were made within the last decade. The Global Seismic Network (GSN), initiated with the founding of the National Science Foundation (NSF)-sponsored Incorporated Research Institutions for Seismology (IRIS) in 1984, is reaching its design goal of 128 broadband, high-dynamic-range stations (as of December 2001, 126 stations had been installed and 122 were operational). The first continuously recording network of Global Positioning System (GPS) stations for measuring tectonic deformation

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was installed in Japan in 1988 by the National Research Institute for Earth Science and Disaster Prevention (1), and the first image of earthquake faulting using interferometric synthetic aperture radar (InSAR) was constructed in 1992. Paleoseismologists produced a preliminary 1000-year history of major ruptures on the San Andreas fault in 1995 and discovered a prehistoric moment magnitude (M) 9 earthquake in the Cascadia subduction zone in 1996. The first three-dimensional simulations of dynamic fault ruptures using laboratory-derived, rate- and state-dependent friction equations were run in 1996. The unprecedented flow of new information opened by these advances is stimulating research on many fronts, from fault-system dynamics and earthquake forecasting to wavefield modeling and the prediction of strong ground motions. This chapter summarizes the state of the art in the main observational disciplines; it focuses on new technologies for observing the active Earth, and it highlights through a few examples the richness of the data sets now becoming available for basic and applied research. 4.1 SEISMOLOGY Seismology lies at the core of earthquake science because its main concern is the measurement and physical description of ground shaking. The central problem of seismology is the prediction of ground motions from knowledge of seismic-wave generation by faulting (the earthquake source) and the elastic medium through which the waves propagate (Earth structure). In order to do this calculation (forward problem), information must be extracted from seismograms to solve two coupled inverse problems: imaging the earthquake source, as represented by its space-time history of faulting, and imaging Earth structure, as represented by three-dimensional models of seismic-wave speeds and attenuation parameters. Because seismic signals can be recorded over such a broad range of frequencies—up to seven decades (2)—seismic signals can be used to observe earthquake processes on time scales from milliseconds to almost an hour, and they provide information about elastic structure at dimensions ranging from centimeters to the size of the Earth itself. Seismometry Seismic waves span a wide range of amplitude, as well as frequency. The ground motions in the vicinity of a large earthquake can have velocities greater than 1 meter per second and accelerations exceeding the pull of gravity (1g = 9.8 m/s2). The lower limit of seismic detection is typically eight orders of magnitude smaller, set by the level of the ambient ground

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noise (3). No single sensor has yet been developed that can faithfully record the violent displacements close to an earthquake and still be capable of detecting small events at the background noise level. For this reason, instruments historically have been divided into weak-motion and strong-motion seismometers. The former have been the principal sensors for studies of Earth structure and remote earthquakes by seismologists (4), while the latter have provided the principal seismological data to earthquake engineers. Technology is closing this gap. Modern force-feedback systems (5) can faithfully record ground motions from the lowest ambient noise at quiet sites to the largest earthquakes at teleseismic distances and achieve a bandwidth that extends from free oscillations with periods of tens of minutes to body waves with periods of tenths of seconds (Figure 4.1). Seismic Monitoring Systems Seismic monitoring systems comprise three basic elements: a network of seismometers that convert ground vibrations to electrical signals, communication devices that record and transmit the signals from the stations to a central facility, and analysis procedures that combine the signals from many stations to identify an event and estimate its location, size, and other characteristics. Monitoring systems are multiple-use facilities; they furnish information about earthquakes and nuclear explosions to operational agencies in near real time, and they also function as the basic data-gathering mechanisms for long-term research and education. With current technology, seismic networks of different types and spatial scales must be deployed to register the Earth’s seismicity over its complete geographic and magnitude range (Table 4.1, Figure 4.2). Since this coverage is typically overlapping, monitoring systems can be effectively organized into nested structures. Global Seismic Networks State-of-the-art seismic stations for global seismic networks comprise three-component sensors with high dynamic TABLE 4.1 Scales of Seismic Monitoring Type Typical Network Size Typical Station Spacing Detection Thresholda Global Global 1000 km 4.5 Regional 500 km 25-50 km 2.0 Local 10 km <1000 m –1.0 aMagnitude of smallest event with a high probability of detection; see examples in Figure 4.2.

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FIGURE 4.1 Plot of acceleration amplitude versus frequency, showing effective bandwidth and dynamic range for current broadband, high-frequency, and low-gain (strong-motion) digital seismometers, as well as for the older short-period and long-period analog seismometers of the World Wide Standardized Seismographic Network. Broadband instruments are capable of faithfully recording ground motions ranging from ambient noise at quiet sites (line labeled low Earth noise) to the peak accelerations generated by an M 9.5 earthquake at an epicentral distance of 3000 kilometers (upper line of stars). Low-gain seismometers are needed to record ground accelerations in the damage zones of large earthquakes, which can exceed 1g, and high-frequency seismometers are needed for periods less than about 0.1 second. SOURCE: R. Butler, IRIS.

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FIGURE 4.2 Number of earthquakes per year greater than a specified magnitude recorded by three networks of the types described in Table 4.1. Solid line shows the seismicity of the entire Earth from the global network of seismic stations cataloged by the International Seismological Centre during the decade 1990-1999. Dotted line is for events in southern California during the same interval recorded by the Southern California Seismic Network. Dashed line is for mining-induced seismicity recorded during 1997-1999 by a local network in the Elandrands gold mine, South Africa. Arrows show the approximate detection thresholds for the three networks, below which the sampling of seismicity is incomplete. All magnitudes are moment magnitudes. SOURCE: M. Boettcher, E. Richardson, and T.H. Jordan, University of Southern California. range (up to 140 decibels) and broadband response (0.0001–10 hertz). Since 1984, more than 300 such stations have been installed at permanent locations worldwide, as elements of global and regional networks (Figure 4.3). Close to half of these stations are part of the GSN, which has been constructed and operated under a cooperative agreement between the U.S. Geological Survey (USGS) and IRIS (6). The GSN is coordinated with other international networks through the Federation of Digital Seismographic Networks (FDSN), and the data are archived and made available

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FIGURE 4.3 Current global distribution of broadband high-performance seismic stations. The Global Seismic Network is part of the Federation of Digital Seismographic Networks (FDSN). SOURCE: IRIS-FDSN.

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on-line by the IRIS Data Management Center (DMC) in Seattle, Washington (7). Some stations still record on local magnetic or optical media that are shipped periodically to the DMC, but direct telemetry is being deployed as communication with remote sites becomes cheaper. At many locations with telephone access, the data can be retrieved via telephone dial-up or Internet connection (108 stations in 2001). The five-year goal is to have all stations on-line all the time. Achieving this goal, especially at remote sites, will depend in part on the cost of satellite communications. The GSN data acquired over the last 15 years have facilitated many advances in the study of global Earth structure and earthquake sources. Seismic tomography has provided dramatic images of subducting slabs, plume-like upwellings, and other features of the mantle convective flow responsible for plate-tectonic motions (Figure 4.4). The GSN data have also improved the plate-tectonic framework for understanding earthquake hazards through better earthquake locations and centroid moment tensor (CMT) solutions (Figure 4.5). Seismologists have used the broadband waveforms to elucidate the details of rupture processes during large earthquakes from a variety of tectonic settings, shedding new light on the geologic and dynamic factors that govern the configuration of seismogenic zones and how earthquakes start and stop. These successes have in no way diminished the need for continued monitoring. Discoveries based on data now being collected by the GSN will undoubtedly continue into the indefinite future. On the rapidly slipping plate boundaries, large earthquakes recur at intervals ranging from decades to centuries, while the recurrence times for significant intraplate events can extend to many millennia. With each passing year, GSN data will thus add new information to the evolving pattern of global seismicity by the direct observation of large, rare events and the delineation of low-level seismicity that may mark the eventual occurrence of such events. The densification of seismic sources through time will also improve tomographic mapping of features in the crust and mantle that control seismicity and may be indicative of the forces causing lithospheric faulting. Global seismological monitoring could be further enhanced by increasing the spatial resolution on land with permanent and temporary deployments of seismometers, expanding the coverage of global networks to the ocean floor, and upgrading the present networks as new technologies become available. However, sustained funding of the global networks will present a continuing challenge. In terms of annualized expenditures, the operation and maintenance of the GSN is projected to be comparable to its initial capitalization. Under current arrangements, the USGS shares a portion of the costs of GSN operations with the NSF. Stable support of the GSN from a federal agency that embraces the mission of global seismic monitoring is essential to the long-term health of earthquake science.

