83, No. 1, pp. 144-159, February 1993. AFTERSHOCK STRESS RELEASE ALONG ACTIVE FAULT ...... Mammoth Lakes, California, earthquake sequence, Bull.
Bulletin ofthe SeismologicalSocietyofAmerica,Vol. 83, No. 1, pp. 144-159, February 1993
A F T E R S H O C K STRESS RELEASE ALONG ACTIVE FAULT PLANES OF T H E 1984 ROUND VALLEY, CALIFORNIA, EARTHQUAKE S E Q U E N C E APPLYING A TIME-DOMAIN STRESS DROP ME T H O D BY KENNETH D. SMITH AND KEITH F. PRIESTLEY ABSTRACT The 23 November 1984 M L 5.8 Round Valley earthquake is one in a series of moderate ( M L ~ 6) earthquakes to have occurred in the Bishop-Mammoth Lakes, California, area since 1978. This earthquake and its aftershock sequence occurred within a dense seismic network, and hypocentral location quality is excellent. In a previous study, we determined that the Round Valley sequence involved faulting on a conjugate set of fault planes; one, a nearvertical plane striking N30°E, the mainshock fault plane showing principally left-lateral strike-slip motion, and another subperpendicular to the mainshock fault plane striking N40°W and dipping 55°NE, exhibiting dominantly right-lateral strike slip. This conjugate fault plane conforms to a postulated extension of the Hilton Creek fault and is the only significant activity on this structure in the 12-year Bishop-Mammoth Lakes earthquake sequence. Source dimensions and stress drops for 87 aftershocks ( M L 2.8 to 4.2) of the Round Valley sequence have been determined using an adaptation of the initial P-wave pulse width time-domain deconvolution technique of Frankel and Kanamori (1983). The aftershock sequence is confined to a limited volume of crust. We have shown that site and instrument effects and not whole-path attenuation control the minimum pulse widths for this limited region. The determination of a site minimum pulse width, rather than a minimum pulse width for each source receiver pair as in the Frankel and Kanamori study, makes the deconvolution procedure practical for processing the large numbers of events in an aftershock sequence. With the large data set available for the Round Valley affershock sequence, patterns of the stress drop along the active fault planes can be seen in detail. Source radii systematically increase with magnitude from about 100 m for events near magnitude 3.0 to 500 m for events near magnitude 4.0. Static stress drops range from 10 to 200 bars and are not strongly correlated with magnitude or depth. The stress release pattern reveals a broad stress drop low ( ~ r -- 10 bars) for aftershocks within the mainshock fault plane that is consistent with other evidence of the rupture surface of the Round Valley mainshock. Higher stress release occurs above and below the mainshock rupture surface and on the shallower, conjugate fault plane. Further distant from the rupture surface of the mainshock, stress drops decrease to average values. On the conjugate fault surface, stress drops are seen to be high in areas that may be interpreted as "off-fault" clusters with respect to the mainshock rupture surface.
INTRODUCTION Source p a r a m e t e r estimates for small e a r t h q u a k e s ( M L -- 3 - 4) are typically made using spectral techniques; the long-period spectral level is related to the seismic m o m e n t of the event and the spectral corner frequency is related to a source dimension t h r o u g h scaling relationships t h a t make assumptions about the source geometry and r u p t u r e velocity. For example, in the application of Brune's (1970, 1971) source theory, instrument-corrected Fourier displacement 144
A F T E R S H O C K S T R E S S R E L E A S E IN R O U N D VALLEY, C A L I F O R N I A
145
spectra for the direct S-wave arrival are parameterized by a constant lowfrequency level, an w-2 high-frequency decay, and a corner frequency, fc, defined as the intersection of these two trends. The low-frequency level is related to the seismic moment Mo, and the corner frequency to the source dimension. The success of such analysis depends on an accurate observation of the seismic source spectrum. Seismograph site response and whole-path attenuation m a y severely distort the observed seismic spectrum, causing interpretation of spectra in terms of the earthquake source, especially of the corner frequency, difficult. The influence of the site response has been considered in very few studies b u t m a y significantly affect the interpretation of the observed spectra radiated from microearthquakes. For example, Archuleta et al. (1982) determined spectral source parameters for a wide magnitude range of events of the 1980 M a m m o t h Lakes earthquake sequence. They found that, for events with seismic moment greater than about I × 1021 dyne-cm ( M n ~ 3.2), the stress drop is nearly constant at around 50 bars, while for events with seismic moment less than 1 × 1021 dyne-cm the stress drop decreased with decreasing seismic moment. This implies that the source radius for small earthquakes decreases much more slowly with seismic moment than predicted by cube root scaling. Archuleta et al. (1982) argue that their observation cannot be a result of whole-path attenuation and conclude that the cube root scaling relationship does not appear to be correct for events with seismic moment less than 1 × 1021 dyne-cm in this area. They extend this conclusion to other regions by citing various other published studies where the same observation was made. H a n k s (1982) found a high-frequency band-limittion for acceleration spectra of California earthquakes and termed this fmax" Regardless of source strength and tectonic environment, fmax w a s found in a narrow frequency band between 10 and 25 Hz. H a n k s attributed this to the local recording site conditions. On the other hand, Papageorgio and Aki (1983a, b) suggest that the observed band limitation and the m a x i m u m corner frequency result from the existence of a minimum source radius, which correlates with a physical property of the fault surface, the "barrier" size. Recent work comparing surface and borehole recordings of the same microearthquakes have given support to the suggestion that the band limitation and m a x i m u m corner frequency is a local recording site effect. Carroll et al. (written comm.) have discussed data from two borehole sites within the Anza seismic array in southern California. For these sites they find that, over a large moment range, downhole displacement spectral shapes agree with those predicted by the Brune spectral model, which is a constant low-frequency spectral level below some corner frequency, above which the spectral level decays as w -2. Borehole spectra show no band-limitation or m a x i m u m corner frequency, at least to 100 Hz. Surface displacement spectra of the same events show resonance effects at lower frequencies ( = 1-10 Hz) and fall off much more steeply (~o-3 to ~o-5) at high frequencies. While the study of Carroll et al. (written comm.) does not discount that there exists some band-limitation associated with source properties, their results do suggest that for the Anza region in Southern California these effects must correspond to frequencies greater than 100 Hz. In several studies (Mueller, 1985; Frankel et al., 1986; Li and Thurber, 1988), the source spectrum of a larger earthquake is recovered by deconvolving from
146
K. D. S M I T H
A N D K. F. P R I E S T L E Y
its observed spectrum the observed spectrum of a small event. This small event is considered to be the impulse response of the propagation path (an empirical Green's function). When both events have the same hypocentral location and focal mechanism, parameters common to both events, whole-path attenuation, instrument response, and local site effects, are removed in the deconvolution procedure, leaving that part of the signal of the larger event describing the actual source process. O'Neill and Healy (1973) proposed a time-domain method for recovering source parameters of microearthquakes using data recorded on the USGS short-period vertical type seismographs. They corrected the initial P-wave pulse width, defined as the time interval from the first arrival to the first zero crossing (T1/2), measured on seismograms recorded at small hypocentral distances, for the effect of the instrument and attenuation. They then estimated the source rupture duration and stress drop in terms of ~1/2 by comparing observed and theoretical pulse widths. O'Neill (1984) used this method to estimate source parameters of 30 small Parkfield, California, earthquakes. Frankel and Kanamori (1983) extended the initial P-wave pulse width method by incorporating a method of accounting for the seismograph site response. They showed that for each seismograph site there was a magnitude level below which the pulse width no longer decreased: a pulse width floor. The frequencydomain deconvolution could be approximated in the time domain by subtracting the initial P-wave pulse width of a co-located event whose magnitude was smaller than this level from that of the larger event. In this way, the P-wave pulse broadening due to factors common to both seismograms, including site response, were eliminated, and the source information of the larger event was thereby isolated. Using this technique, Frankel and Kanamori (1983) determined the source dimensions of several M L 3.5 to 4 events over a large area within the southern California network. In this paper, we analyze initial pulse widths to determine source parameters for aftershocks of the 1984 Round Valley, California, earthquake sequence. Frankel and Kanamori (1983) defined minimum T1/2 values for the path associated with each source receiver pair. The events we have studied occur in a limited crustal volume and are recorded on a dense local (epicentral distances less than 40 km) seismic network (Fig. 1). This allows for some simplification in the technique proposed by Frankel and Kanamori (1983), in particular establishing the ability to define individual station corrections rather than defining individual event corrections. We have applied this time-domain technique to 87 aftershocks of the 1984 Round Valley, California, earthquake sequence. From these measurements, we have been able to map the stress drops and stress release along the two active faults of the sequence. THE 1984 ROUNDVALLEYEARTHQUAKESEQUENCE Details of the temporal and spatial development of the Round Valley earthquake sequence have been discussed by Priestley e t al. (1988) and are only summarized here. They located more than 1100 events with a standard error better than 1.0 km in epicenter and 2.0 km in depth during the first 3 weeks of the sequence. This proved adequate to resolve the three-dimensional complexities in the faulting process. The sequence began on 23 November 1984 with an M L 2.8 foreshock followed 4 sec later by the M L 5.8 mainshock. Aftershocks of
AFTERSHOCK
STRESS
RELEASE
IN ROUND 119.0 °
VALLEY,
CALIFORNIA
147
118.5 °
FIG. 1. Round Valley aftershock pattern and short period seismograph stations of the UNR-USGS in the Bishop-Mammoth Lakes California area. The seismograph stations used in determining earthquake locations are indicated by solid triangles in the upper plot. The 11 seismograph used in determining the P-wave pulse widths are shown with larger triangles and the major mapped faults in the region are denoted by solid lines. Topographic contours shaded with white represent elevations below 1000 meters, and dark shading represents elevations above 3000 meters. The large scale plot shows the aftershock patterns for the Round Valley sequence. The octagons labeled 1, 2, 3 are the locations of 1, M L 5.8 main shock, 2, ML 5.2 November 23, 1984; 19:12 aftershock, 3, ML 4.8 November 26, 1984; 16:23 aftershock. The cross section views shown in Figure 2 are oriented by lines A-A' and B*B'.