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FIGURE 4.4 Depth cross section showing the African plume as seen in four recent global tomographic models of S velocity: SAW23B16 (Megnin and Romanowicz, 2000), SB4L18 (Masters et al., 1999), S362D1 (Gu et al., 2001), and S20RTS (Ritsema et al., 1999). Although the models differ in detail, they are in agreement on the broad characteristics of this major upwelling. SOURCE: Y.J. Gu, A.M. Dziewonski, W.-J. Su, and G. Ekström, Models of the mantle shear velocity and discontinuities in the pattern of lateral heterogeneities, J. Geophys. Res., 106, 11,169-11,199, 2001; G. Masters, H. Bolton, and G. Laske, Joint seismic tomography for P and S velocities: How pervasive are chemical anomalies in the mantle? Eos, Trans. Am. Geophys. Union, 80, S14, 1999; C. Megnin and B. Romanowicz, The 3D shear velocity structure of the mantle from the inversion of body, surface and higher mode waveforms, Geophys. J. Int.,143, 709-728, 2000; J. Ritsema, H. van Heijst, and J. Woodhouse, Complex shear wave velocity structure imaged beneath Africa and Iceland, Science, 286, 1925-1928, 1999. SOURCE: B. Romanowicz and Y. Gung, University of California, Berkeley.

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FIGURE 4.5 25 years of CMT solutions (1976-2000) for the region surrounding Africa. The availability of broadband data has made it possible to describe global seismicity in terms not only of location and magnitude, but also of fault mechanism, thereby greatly enhancing our view of active tectonics. SOURCE: Harvard CMT group. Regional Seismic Networks Owing to their sparse station coverage, global networks do a poor job of detecting and locating events with magnitudes less than about 4.5 (Figure 4.2), and their sampling is too crude for investigating how waves are produced by fault ruptures, especially the near-fault radiation that generates the complex patterns of strong ground motions observed in large earthquakes. To deal with these problems, seismologists have densified station arrays in areas of high (or otherwise interesting) seismicity. Regional networks are collections of seismographic stations distributed over tens to hundreds of kilometers, usually as permanent facilities. The information supplied by regional networks services three overlapping but distinct communities: (1) scientists and engineers

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engaged in basic and applied research; (2) engineers, public officials, and other decision makers charged with the management of earthquake risk and emergency response; and (3) public safety officials, news media, and the general public. As information technology has transformed the regional networks into integrated monitoring systems, they have become centers for educating the general public about earthquake hazards, as well as key facilities for training graduate students in seismology (8). The short-period, high-gain instruments historically used in regional networks (9) brought seismicity patterns into much clearer focus (Figure 4.6), but the dynamic range of these instruments was too low to furnish useful recordings of large regional events. In the last decade, deployments of broadband, high-dynamic-range seismometers have begun to transform the regional networks into much more powerful tools for investigating the basic physics of the earthquake source, the detailed structure of the Earth’s crust and deep interior, and the patterns of potentially destructive ground motions. With these data, seismologists can now map the patterns of slip during earthquakes using seismic tomography, just as they map Earth structure. Images of fault ruptures during the more recent earthquakes in the Los Angeles, San Francisco, and Seattle regions have all been captured by high-performance networks (Figure 4.7). Long-term funding has been a persistent problem for regional network operators, and new investments in equipment are badly needed (10). In particular, the implementation of new broadband technologies in regional monitoring has been lagging in the United States, especially when compared to the investments made by other high-risk countries such as Japan (Box 4.1) and Taiwan. Two exceptions are the Berkeley Digital Seismic Network in northern California and Caltech’s TERRAscope Network in southern California. Both are equipped with a combination of three-component broadband seismometers and three-component strong-motion accelerometers; they have digital station processors and feed continuous data streams via real-time telemetry to central processing sites. Although these networks have developed independently, a major effort is under way, with some support from the State of California, to modernize the earthquake monitoring infrastructure throughout the region by integrating the regional networks into a California Integrated Seismic Network monitoring in the United States. Local Networks Networks have been deployed with seismometers distributed over a few tens of kilometers or less for specialized purposes such as seismic monitoring of critical facilities (e.g., dams and nuclear power plants) or localized source zones (e.g., volcanoes or geothermal reservoirs). Local networks are important instruments for the study of natural earthquake laboratories such as deep mines. Digital arrays of very

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FIGURE 4.6 Vertical along-strike cross sections of earthquakes located along a 30-kilometer section of the San Andreas fault. Circles are an estimate of event size. (a) Locations of earthquakes between May 1984 and December 1997 reported in the Northern California Seismic Network catalog. (b) Relocated earthquakes show a linear, nearly horizontal pattern of seismicity. SOURCE: A.M. Rubin, D. Gillard, and J.-L. Got, Streaks of microearthquakes along creeping faults, Nature, 400, 635-641, 1999. Reprinted by permission from Nature copyright 1999 Macmillan Publishers Ltd. high frequency sensors have been deployed in deep mines in Canada, Poland, and South Africa to monitor mine tremors and rock bursts induced by mining activities (11), and they have furnished unique, close-in observations of earthquakes as large as M 5 and at depths as great as 4 kilometers. Recent research has shown that in the deep gold mines of South Africa, mine tremors caused by friction-controlled slip on faults