t h e R o u n d Valley s e q u e n c e occurred on two conjugate f a u l t p l a n e s a n d involved little, i f any, n o r m a l f a u l t i n g along t h e e a s t e r n e s c a r p m e n t of t h e S i e r r a N e v a d a . F i g u r e 1 is a m a p of seismicity from t h e U n i v e r s i t y of N e v a d a catalog for t h e R o u n d Valley a r e a from 23 N o v e m b e r t h r o u g h 31 D e c e m b e r 1984. T h e octagons labeled 1,2,3 are the locations of t h e principle e v e n t s of t h e s e q u e n c e (see figure caption). T h e first f a u l t p l a n e of t h e c o n j u g a t e p a i r to develop in t h e R o u n d V a l l e y s e q u e n c e was associated w i t h t h e m a i n s h o c k , is n e a r l y vertical, a n d strikes N30°E. T h e second, which b e c a m e active w i t h i n t h e first 24 h o u r s of t h e sequence, strikes N40°W a n d dips 55 ° to the n o r t h e a s t . We will r e f e r to t h e first p l a n e to develop in t h e sequence, associated w i t h t h e m a i n s h o c k , as t h e p r i m a r y f a u l t plane, a n d t h e second f a u l t p l a n e to develop as t h e c o n j u g a t e f a u l t plane. F i g u r e s 2a a n d b shows t h e seismicity cross sections p e r p e n d i c u l a r a n d p a r a l l e l to t h e p r i m a r y f a u l t plane. T h e seismicity t h a t o c c u r r e d on t h e p r i m a r y p l a n e is m a i n l y r e s t r i c t e d to t h e footwall block of the conjugate fault plane, a n d the fault m o t i o n is d o m i n a n t l y left-lateral s t r i k e slip. In t h e h a n g i n g wall block of t h e conjugate plane, t h e r e is a less d e n s e c o n c e n t r a t i o n of a f t e r s h o c k activity. Activity on t h e conjugate p l a n e s p r e a d to t h e w e s t a n d shallowed as t h e s e q u e n c e progressed. T h e d o m i n a n t sense of m o t i o n on t h e conjugate p l a n e is r i g h t - l a t e r a l s t r i k e slip, a l t h o u g h the w e s t e r n e x t e n t of t h e a f t e r s h o c k distribution is c h a r a c t e r i z e d b y shallow ( = 5 km) n o r m a l faulting.
148
K. D. S M I T H AND K. F. P R I E S T L E Y Main Shock Fault Plane A
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(b) FIG. 2. (a) A-A' seismic cross section along the N30°E m a i n shock fault plane. The subcircular area largely free of aftershock activity, (dashed line) which is located slightly above the m a i n shock hypocenter is interpreted as the mainshock r u p t u r e surface (Priestley et al., 1988). (b) B-B' along the N40°W conjugate fault plane t h a t conforms to a n extension of the Hilton Creek Fault. The numbers in (a) a n d (b) correspond to the events in Figure 1.
Moment tensor inversion of seismic body waves, the first-motion focal mechanism of the main shock, and the narrow vertical distribution of early aftershock activity all support left-lateral strike slip on the near-vertical primary plane as the source mechanism of the mainshock (Priestley et al., 1988). In addition, early aftershock activity is consistent with unilateral rupture to the southwest. A cross section of the mainshock fault plane (Fig. 2a) indicates a lack of aftershock activity in a subcircular region of the primary plane to the southwest
AFTERSHOCK STRESS RELEASE IN ROUND VALLEY, CALIFORNIA
149
of the mainshock hypocenter (dashed line). This was interpreted by Priestley et al. (1988) as the section of the fault that ruptured during the main event. The area of this subcircular region is approximately 30 km 2. This value and the surface-wave moment of the mainshock of 8.0 × 1024 dyne-cm corresponds to a stress drop of 23 bars and total slip of 86 cm (Priestley et al., 1988). The mainshock source spectrum, constructed from near-source strong-motion data and long-period body waves and surface waves (Priestley et al., 1988), indicates a corner frequency of 0.2 Hz, consistent with the source area suggested by the aftershock distribution. For the present study, stress drops determined from the analysis of P-wave pulse widths will be interpreted within the framework of the aftershock distribution and the rupture characteristics of the mainshock that were found in the Priestley et al. (1988) study. PULSE WIDTH DETERMINATION
The data used in determining the P-wave pulse widths are short-period vertical-component seismograms from the University of Nevada, Reno--U.S. Geological Survey (UNR-USGS) telemetered seismic array. Data were selected from seismographs at hypocentral distances less than 40 km so as to ensure that the first arrivals were direct arrivals, and from a range of azimuths to adequately sample the focal sphere. The seismograph stations used in the study (Fig. 1) consist of vertical-component 1-Hz seismometers. These signals are telemetered to a central recording site on the U N R campus, digitized at 50 Hz, and digitally recorded. The seismometer output at the frequencies of interest to this study ( = 1 to 5 Hz) is flat to velocity. For the same instrument response, Frankel and Kanamori (1983) found that the pulse widths measured from the seismograms were within 0.012 sec, equivalent to the pulse widths of the true ground velocity. At 40 km distance, a M n 3 event normally exceeds the dynamic range of the USGS analog seismograph and the records are clipped. However, since the zero crossing times are preserved even for signals that saturate the seismograph electronics (Ellis and Lindh, 1976), accurate pulse widths can be determined from clipped data. We have measured over 2000 P-wave pulse widths from more than 200 Round Valley aftershocks ranging from MLD (local duration magnitude scale) 1.5 to 4.5. The resolution in the pulse width determination is somewhat limited by the sample rate of 50 Hz. This is due to the variation of the position of the first arrival between two discrete samples in the digitized data. Frankel and Kanamori (1983) suggested the resolution for this sample rate could be interpolated to approximately one quarter of a sample (0.005 sec for 50-Hz data). For example, if the initial slope between the first two samples, bracketing the first arrival, is steep, the P-wave arrival is more near in time to the first sample; on the other hand, if the initial slope is very gradual then the P-wave arrival is more near the latter sample. There are clear examples of small events with identical waveforms at a particular station that do show this initial slope effect in the first arrivals. It becomes a simple m a t t e r of extrapolation between these two samples to estimate, to a much greater precision than the sample rate would suggest, the precise P-wave arrival time. The time to the first zero crossing is then determined by calculating the pre-event average trace amplitude and again extrapolating between samples at the zero crossing. There are
150
K. D. S M I T H AND K. F. P R I E S T L E Y
variations in noise level at some of the stations during the period of this study that affect the estimates of pulse width. Figure 3 compares P-wave pulses for MLD 3.1 and 4.1 events. The initial pulse width measured on the seismogram is a function of the rupture duration, the instrument response, and the broadening caused by the apparent attenuation along the path, including both intrinsic attenuation, scattering, and site effects. Frankel and Kanamori (1983) noted that, as the magnitude of southern California earthquakes decreased to about 2.2, there was a proportional decrease in the initial pulse width of the first arrival. Below about M L 2.2, the pulse width remained about constant as magnitude decreased further. They assumed that these small events ( M L < 2.2) could be taken as point sources, and hence their seismograms were the impulse response of the combination of the path and instrument factors. Earthquake source durations can only be determined for events whose rupture times are sufficiently longer than the broadening effect caused by the path and instrument. To correct the pulse width measure for instrument and path effects in our study, minimum ~1/2's have been determined for each station. Figure 4 is an example for two stations used. The minimums are associated with the lowest magnitude events ( M L < 2.0). These plots include all pulses widths recorded at a particular site for the entire magnitude range ( M L 1.5 to 4), to not only illustrate the minimum but the distribution over the hypocentral range. We have chosen the minimum through the scatter of the lowest values at each station to account for picking errors introduced by noise and the limited resolution due to the 50-Hz sample rate. Pulse widths are plotted versus hypocentral distance with y-axis divisions of one sample, 0.02 sec. The lack of variation in the minimum pulse width with hypocentral distance or hypocentral depth for the earthquakes in this limited crustal volume for all source receiver pairs suggests a minimum pulse width for each station. This indicates that, for these data, variations in whole-path Q is not a significant factor in the observed variations in broadening of the P-wave pulse at the studied sites. Following the path attenuation arguments of O'Neill (1984) in his pulse width study in the Parkfield area, low Q values in the neighborhood of 200 would return a t* (t* = travel time/average Q) of approximately 0.04 sec for a 40-km travel path. This variation is not seen in the pulse width versus distance plots. The consistency in the pulse width with distance is evidence that whole-path Q is much higher for the study area and is not a significant factor.