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    nia: An alternative interpretation, Science, 210, 534-536, 1981; W.E. Strange, The impact of refraction correction on leveling interpretations in southern California, J. Geophys. Res., 86, 2809-2834, 1981; R.S. Stein, Role of elevation-dependent errors on the accuracy of geodetic leveling in the southern California uplift, 1953-1979, in Earthquake Prediction—An International Review, D.W. Simpson and P.G. Richards, eds., American Geophysical Union, Maurice Ewing Series, 4, Washington, D.C., pp. 441-456, 1981). 39.   Errors in leveling lines accumulate as a constant times the square root of the line length L. The constant of proportionality is typically 3 millimeters when L is expressed in kilometers, such that for L = 25 kilometers, the standard error is about 15 millimeters. Measurement precision with GPS is about 1 to 2 millimeters in horizontal position and 5 to 10 millimeters in vertical position, which can be achieved over short length scales for observation periods of a few hours and at length scales of hundreds of kilometers with day-long observations. Thus, several years of continuous GPS recording can yield steady-state velocity estimates with a precision of under 1 millimeter per year. The precision of such long-term measurements depends critically on stable, deeply anchored monuments. 40.   T.A. Herring, I.I. Shapiro, T.A. Clark, C. Ma, J.W. Ryan, B.R. Schupler, C.A. Knight, G. Lundqvist, D.B. Shaffer, N.R. Vandenberg, B.E. Corey, H.F. Hinteregger, A.E.E. Rogers, J.C. Webber, A.R. Whitney, G. Elgered, B.O. Ronnang, and J.L. Davis, Geodesy by radio interferometry: Evidence for contemporary plate motion, J. Geophys. Res., 91, 8341-8347, 1986; T.H. Jordan and J.B. Minster, Measuring crustal deformation in the American West, Sci. Am., 256, 48-58, 1988. 41.   The utility of GPS for a variety of military, commercial, and recreational purposes has reduced the price of navigation-quality receivers with resolutions of about 10 meters to a few hundred dollars. However, geodetic-quality receivers that operate off the GPS carrier phase are an order of magnitude more expensive, partly because they require better electronics and must process two frequencies to correct for signal delays caused by charged particles in the Earth’s ionosphere. Errors in GPS data come from errors in the reference frame, drift of the clocks onboard the satellites, refraction in the ionosphere and troposphere, multipath reflection of radio waves from the satellites, and so forth. These error sources can be included as terms in the basic equations to model GPS signals, and with sufficiently redundant data, the errors can be reduced dramatically. Large continuous networks are especially valuable for this purpose. In tectonic geodesy an additional error arises from nontectonic motions of the survey points caused by soil motions and fluid withdrawal in the region of the survey point. To reduce site instability, the Southern California Continuous GPS Network has developed an effective type of monument fixed at four points, each more than 10 meters below ground, by stainless steel rods welded to the surface monument. 42.   K. Feigl, D. Agnew, Y. Bock, D. Dong, A. Donnellan, B. Hager, T. Herring, D. Jackson, T. Jordan, R. King, S. Larsen, K Larson, M. Murray, Z. Shen, and F. Webb, Space geodetic measurement of crustal deformation in central and Southern California, 1984-1992, J. Geophys. Res., 98, 21,677-21,712, 1993. 43.   The International GPS Service for Geodynamics (IGS) integrates several forms of geodetic measurement to provide precise estimates of the satellite orbital parameters, the locations of selected tracking stations, and other important information needed in GPS processing. For more information, see <http://igscb.jpl.nasa.gov/>. 44.   Because of the need to integrate information into a global network, GPS data processing and interpretation depend critically on data sharing. The International GPS Services for Geodynamics, the Universities NAVSTAR Consortium, and the Southern California Earthquake Center, among many organizations, have made great progress in publishing GPS data freely over the Internet, contributing greatly to scientific progress in tectonic geodesy.

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45.   The first InSAR image of an earthquake displacement field was published by D. Massonnet, M. Rossi, C. Carmona, F. Adragna, G. Peltzer, K. Feigl, and T. Rabaute (The displacement field of the Landers earthquake mapped by radar interferometry, Nature, 364, 138-142, 1993), who used a series of radar images acquired by the European Remote Sensing (ERS) satellites to construct an interferogram of the 1992 Landers earthquake (M 7.3). 46.   G. Peltzer, P. Rosen, F. Rogez, and K. Hudnut, Postseismic rebound in fault stepovers caused by pore fluid flow, Science, 273, 1202-1204, 1996. 47.   The proposed ECHO mission would be carried out jointly between NASA, NSF, and the USGS to provide spatially continuous strain measurements over wide geographic areas. The design goals of the proposed InSAR mission are dense spatial (100 meters) and temporal (every eight days) coverage of the entire North American-Pacific plate boundary with vector solutions accurate to 2 millimeters on spatial scales of 50 kilometers over all terrain types, which exceeds the capabilities of existing and planned international SAR missions. Spatially continuous, but intermittent, InSAR images complement continuous GPS point measurements and will therefore contribute to the EarthScope science objectives. 48.   Stable aseismic slip was discovered at the Cienega Winery, which straddles the San Andreas fault south of Hollister, California (K.V. Steinbrugge, E.G. Zacher, D. Tocher, C.A. Whitten, and C.N. Clair, Creep on the San Andreas fault [California]—Analysis of geodetic measurements along the San Andreas fault, Bull. Seis. Soc. Am., 50, 396-404, 1960). The walls of the winery building have been progressively offset at a rate of 11 millimeters per year since it was built in 1948. This “creeping section” of the San Andreas extends 160 kilometers from San Juan Bautista to Parkfield, California. 49.   Near-fault tectonic geodesy is reviewed by A.G. Sylvester in National Research Council, Active Tectonics, National Academy Press, Washington, D.C., pp. 164-180, 1986. 50.   C.R. Allen, M. Wyss, J.N. Brune, A. Grantz, and R.E. Wallace, Displacements on the Imperial, Superstition Hills, and San Andreas faults triggered by the Borrego Mountain earthquake, U.S. Geological Survey Professional Paper 787, Reston, Va., pp. 87-104, 1972; S.S. Schulz, G. Mavco, R.O. Burford, and W.D. Smith, Long-term fault creep observations in central California, J. Geophys. Res., 87, 6977-6982, 1982; R.O. Burford, The response of creeping parts of the San Andreas fault to earthquakes on nearby faults; Two examples, Pure Appl. Geophys., 126, 499-529, 1988; C.H. Thurber, Creep events preceding small to moderate earthquakes on the San Andreas fault, Nature, 380, 425-428, 1996. 51.   D.C. Agnew, Strainmeters and tiltmeters, Rev. Geophys., 24, 579-624, 1986. 52.   I.S. Sacks, S. Suyehiro, A.T. Linde, and J.A. Snoke, Slow earthquakes and stress redistribution, Nature, 275, 599-602, 1978; A.T. Linde, S. Suyehiro, I. Miura, I.S. Sacks, and A. Takagi, Episodic aseismic earthquake precursors, Nature, 334, 513-515, 1988; M.T. Gladwin, High-precision multicomponent borehole deformation monitoring, Rev. Sci. Instr., 55, 2011-2016, 1984. 53.   PBO Steering Committee, The Plate Boundary Observatory: Creating a four-dimensional image of the deformation of western North America, White paper providing the scientific rationale and deployment strategy for a Plate Boundary Observatory based on a workshop held October 3-5, 1999. Available at <http://www.earthscope.org>. 54.   The Southern California Earthquake Center employed this strategy in its 1995 earthquake hazard estimate. See Working Group on California Earthquake Probabilities, Seismic hazards in southern California: Probable earthquake 1994 to 2024, Bull. Seis. Soc. Am., 85, 379-439, 1995. In combination with geologic estimates of fault slip rate, they used GPS estimates from K. Feigl, D. Agnew, Y. Bock, D. Dong, A. Donnellan, B. Hager, T. Herring, D. Jackson, T. Jordan, R. King, S. Larsen, K Larson, M. Murray, Z. Shen, and F. Webb, Space geodetic measurement of crustal deformation in Central and Southern California, 1984-1992, J. Geophys. Res., 98, 21,677-21,712, 1993.