Pulse Width Comparison
FIG. 3. Example of P-wave pulses a n d picks for events differing in size by about 1 magnitude unit.
AFTERSHOCK
S T R E S S R E L E A S E IN R O U N D VALLEY, C A L I F O R N I A
151
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(b) FIG. 4. Determination of m i n i m u m pulse width for two of the 10 stations used in the study. Divisions of Zl/2 y axis, r e p r e s e n t one sample (0.02 sec) of the digital record. The x axis is t h e hypocentral digt~nce in km. The dashed line denotes the value of zl/2~ chosen for the station.
152
K. D. SMITH AND K. F. PRIESTLEY RUPTURE DURATION AND SOURCE PARAMETERS
The pulse width A ( t ) observed on a seismograph is a convolution of the source-time function S(t), the path Greens Function G(t), and the i n s t r u m e n t response I(t): A(t) = S(t)*G(t)*I(t).
Several published studies have assumed S ( t ) ~ ~(t) for very small events and determined S ( t ) for larger co-located events by deconvolution. This procedure cannot be applied directly to clipped seismograms. Frankel and Kanamori (1983) showed t h a t the deconvolution could be approximated by subtracting the pulse width of the smaller event from t h a t of the larger event, T1/2sourc e ~
T1/2obser,e d --
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where rl/2 is the estimated source duration, 71/2ob..... ~ is the observed pulse width of the larger event, and T1/2~ is the path, instrument, and site contribution. Since the zero crossing times are preserved for clipped records, this procedure can be used to approximate the deconvolution and estimated source duration component of the source-time function from clipped seismograms. The pulse width minima determined for each site from Figure 4 were subtracted from the measured pulse widths of the larger events (2.8 < M L < 4.2) recorded at those station. Events for which five or more individual station ~1/2's were available were averaged, and the m e a n T1/2 was t a k e n as the effective source duration for the event. Averaging T1/2 values for a particular event over a range of azimuths tends to reduce the effects of trace noise and timing errors and smooth over the radiation pattern. As in the Frankel and Kanamori (1983) study, the source radius, r, was calculated from Boatwright's (1980) relationship: ......
r-
• 1/2 v ~)
1 + -sin 0 ¢
w h e r e v is the rupture velocity (0.80 of an assumed S-wave velocity of 3.2 km/sec), c is the P-wave velocity (assumed 6 km/sec), and 0 is the azimuth between the unit vector normal to the fault plane and the seismograph site (assumed to be 45°). Source dimensions have been calculated for 87 events and are summarized in Table 1. Since all of the events come from a limited volume of the upper crust, the relative scaling of the source dimensions between individual events should be preserved. The true measure of the source dimension is difficult to determine, since the rupture velocity and rupture propagation geometry will vary between events from our assumed values. Since the ray paths for each event cover a complete range of azimuths, attenuation for paths through the fault plane of the mainshock will be minimized by averaging each value of T1/2...... at all stations for a particular event. A wide range of ray paths are sampled in the determination of the pulse minimum, T1/2~, for each station, which tends to average attenuation effects across the fault plane and smooths any differential attenuation within the limited source region. These averaging processes and sampling of a wide variety
153
A F T E R S H O C K S T R E S S R E L E A S E IN R O U N D VALLEY, C A L I F O R N I A TABLE 1 Origin Time
Stress Drop (bars)
Source Radius (meters)
Log M o
MLD