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55.   See, for example, E. Field, D. Jackson, and J.F. Dolan, A new look at earthquake occurrence in southern California: No deficit or huge earthquake required, Bull. Seis. Soc. Am., 89, 559-578, 1999. 56.   Harris and Segall (R. Harris and P. Segall, Detection of a locked zone at depth in the Parkfield, California segment of the San Andreas fault, J. Geophys. Res., 92, 7945-7962, 1987) identified a locked patch apparently surrounded by a creeping region. They identified the locked patch with the rupture zone of the 1966 Parkfield earthquake, suggesting that the geodetic data provided evidence for a future earthquake on the same rupture surface. Sung and Jackson (L. Sung and D. Jackson, Geodetic evidence of the seismic potential at Parkfield, California, Geophys. Res. Lett., 15, 820-823, 1988) showed that a smooth transition from creeping to locked, without the asperity, also fit the geodetic data within their accuracy. 57.   For example, models of the 1992 Landers, California, earthquake (K.W. Hudnut, Y. Bock, M. Cline, P. Fang, Y. Feng, J. Freymueller, X. Ge, W.K. Gross, D. Jackson, M. Kim, N.E. King, J. Langbein, S.C. Larsen, M. Lisowski, Z.K. Shen, J. Svarc, and J. Zhang, Coseismic displacements of the 1992 Landers earthquake sequence, Bull. Seis. Soc. Am., 84, 625-645, 1994) confirm the strong spatial variability and location of slip derived from modeling of the strong-motion data. In the case of the 1989 Loma Prieta earthquake, a geodetically derived slip model (T. Arnadottir and P. Segall, The 1989 Loma Prieta earthquake imaged from inversion of geodetic data,J. Geophys. Res., 99, 21,835-21,855, 1994) supported seismic models that found a variation of rate with distance along the fault in that earthquake (G.C. Beroza and H. Krawinkler, Near-source modeling of the Loma Prieta earthquake; Evidence for heterogeneous slip and implications for earthquake hazard, Bull. Seis. Soc. Am., 81, 1603-1621, 1991). 58.   P. Segall and M. Matthews, Time dependent inversion of geodetic data, J. Geophys. Res., 102, 22,391-22,409, 1997. 59.   T. Arnadottir and P. Segall, The 1989 Loma Prieta earthquake imaged from inversion of geodetic data,J. Geophys. Res., 99, 21,835-21,855, 1994. 60.   G.C. Beroza and H. Krawinkler, Near-source modeling of the Loma Prieta earthquake; Evidence for heterogeneous slip and implications for earthquake hazard, Bull. Seis. Soc. Am., 81, 1603-1621, 1991. 61.   J. Freymueller, N.E. King, and P. Segall, The co-seismic slip distribution of the Landers earthquake, Bull. Seis. Soc. Am., 84, 646-659, 1994. 62.   D. Massonnet, M. Rossi, C. Carmona, F. Adragna, G. Peltzer, K. Feigl, and T. Rabaute, The displacement field of the Landers earthquake mapped by radar interferometry, Nature, 364, 138-142, 1993; G. Peltzer, P. Rosen, F. Rogez, and K. Hudnut, Postseismic rebound in fault step-overs caused by pore fluid flow, Science, 273, 1202-1204, 1996. 63.   T. Arnadottir and P. Segall, The 1989 Loma Prieta earthquake imaged from inversion of geodetic data,J. Geophys. Res., 99, 21,835-21,855, 1994; K.W. Hudnut, Y. Bock, M. Cline, P. Fang, Y. Feng, J. Freymueller, X. Ge, W.K. Gross, D. Jackson, M. Kim, N.E. King, J. Langbein, S.C. Larsen, M. Lisowski, Z.K. Shen, J. Svarc, and J. Zhang, Co-seismic displacements of the 1992 Landers earthquake sequence, Bull. Seis. Soc. Am., 84, 625-645, 1994; Z.-K. Shen, B.X. Ge, D.D. Jackson, D. Potter, M. Cline, and L.Y. Sung, Northridge earthquake rupture models based on the Global Positioning System measurements, Bull. Seis. Soc. Am., 86, S37-S48, 1996. 64.   I.S. Sacks, A.T. Linde, J.A. Snoke, and S. Suyehiro, A slow earthquake sequence following the Izu-Oshima earthquake of 1978, in Earthquake Prediction—An International Review, D.W. Simpson and P.G. Richards, eds., American Geophysical Union, Maurice Ewing Series, 4, Washington, D.C., pp. 617-628, 1981. 65.   H. Dragert, K. Wang, and T.S. James, A silent slip event on the deeper Cascadia subduction interface, Science, 292, 1525-1528, 2001.

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66.   M.M. Miller, T. Melbourne, D.J. Johnson, and W.Q. Sumner, Periodic slow earthquakes from the Cascadia subduction zone, Science, 295, 2432, 2002. 67.   W. Thatcher, Present-day crustal movements and the mechanics of cyclic deformation, in San Andreas Fault System, California, R.E. Wallace, ed., U.S. Geological Survey Professional Paper 1515, Washington, D.C., pp. 189-205, 1990. 68.   D.D. Jackson, Z.-K. Shen, D. Potter, X.-B. Ge, and L. Sung, Southern California deformation, Science, 277, 1621-1622, 1997. 69.   Two limnographs on different sides of Rapel Lake showed water-level differences, implying tilting, in response to the 1985 earthquake in Chile. The tilting had a characteristic time of 10 months. A similar signal was also seen on tide gauge records of sea level recorded at a distance of 100 kilometers. There was no co-seismic change in lake level, so the postseismic and co-seismic deformations were generated differently. Based on these observations Barrientos concluded that the postseismic slip occurred updip from the co-seismic slip. See S.E. Barrientos, Dual seismogenic behavior; The 1985 central Chile earthquake, Geophys. Res. Lett., 22, 3541-3544, 1995. 70.   For the Loma Prieta earthquake, fault-normal compression after the earthquake exceeded that from the earthquake itself. Savage and others explain this observation with a model of postseismic fault-zone collapse with a time constant of 1.3 years (J.C. Savage, M. Lisowski, and J.L. Svarc, Postseismic deformation following the 1989 (M = 7.1) Loma Prieta, California, earthquake, J. Geophys. Res., 99, 13,757-13,765, 1994). Bürgmann and others model the same observations with a combination of continuing slip on the mainshock plane and aseismic slip on reverse faults located to the northeast of the mainshock fault plane (R. Burgmann, P. Segall, M. Lisowski, and J.P. Svarc, Postseismic strain following the 1989 Loma Prieta earthquake from GPS and leveling measurements, J. Geophys. Res., 102, 4933-4955, 1997). 71.   Z.-K. Shen, D.D. Jackson, Y. Feng, M. Cline, M. Kim, P. Fang, and Y. Bock, Postseismic deformation following the Landers earthquake, California, 28 June 1992, Bull. Seis. Soc. Am.,84, 780-791, 1994. These authors found a postseismic signal with a decay time of about 1 month following the Landers earthquake, with slip in their model concentrated on the southern half of the mainshock rupture zone. They also inferred slip at depth below the Banning segment of the San Andreas fault, which could have serious consequences for possible earthquakes on the San Andreas. Savage and Svarc inferred that postseismic slip was concentrated on the northern half of the mainshock rupture zone and that postseismic deformation is still ongoing (J.C. Savage and J.L. Svarc, Postseismic deformation associated with the 1992 Mw = 7.3 Landers earthquake, Southern California, J. Geophys. Res., 102, 7565-7577, 1997); they also found evidence for fault-zone collapse in the Landers postseismic zone. 72.   F.K. Wyatt, D.C. Agnew, and M.T. Gladwin, Continuous measurements of crustal deformation for the 1992 Landers earthquake sequence, Bull. Seis. Soc. Am., 84, 768-779, 1994. 73.   D. Massonnet, W. Thatcher, and H. Vadon, Detection of postseismic fault-zone collapse following the Landers earthquake, Nature, 382, 612-616, 1996. Interferograms spanning intervals after the Landers mainshock show clear postseismic signals consistent with ongoing slip in the rupture zone as well as motions concentrated within fault offsets. The latter have been attributed to pressure-driven fluid flow into dilatant regions (G. Peltzer, P. Rosen, F. Rogez, and K. Hudnut, Postseismic rebound in fault step-overs caused by pore fluid flow, Science, 273, 1202-1204, 1996). 74.   P.A. Rydelek and I.S. Sacks, Asthenospheric viscosity and stress diffusion; A mechanism to explain correlated earthquakes and surface deformations in NE Japan,Geophys. J. Int., 100, 39-58, 1990; I.S. Sacks and F.F. Pollitz, Analysis of Postseismic Crustal Motions in California, Carnegie Institution, Washington, D.C., 5 pp. + 7 plates, 1992; F.F. Pollitz and I.S.