41123 2342 41124 0029 41124 0031 41124 0235 41124 0246 41124 0316 41124 0419 41124 0510 41124 0740 41124 0748 41124 0800 41124 0834 41124 0841 41124 0844 41124 0851 41124 0853 41124 0921 41124 1013 41124 1025 41124 1122 41124 1215 41124 1232 41124 1255 41124 1308 41124 1349 41124 1646 41124 1801 41124 1804 41124 1820 41124 1830 41124 1857 41124 1939 41124 1946 41124 2025 41124 2029 41124 2033 41124 2103 41124 2110 41124 2135 41124 2324 41125 0014 41125 0058 41125 0145 41125 0230 41125 0241 41125 0827 41125 0844 41125 1037 41125 1555 41125 1615 41125 1617 41125 1625 41125 1914 41125 1918 41125 2148
19 18 88 11 47 22 23 12 11 23 9 24 4 52 31 178 129 33 89 33 11 25 19 9 13 3 9 47 3 9 10 20 5 17 79 110 8 9 6 9 11 42 26 29 11 32 282 13 225 106 149 33 3 64 28
408 347 196 290 189 157 147 244 349 221 238 140 218 252 110 133 386 273 114 232 270 285 127 166 203 144 221 104 214 180 191 195 143 417 174 134 111 346 197 166 227 151 106 147 138 160 114 236 123 354 121 129 327 173 247
21.48 21.24 21.18 20.80 20.86 20.28 20.22 20.60 21.03 20.75 20.43 20.18 20.01 21.28 19.98 20.98 22.23 21.18 20.48 20.98 20.68 21.13 19.96 19.98 20.40 19.37 20.34 20.08 19.77 20.10 20.18 20.54 19.51 21.46 20.98 20.78 19.41 20.94 20.04 19.98 20.45 20.52 19.85 20.33 19.83 20.48 20.98 20.58 20.98 22.03 20.78 20.21 20.38 20.88 20.98
3.1 3.1 2.9 3.4 3.0 2.8 3.2 3.0 3.0 3.1 3.2 3.0 2.7 3.1 2.9 2.8 3.7 3.2 2.9 2.9 3.4 3.3 3.0 2.9 3.3 2.8 3.3 2.8 3.3 3.4 2.8 3.5 3.0 4.2 3.1 3.1 2.9 3.7 3.1 3.1 3.3 3.5 2.8 3.2 2.8 3.2 3.3 3.1 3.4 3.8 2.9 3.1 2.9 3.3 3.4
154
K. D. SMITH AND K. F. PRIESTLEY TABLE 1--Continued
Origin Time
41125 2254 41125 2309 41126 1030 41126 1628 41126 1749 41126 1801 41126 1824 41126 2116 41127 0954 41127 1215 41127 1253 41127 1337 41127 1406 41127 1647 41127 1731 41127 1928 41127 2006 41127 2125 41128 0241 41128 0054 41128 0821 41129 1042 41129 1045 41129 1309 41129 1750 41130 0003 41201 0455 41201 2014 41203 2038 41207 2205
Stress Drop (bars)
Source Radius (meters)
Log M o
MLD
17 10 16 49 6 2 14 11 17 25 10 16 12 18 23 41 30 41 77 9 29 61 71 21 12 32 3 56 4 4
280 789 203 234 211 384 307 157 156 336 219 149 189 193 373 113 121 127 88 178 166 222 203 272 204 174 764 256 210 220
20.93 22.05 20.48 21.16 20.13 20.37 20.97 19.98 20.18 21.33 20.38 20.08 20.28 20.48 21.43 20.13 20.08 20.28 20.08 20.08 20.48 21.18 21.13 20.98 20.38 20.58 21.48 21.33 19.97 19.97
3.7 4.4 2.9 3.4 2.8 3.7 3.7 2.8 2.9 3.7 3.4 2.8 3.4 3.5 4.1 3.1 2.8 3.2 2.8 3.0 2.8 3.5 3.4 3.3 3.1 3.1 3.5 3.1 2.9 2.9
of a z i m u t h s b o t h in d e t e r m i n a t i o n of T1 / 2 . . . . . . a n d Ti/2~ t e n d to m i n i m i z e r a n d o m e r r o r s a n d i n c r e a s e t h e s t a b i l i t y of t h e e s t i m a t e d s t r e s s d r o p s . SEISMIC MOMENTS W e h a v e e s t i m a t e d t h e s e i s m i c m o m e n t s , Mo, f r o m t h e e a s t - w e s t c o m p o n e n t of m o t i o n r e c o r d e d on t h e L a w r e n c e L i v e r m o r e N a t i o n a l L a b o r a t o r y ( L L N L ) b r o a d b a n d d i g i t a l s e i s m i c s t a t i o n a t M i n a , N e v a d a , l o c a t e d 110 k m N N W of t h e R o u n d V a l l e y a f t e r s h o c k zone. T h i s c o m p o n e n t of m o t i o n is n e a r l y t r a n s v e r s e to t h e R o u n d V a l l e y a f t e r s h o c k r e g i o n , a n d t h e t r a v e l p a t h is n e a r l y c o m m o n for all e v e n t s . T h e r e f o r e a n y r e l a t i v e v a r i a t i o n in t h e s p e c t r a b e t w e e n e v e n t s will b e d u e to v a r i a t i o n i n s e i s m i c m o m e n t , t h e s o u r c e d e p t h , a n d focal m e c h a n i s m . T h e Lg a m p l i t u d e s p e c t r a w e r e c o r r e c t e d for w h o l e p a t h a t t e n u a t i o n u s i n g t h e G r e a t B a s i n Lg a t t e n u a t i o n f u n c t i o n , Q ( f ) , d e t e r m i n e d b y C h a v e z a n d P r i e s t l e y (1985). W e h a v e c h o s e n a n a r r o w f r e q u e n c y b a n d (0.8 to 1.2 Hz) o v e r w h i c h to e s t i m a t e t h e s e i s m i c m o m e n t , since t h e s i g n a l i n t h i s b a n d is c o n s i s t e n t l y a b o v e the noise and substantially lower in frequency than corner frequencies expected for e v e n t s in t h e m a g n i t u d e r a n g e we h a v e s t u d i e d ( S a v a g e , 1974). The seismic moment scales linearly with the low-frequency spectral amplit u d e b e l o w t h e c o r n e r f r e q u e n c y for e v e n t s w i t h a c o m m o n t r a v e l p a t h a n d a