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    Sacks, Modeling of postseismic relaxation following the great 1857 earthquake, Southern California, Bull. Seis. Soc. Am.,82, 454-480, 1992; F.F. Pollitz and I.S. Sacks, Consequences of stress changes following the 1891 Nobi earthquake, Japan, Bull. Seis. Soc. Am., 85, 796-807, 1995; F. Press and C. Allen, Patterns of seismic release in the southern California region,J. Geophys. Res., 100, 6421-6430, 1995. 75.   The best bounds on the sizes of strain precursors are from recent earthquakes in California and Japan. See M.L.S. Johnson, A.T. Linde, and M.T. Gladwin (Near-field high resolution strain measurements prior to the October 18, 1989 Loma Prieta Ms 7.1 earthquake, Geophys. Res. Lett., 17, 1777-1780, 1990) for the Loma Prieta earthquake, and F.K. Wyatt, D.C. Agnew, and M. Gladwin (Continuous measurements of crustal deformation for the 1992 Landers earthquake sequence, Bull. Seis. Soc. Am., 84, 768-779, 1994) for the Landers earthquake. 76.   Developments in this field are summarized in the textbook The Geology of Earthquakes, by R.S. Yeats, K. Sieh, and C.R. Allen (Oxford University Press, Oxford, U.K., 568 pp., 1997). 77.   National Research Council, Active Tectonics, National Academy Press, Washington, D.C., 266 pp., 1986. 78.   The SRTM mission and results are described at <http://jpl.nasa.gov/srtm>. 79.   An example is the Airborne Topographic Mapper, mounted on an Otter aircraft and used in the U.S. Topographic Change Mapping Project, a joint venture among NOAA, NASA, and USGS; see <http://www.csc.noaa.gov/crs/tcm/>. 80.   In deep water (3 kilometers), oceanographic swath-mapping systems yield bathymetric maps with a horizontal resolution of about 60 meters and a vertical precision of a few meters. The resolution and precision improve more or less linearly with decreasing water depth. 81.   For example, N.N. Ambraseys and C.P. Melville, A History of Persian Earthquakes, Cambridge University Press, Cambridge, U.K., 219 pp., 1982; Y. Sugiyama, Neotectonics of southwest Japan due to the right-oblique subduction of the Philippine Sea plate, Geof. Int., 33, 53-76, 1994; D.C. Agnew and K. Sieh, A documentary study of the felt effects of the great California earthquake of 1857, Bull. Seis. Soc. Am., 68, 1717-1729, 1978; 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. 82.   See R. Page, Dating episodes of faulting from tree rings: Effects of the 1958 rupture of the Fairweather fault on tree growth, Geol. Soc. Am. Bull., 81, 3085-3094, 1970; V.C. LaMarche, Jr. and R.E. Wallace, Evaluation of effects of trees on past movements on the San Andreas fault, northern California, Geol. Soc. Am. Bull., 83, 2665-2676, 1972; G.C. Jacoby, P.R. Sheppard, and K.E. Sieh, Irregular recurrence of large earthquakes along the San Andreas fault; Evidence from trees, Science, 241, 196-199, 1988; 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. 83.   The age range, resolution, and applications of different techniques for dating surficial materials are described in J.S. Noller, J.M. Sowers, and W.R. Lettis, Quaternary Geochronology: Methods and Applications, American Geophysical Union, Washington, D.C., 582 pp., 2000. 84.   J.A. Spotila, K.A. Farley, and K. Sieh, Uplift and erosion of the San Bernardino Mountains associated with transpression along the San Andreas fault, California, as constrained by radiogenic helium thermochronometry, Tectonics,17, 360-378, 1998; C.R. Bacon, M.A. Lanphere, and D.E. Champion, Late Quaternary slip rate and seismic hazards of the West Klamath Lake fault zone near Crater Lake, Oregon Cascades, Geology, 27, 43-46, 1999; J. Lee, C.M. Rubin, and A. Calvert, Quaternary faulting history along the Deep Springs fault, California, Geol. Soc. Am. Bull., 113, 855-869, 2001; A.K. Jain, D. Kumar; S. Singh, A.

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    Kumar, and N. Lal, Timing, quantification and tectonic modelling of Pliocene-Quaternary movements in the NW Himalaya; Evidence from fission track dating, Earth Planet. Sci. Lett., 179, 437-451, 2000; B.J. Szabo and J.N. Rosholt, Uranium-series nuclides in the Golden fault, Colorado, U.S.A.: Dating latest fault displacement and measuring recent uptake of radionuclides by fault-zone materials, Appl. Geochem., 4, 177-182, 1989. 85.   For example, K.E. Sieh, Prehistoric large earthquakes produced by slip on the San Andreas fault at Pallet Creek, California, J. Geophys. Res., 83, 3907-3939, 1978; K. Sieh, M. Stuiver, and D. Brillinger, A more precise chronology of earthquakes produced by the San Andreas fault in southern California, J. Geophys. Res., 94, 603-623, 1989. See Box 4.5. 86.   E. Bard, B. Hamelin, R.G. Fairbanks, and A. Zindler, Calibration of the 14C timescale over the past 30,000 years using mass spectrometric U-Th ages from Barbados corals, Nature, 345, 405-410, 1990. 87.   F.W. Taylor, C. Frohlich, J. Lecolle, and M. Strecker, Analysis of partially emerged corals and reef terraces in the central Vanuatu arc: Comparison of contemporary coseismic and nonseismic with Quaternary vertical movements, J. Geophys. Res., 92, 4905-4933, 1987; J. Zachariasen, K. Sieh, F. Taylor, R.L. Edwards, and W.S. Hantoro, Submergence and uplift associated with the giant 1833 Sumatran subduction earthquake: Evidence from coral microatolls, J. Geophys. Res., 104, 895-919, 1999. 88.   L.A. Perg, R.S. Anderson, and R.C. Finkel, Use of a new 10Be and 26Al inventory method to date marine terraces, Santa Cruz, California, USA, Geology, 29, 879-882, 2001; J. Van der Woerd, F.J. Ryerson, P. Tapponnier, Y. Gaudemer, R. Finkel, A.S. Meriaux, M. Caffee, G. Zhao, and Q. He, Holocene left-slip rate determined by cosmogenic surface dating on the Xidatan segment of the Kunlun fault (Qinghai, China), Geology, 26, 695-698, 1998; L.C. Benedetti, R.C. Finkel, G.C.P. King, R. Armijo, D. Papanastassiou, F.J. Ryerson, F. Flerit, and D. Farber, Earthquake time-slip history of the Kaparelli fault (Greece) from in situ chlorine-36 cosmogenic dating, EOS Trans. Am. Geophys. Union, 82, F931, 2001. 89.   Examples include neotectonic maps of Japan (Y. Kinugasa, E. Tsukada, and H. Yamazaki, Neotectonic map of Japan, Geological Atlas of Japan, 2nd ed., Asakura Publishing Company, Ltd., Tokyo, sheet 5, 1992), Turkey (F. Saroglu, O. Emre, and I. Kuscu, The east Anatolian fault zone of Turkey, Annales Tectonicae, Suppl. 6 (Special Issue), 99-125, 1992), Alaska (G. Plafker, L.M. Gilpin, and J.C. Lahr, Neotectonic map of Alaska, in The Geology of Alaska, G.B. Plafker and H. Berg, eds., Decade of North American Geology, G-1, Geological Society of America, Boulder, Colo., pl. 12 (map), 1994), southern Tibet (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), and Sumatra (K. Sieh and D. Natawidjaja, Neotectonics of the Sumatran fault, J. Geophys. Res., 105, 28,295-28,336, 2000). 90.   The World Map of Major Active Faults being compiled under Project II-2 of the International Lithosphere Program is a step in this direction; see<http://www.gfzpotsdam.de/pb4/ilp96/projects.htm>. 91.   L.A. Reinen, J.D. Weeks, and T.E. Tullis, The frictional behavior of serpentinite: Implications for aseismic creep on shallow crustal faults, Geophys. Res. Lett., 18, 1921-1924, 1991; The frictional behavior of lizardite and antigorite serpentinites: Experiments, constitutive models, and implications for natural faults, Pure Appl. Geophys., 143, 317-358, 1994. 92.   J. Van der Woerd, F.J. Ryerson, P. Tapponnier, A.-S. Meriaux, Y. Gaudemer, B. Meyer, R.C. Finkel, M.W. Caffee, Z. Guoguang, and X. Zhiqin, Uniform slip-rate across the Kunlun fault: Implications for seismic behavior and large-scale tectonics, Geophys. Res. Lett., 27, 2353-2356, 2000. 93.   J.-C. Lee, Y.-G. Chen, K. Sieh, K. Mueller, W.-S. Chen, H.-T. Chu, Y.-C. Chan, C. Rubin, and R. Yeats, A vertical exposure of the 1999 surface rupture of the Chelungpu fault at Wufeng, western Taiwan: Structural and paleoseismic implications for an active thrust fault, Bull. Seis. Soc. Am., 91, 914-929, 2001.

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94.   R.S. Yeats, Large-scale Quaternary detachments in Ventura basin, southern California, J. Geophys. Res., 88, 569-583, 1983. 95.   Developing a unified structural representation for southern California has been set as a high-priority goal of the Southern California Earthquake Center; see Southern California Earthquake Center, Science Plan for 2002-2007, University of Southern Calif., 9 pp., 2001, available at <http://www.scec.org/aboutSCEC/documents/science.plan.2002/>. 96.   This definition of paleoseismology is offered in the historical overview by R.S. Yeats and C.S. Prentice, Introduction to special session: Paleoseismology, J. Geophys. Res., 101, 5847-5853, 1996. A survey of the subject is given by J.P. McCalpin, ed., Paleoseismology, International Geophysics Series 62, Academic Press, San Diego, Calif., 588 pp., 1996. 97.   D.C. Agnew and K. Sieh, A documentary study of the felt effects of the great California earthquake of 1857, Bull. Seis. Soc. Am., 68, 1717-1729, 1978. Geologic features corresponding to individual earthquake offsets on the San Andreas, including the 1857 event, were first recognized by R.E. Wallace (Notes on stream channels offset by the San Andreas fault, southern Coast Ranges, California, in Proceedings of a Conference on Geological Problems of the San Andreas Fault System, W.R. Dickinson and A. Grantz, eds., Stanford University Publications in Geological Science 11, Stanford, Calif., pp. 6-21, 1968). 98.   K. Sieh, M. Stuvier, and D. Brillinger, A more precise chronology of earthquakes produced by the San Andreas fault in southern California, J. Geophys. Res., 94, 603-623, 1989. 99.   See K.R. Berryman, S. Beanland, A. Cooper, H. Cutten, R. Norris, and P. Wood, The Alpine fault, New Zealand: Variation in Quaternary structural style and geomorphic expression, Annales Tectonicae, Suppl. 6 (Special Issue), 126-163, 1992; K. Sieh, A review of geological evidence for recurrence times of large earthquakes, in Earthquake Prediction—An International Review, D. Simpson and P. Richards, eds., American Geophysical Union, Maurice Ewing Series 4, Washington, D.C., pp. 181-207, 1981; 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; Q.-D. Deng and P.-Z. Zhang, Research on the geometry of shear fracture zones, J. Geophys. Res., 89, 5699-5710, 1984. 100.   R.E. Wallace, Profiles and ages of young fault scarps, north-central Nevada, Geol. Soc. Am. Bull., 88, 1267-1281, 1977. 101.   K. Mueller, J. Champion, M. Guccione, and K. Kelson, Fault slip rates in the modern New Madrid seismic zone, Science, 286, 1135-1138, 1999. 102.   J. Clague, Evidence for large earthquakes at the Cascadia subduction zone, Rev. Geophys., 35, 439-460, 1997. 103.   F.W. Taylor, C. Frohlich, J. Lecolle, and M. Strecker, Analysis of partially emerged corals and reef terraces in the central Vanuatu arc: Comparison of contemporary coseismic and nonseismic with Quaternary vertical movements, J. Geophys. Res., 92, 4905-4933, 1987; R.L. Edwards, F.W. Taylor, and G.J. Wasserburg, Dating earthquakes with high-precision thorium-230 ages of very young corals, Earth Planet. Sci. Lett., 90, 371-381, 1988. 104.   J. Zachariasen, K. Sieh, F. Taylor, R.L. Edwards, and W.S. Hantoro, Submergence and uplift associated with the giant 1833 Sumatran subduction earthquake: Evidence from coral microatolls, J. Geophys. Res., 104, 895-919, 1999; K. Sieh, S. Ward, D. Natawidjaja, and B. Suwargadi, Crustal deformation at the Sumatran subduction zone revealed by coral rings, Geophys. Res. Lett., 26, 3141-3144, 1999. 105.   See R.S. Yeats, K. Sieh, and C.R. Allen, The Geology of Earthquakes, Oxford University Press, Oxford, U.K., 568 pp., 1997, for a more extensive enumeration and discussion. 106.   W.B. Bull, J. King, F. Kong, T. Moutoux, and W.M. Phillips, Lichen dating of coseismic landslide hazards in alpine mountains, Geomorph., 10, 253-264, 1994. 107.   J. Adams, Paleoseismicity of the Cascadia subduction zone: Evidence from turbidites off the Oregon-Washington margin, Tectonics, 9, 569-583, 1990.