AFTERSHOCK STRESS RELEASE IN ROUND VALLEY, CALIFORNIA
155
common focal mechanism. We have calibrated the Lg spectral level with seismic moments by fitting the lower-frequency ( -- 0.1 Hz) surface waves for the largest (M L 3.8 to 4.2) events for which we had a P-wave first-motion focal mechanism, with synthetic surface-wave seismograms. These seismograms were computed using the mode summation code of Gomberg et al. (1988), the Basin and Range velocity structural model of Priestley and Brune (1978), and the Basin and Range attenuation model of Patton and Taylor (1984). Seismic moments determined from surface-wave modeling were consistent with the seismic moment estimates obtained using the moment-magnitude relationship for the BishopMammoth Lakes region determined by Chavez and Priestley (1985). The seismic moment determined by this procedure was then used to determine the relationship between seismic moment and the Lg spectral amplitudes for the Round Valley aftershocks. We have not accounted for the focal mechanism or focal depth in deriving the scaling relationship between Lg spectral amplitude and the seismic moment. Although the amplitude of fundamental-mode Love waves is sensitive to the focal mechanism, the 0.8- to 1.2-Hz band is dominated by Lg higher-mode surface waves, which sample a wide portion of the focal sphere and are therefore much less sensitive to variations in the focal mechanism. Source depth and hypocentral elastic parameters m a y have a much more significant effect on the Lg spectral amplitudes. To estimate the magnitude of these, we computed a suite of transverse Lg synthetic seismograms for a range of depths extending from 3 to 13 km and the two principle focal mechanisms observed in the Round Valley aftershock sequence. The synthetics were computed using the reflectivity method (Fuchs and Mueller, 1971) and for the Great Basin velocity structure of Benz et al. (1990). The Lg spectral amplitudes determined from these synthetic seismograms showed no systematic variation with focal mechanism and, at most, a factor of 2 variation in amplitude in the 0.8- to 1.2-Hz band with depth. This is less than other uncertainties in determining the seismic moment. Determination of the absolute seismic moment for M L 3 to 4 is difficult, and we feel that at least there is consistency in the relative scaling of the moments between events for this localized source region. STRESS DROP DISTRIBUTION
Once the seismic moment, M0, and source radius, r, are determined, the stress drop, A ~r can be calculated from the relationship: Ao---
7 M0 16 r 3 "
Stress drops for the 87 aftershocks of the Round Valley sequence have been determined (Table 1), and these values have been intepreted with respect to the spatial seismicity distribution. Those events occurring on each of the conjugate fault planes have been isolated and plotted in cross sectional views. Figure 5a is a map of stress drop values on the primary fault plane. This plane is the same view as seismicity cross-section Figure 2a. The hypocenter of the mainshock is indicated by the octagon containing the '1.' The low stress-drop values are concentrated in the center of the plot, and the minimum to the southwest of the mainshock hypocenter is consistent with what Priestley et al. (1988) have interpreted as the section along the mainshock fault plane that
156
K. D. S M I T H
AND
K. F. P R I E S T L E Y
Main Shock Fault Plane A
A'
0
47 2~26 32 111~ 23
5
1719 17
111o
56 14
•"
4
52 2521 9
178
.12 '0 3
31 3Z
l~S
".