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108.   P.J. Munson, S.F. Obermeier, C. Munson, and E.R. Hajic, Liquefaction evidence for Holocene and latest Pleistocene seismicity in the southern halves of Indiana and Illinois: A preliminary overview, Seis. Res. Lett., 68, 521-536, 1997. 109.   S.F. Obermeier, G.S. Gohn, R.F. Weems, R.L. Gelinas, and M. Rubin, Geologic evidence for recurrent moderate to large earthquakes near Charleston, South Carolina, Science, 227, 408-411, 1985. 110.   S. Marco, M. Stein, A. Agnon, and H. Ron, Long-term earthquake clustering: A 50,000-year paleoseismic record in the Dead Sea Graben, J. Geophys. Res., 101, 6179-6191, 1996. 111.   Most laboratory experiments in triaxial, direct-shear, and rotary-shear machines that can attain high pressures, temperatures, and fluid pressures involve rock samples with maximum dimensions of a few centimeters. The USGS has conducted stick-slip experiments at room conditions on a biaxially loaded granite sample of dimension 1.5 × 1.5 × 0.4 cubic meters; the sample was saw-cut on the diagonal and loaded by jacks at the edges, allowing both shear and normal stresses to be varied on a precut fault area of 2 × 0.4 square meters (J.H. Dieterich, N.G.W. Cook, and H.C. Heard, Potential for geophysical experiments in large scale tests, Geophys. Res. Lett., 8, 653-656, 1981; D.A. Lockner and P.G. Okubo, Measurements of frictional heating in granite, J. Geophys. Res., 88, 4313-4320, 1983). Slip events were observed with seismic moments as large as 3 × 106 newton-meters; this corresponds to a moment magnitude of about –1.7, which overlaps with the sizes of microearthquakes that have been recorded by seismic sensors in deep mines. 112.   M.S. Paterson, Experimental Rock Deformation—The Brittle Field, Springer-Verlag, Berlin, 254 pp., 1978. The development of double-direct-shear, rotary-shear, and other types of testing machines was discussed by T.E. Tullis and J.D. Weeks (Constitutive behavior and stability of frictional sliding of granite, Pure Appl. Geophys., 124, 383-414, 1986). 113.   The micromechanics of rate-state friction has been discussed by T.E. Tullis (Rock friction constitutive behavior from laboratory experiments and its implications for an earthquake prediction field monitoring program, Pure Appl. Geophys., 126, 555-588, 1988) and C.G. Sammis and S.J. Place (The micromechanics of friction in a granular layer, Pure Appl. Geophys., 142, 777-794, 1994). Recent advances in atomic-scale tribology (the study of friction, lubrication, and wear) are described by G. Hähner and N. Spencer (Rubbing and scrubbing, Physics Today, 51, 22-27, 1998). 114.   See the review by C. Marone, Laboratory-derived friction laws and their application to seismic faulting, Ann. Revs. Earth Planet. Sci., 26, 643-696, 1998. 115.   P. Okubo, Dynamic rupture modeling with laboratory-derived constitutive relations, J. Geophys. Res., 94, 12,321-12,335, 1989; A. Cochard and R. Madariaga, Dynamic faulting under rate-dependent friction, Pure Appl. Geophys., 142, 419-445, 1994; N. Lapusta, J.R. Rice, Y. Ben-Zion, and G. Zheng, Elastodynamic analysis for slow tectonic loading with spontaneous rupture episodes on faults with rate and state dependent friction, J. Geophys. Res., 105, 23,765-23,789, 2000. 116.   J.H. Dieterich, A constitutive law for rate of earthquake production and its application to earthquake clustering, J. Geophys. Res., 99, 2601-2618, 1994; J.H. Dieterich, V. Cayol, and P. Okubo, The use of earthquake rate changes as a stress meter at Kilauea volcano, Nature, 408, 457-460, 2000. 117.   K. Mair and C. Marone, Friction from simulated fault gouge at a wide range of velocities, J. Geophys. Res., 104, 28,888-28,894, 1999. As first hypothesized by D.J. Andrews (Rupture propagation with finite stress in antiplane strain, J. Geophys. Res., 81, 3579-3587, 1976), the critical slip distance for fault zones of finite width W may obey Dc = ?cW, where ?c is a “critical strain.” Experiments by C. Marone and B. Kilgore (Scaling of the critical slip distance for seismic faulting with shear strain in fault zones, Nature, 362, 618-621, 1993) gave ?c = 10–2, provided W was interpreted to be the width of the “active gouge” containing shear bands actually involved in the slip.

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118.   J.H. Dieterich and B. Kilgore, Implications of fault constitutive properties for earthquake prediction, Proc. Natl. Acad. Sci., 93, 3787-3794, 1996. 119.   B. Evans, J.T. Frederich, and T.F. Wong, The brittle-ductile transition in rocks: Recent experimental and theoretical progress, in The Brittle-Ductile Transition in Rocks: The Heard Volume, A.G. Duba, W. Durham, J. Handin, and H. Wang, eds., American Geophysical Union, Washington, D.C., pp. 1-20, 1990. 120.   N.L. Carter and S.H. Kirby, Transient creep and semibrittle behavior of crystalline rocks, Pure Appl. Geophys., 116, 807-839, 1978; C.H. Scholz, The Mechanics of Earthquakes and Faulting, Cambridge University Press, Cambridge, U.K., 439 pp., 1990. 121.   D.L. Kohlstedt, B. Evans, and S.J. Mackwell, Strength of the lithosphere: Constraints imposed by laboratory experiments, J. Geophys. Res., 100, 17,587-17,602, 1995. 122.   B.R. Lawn and T.R. Wilshaw, Fracture of Brittle Solids, Cambridge University Press, Cambridge, U.K., 204 pp., 1975. 123.   C.H. Scholz, Mechanics of faults, Ann. Rev. Earth Planet Sci., 17, 309-334, 1989. 124.   J.J. Walsh and J Waterson, Analysis of the relation between displacements and dimensions of faults, J. Struct. Geol., 10, 238-247, 1988. 125.   W.L. Power, T.E. Tullis, and J.D. Weeks, Roughness and wear during brittle faulting, J. Geophys. Res., 93, 15,268-15,278, 1988. 126.   P. Segall and D.D. Pollard, Nucleation and growth of strike slip faults in granite, J. Geophys. Res., 88, 555-568, 1983. 127.   A. Nur, H. Ron, and O. Scotti, Kinematics and mechanics of tectonic block rotations, in Slow Deformation and Transmission of Stress in the Earth, S.C. Cohen and P. Vanicek, eds., American Geophysical Union, Geophysics Monograph 49, Washington, D.C., pp. 31-46, 1989. 128.   R.H. Sibson, An assessment of field evidence for “Byerlee” friction, Pure Appl. Geophys., 142, 645-662, 1994. 129.   See R.H. Sibson, Earthquakes and rock deformation in crustal fault zones, Ann. Rev. Earth Planet. Sci., 14, 149-175, 1986. 130.   C. Lapworth, The Highland controversy in British geology, Nature,32, 558-559, 1885; J.M. Christie, Mylonitic rocks of the Moine thrust-zone in the Assynt region, north-west Scotland, Trans. Geol. Soc. Edinburgh,18, 79-93, 1960; T.H. Bell and M.A. Ethridge, Microstructure of mylonites and their descriptive terminology, Lithos., 6, 337-348, 1973. 131.   R.H. Sibson, Fault rocks and fault mechanisms, J. Geol. Soc. London,133, 191-213, 1977; J. Magloughlin, F.M. Chester, and J. Spray, Penrose conference report: Fine-grained fault rocks, GSA Today, 6, 33-37, 1996. 132.   M.A. Etheridge and J.C. Wilkie, Grain size reduction, grain boundary sliding, and the flow strength of mylonites, Tectonophysics, 58, 159-178, 1979. 133.   R.H. Sibson, Fault zone models, heat flow, and the depth distribution of earthquakes in the continental crust of the United States, Bull. Seis. Soc. Am., 72, 151-163, 1982; R.H. Sibson, Roughness at the base of the seismogenic zone: Contributing factors, J. Geophys. Res., 89, 5791-5800, 1984. 134.   J.P. Evans and F.M. Chester, Fluid-rock interaction in faults of the San Andreas system: Inferences from San Gabriel fault rock geochemistry and microstructures, J. Geophys. Res., 100, 13,007-13,020, 1995; F.M. Chester and J.S. Chester, Ultracataclasite structure and friction processes of the Punchbowl fault, San Andreas system, California, Tectonophysics, 295, 199-221, 1998. 135.   Y.-G. Li, K. Aki, D. Adams, A. Hasemi, and W.H.K. Lee, Seismic guided waves trapped in the fault zone of the Landers, California earthquake of 1992, J. Geophys. Res., 99, 11,705-11,722, 1994. 136.   S.E. Schulz and J.P. Evans, Mesoscopic structure of the Punchbowl fault, southern

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    California, and the geological and geophysical structure of active strike-slip faults, J. Struct. Geol., 22, 913-930, 2000. 137.   H. Kanamori and T.H. Heaton, Microscopic and macroscopic physics of earthquakes, in Physics of Earthquakes, J.B. Rundle, D.L. Turcotte, and W. Klein, eds., American Geophysical Union Monograph, Washington, D.C., pp. 147-163, 2000. 138.   The other energy terms include seismic-wave radiation, gravitational energy, and the energy in the formation of new surfaces, all of which are thought to be small relative to heat dissipation (C.H. Scholtz, The Mechanics of Earthquakes and Faulting, Cambridge University Press, Cambridge, U.K., 439 pp., 1990). 139.   R.H. Sibson, Generation of pseudotachylyte by ancient seismic faulting, Geophys. J. R. Astron. Soc., 43, 775-794, 1975; R.H. Maddock, Melt origin of fault-generated pseudotachylytes demonstrated by textures, Geology, 11, 105-108, 1983. 140.   C.H. Scholz, Shear heating and the state of stress on faults, J. Geophys. Res., 85, 6174-6184, 1980. 141.   C.H. Scholz, J. Beavan, and T.C. Hanks, Frictional metamorphism, argon depletion, and tectonic stress on the Alpine fault, New Zealand, J. Geophys. Res., 84, 6770-6782, 1979. 142.   P. Molnar, W.-P. Chen, and E. Padovani, Calculated temperatures in overthrust terrains and possible combinations of heat sources responsible for the Tertiary granites in the Greater Himalaya, J. Geophys. Res., 88, 6415-6429, 1983. 143.   The fault-slip directions determined from historic offsets or from grooves and slickensides on exposed fault surfaces can be used as stress indicators in both extensional and compressional regimes. The horizontal component of normal-fault slip is taken as the direction of least principal stress, while the horizontal component of reverse-fault slip determines the direction of greatest principal stress. As in the case of seismologically derived focal mechanisms, these geological interpretations can be biased if the faulting occurs on misoriented planes of weakness; however, the bias associated with normal and reverse faults tends to be small, because strength anisotropy tends to be in the s1-s3 plane and not in planes containing the s2 axis. 144.   Dike intrusions, which are common in extensional regimes, are particularly reliable estimators of stress orientation (±5-10 degrees); even in regions with abundant geologic fabric, the orientation of the intrusions follows planes perpendicular to the axis of least principal stress (R.L. Christiansen and E.H. McKee, Late Cenozoic volcanic and tectonic evolution of the Great Basin and Columbia intermountain regions, in Cenozoic Tectonics and Regional Geophysics of the Western Cordillera, R.B. Smith and G.P. Eaton, eds., Geological Society of America Memoir 152, Boulder, Colo., pp. 283-312, 1978). 145.   This problem was first recognized by D.P. McKenzie (Relation between fault plane solutions and direction of principal stresses, Bull. Seis. Soc. Am., 59, 591-601, 1969). In his analysis, he assumed that the deviatoric stresses in the crust were much smaller than those needed to fracture intact rock and therefore faulting must occur on weak, preexisting fractures; he showed that for the general case, s1 must lie in the same quadrant as the P axis, but could differ by as much as 90 degrees. Using more realistic constraints on rock strength, C.B. Raleigh, J.H. Healy, and J.D. Bredehoeft (Faulting and crustal stress at Rangely, Colorado, in Flow and Fracture of Rocks, H.C. Heard, I.Y. Borg, and N.L. Carter, eds., American Geophysical Union, Geophysical Monograph Series 16, Washington, D.C., pp. 275-284, 1972) concluded that the bias was not likely to be more than 35-40 degrees. 146.   M.L. Zoback and M.D. Zoback (Faulting patterns in north-central Nevada and strength of the crust, J. Geophys. Res., 85, 275-284, 1980) found that composite fault-plane solutions for a large number of earthquakes agreed to within 3 degrees with the observed fault orientations from the northern Basin and Range, even though the standard deviation in the T-axis azimuth for an individual mechanism was substantial (25 degrees). 147.   J. Angelier, Tectonic analysis of fault slip data sets, J. Geophys. Res., 89, 5835-5848,

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    1984; J.W. Gephart and D.W. Forsyth, An improved method for determining the regional stress tensor using earthquake focal mechanism data: Application to the San Fernando earthquake sequence, J. Geophys. Res., 89, 9305-9320, 1984. This method can also be used to constrain the relative magnitude of the maximum deviatoric stress, s1 – s3, which governs the amount of scatter. 148.   B.C. Haimson and C. Fairhurst, In-situ stress determination at great depth by means of hydraulic fracturing, in Rock Mechanics—Theory and Practice, Proceedings, 11th Symposium on Rock Mechanics, Berkeley, 1969, W.H. Somerton, ed., Society of Mining Engineers of AIME, New York, pp. 559-584, 1970; A. McGarr and N.C. Gay, State of stress in the Earth’s crust, Ann. Rev. Earth Planet. Sci.,6, 405-436, 1978. Stress orientations can also be estimated from the strain observed inside boreholes after overcoring relieves a component of the in situ stress, although this method cannot be used far from the free surface and is more sensitive to local inhomogeneities. 149.   See summary comments in the special volume on mechanical involvement of fluids in faulting by S. Hickman, R. Sibson, and R. Bruhn (Introduction to special session: Mechanical involvement of fluids in faulting, J. Geophys. Res., 100, 12,831-12,840, 1995). 150.   R. Emmermann and J. Lauterjung, The German Continental Deep Drilling Program KTB: Overview and major results, J. Geophys. Res., 102, 18,179-18,201, 1997. 151.   See M.D. Zoback and R. Emmermann, eds., Scientific Rationale for Establishment of an International Program of Continental Scientific Drilling, Report of the International Conference on Scientific Drilling, Potsdam, Germany, 126 pp., August 20-September 1, 1993; H.C. Larsen and I. Kushiro, eds., Report of the Conference of Cooperative Ocean Riser Drilling (CONCORD), Tokyo, 116 pp., July 22-24, 1997.