...
\
9
10
Io
19
}6
1 1 .....................[ ~ - . . . - ~ . . . [ ~ 42 17 71 18 79 61 13 23 110
Z0
15
71
0
15
10
Distance (km) (a) B
B;
o
3 4
47
5
1~ %8
6
47 24
12
6149 106282
77
e~
12 49 2
225
33
ed
"~ 12
29 11
13 9 16
@ 23
18 0
5
10
15
Distance (km) (b) FIG. 5. (a) Stress drops values along the vertical m a i n s h o c k fault plane striking N30°E. This is t h e s a m e view as Figure 2a. The interpretation of t h e rupture surface is s h o w n as a dashed line as in Figure 2a. The m a i n shock and primary aftershock are also s h o w n as in Figure 2a. (b) Stress drop values along the conjugate fault plane striking N40°W and dipping 55°NE. T h e m a i n s h o e k is labeled as 1. This is a depth projection of t h e dipping plane and not a true vertical cross section projection as s h o w n in Figure 2b.
ruptured during the M L 5.8 main event and the primary aftershock. This zone of low stress release that exists throughout the aftershock sequence defines this area as well as does the seismicity cross section (Fig. 2a). These two independent lines of evidence suggest nearly complete stress release within this area of the mainshock fault plane and suggest that the Round Valley mainshock may have resulted from a single asperity failure. Also consistent with theories of asperity
AFTERSHOCK STRESS RELEASE IN ROUND VALLEY, CALIFORNIA
157
failure and fault rupture (Madariaga, 1974), stress drops are relatively higher around the edges of the proposed rupture surface, where the concentration of aftershocks activity and stress drops are higher. Within this outline of relatively higher stress drops are also lower values in the range of those found within the proposed asperity failure. In other words, local stress drop maximums along the mainshock fault plane are confined to the edges of the interpreted mainshock rupture surface. Below the zone of low stress release are higher values associated with the base of the rupture area. At shallow depth, the stress drop values are less consistent. This m a y be due to some degree to the fact that the conjugate plane intersects the mainshock plane along a particular line, and it is difficult to isolate which events are on a particular fault plane; short-period focal mechanisms are nearly identical for events on the two planes. A projection of stress drops on the conjugate plane is shown in Figure 5b. This plot is a projection of the stress release along the conjugate plane and does not directly correspond to the seismicity cross section of Figure 2b. The mainshock occurred near the intersection of the two planes and is denoted by the '1' within the octagon in the projection. The seismicity associated with this structure clearly shallowed to the west and conforms to a proposed extension of the Hilton Creek Fault. This would be the first significant activity on this major holocene structure since the present Bishop-Mammoth Lakes sequence began 12 years ago. The stress drops patterns do not directly indicate an increase in stress drop with depth but are highest near the base of the westward shallowing seismicity. Also there is a well-defined high (several events) approximately 5 km westward from the intersection of the two planes. This m a y be due to an "off-fault" shear stress increase as a result of displacement on the mainshock fault plane (Das and Scholz, 1981) or m a y represent a state of stress along this southern extension of the Hilton Creek. In Das and Scholz' calculations, a region approximately three quarters of the source diameter on either side of the slip surface experiences an increase in shear stress. Smith and Priestley (1988) have implicated "off-fault" clustering in the triggering mechanism of the M L 6.4 Chalfant Valley, California, earthquake. The Chalfant sequence, 15 km northeast of the Round Valley area, also exhibited conjugate strike-slip faulting (Smith and Priestley, 1988). DISCUSSION AND CONCLUSIONS
Integrating the results from a detailed seismotectonic analysis of the Round Valley sequence (Priestley et al., 1988) allows us to interpret the stress drop results with respect to the operative faults. The consistency between the details of the mainshock rupture characteristics, and the stress drop pattern along the mainshock fault plane, increase our confidence in the pulse width technique. An area of low stress drop is coincident with the section of the fault that was interpreted to have ruptured during the mainshock. This area was isolated in the previous study due to its proximity to the hypocenter of the mainshock, the distribution of early aftershock activity, the focal mechanism of the mainshock, and the low level of aftershock activity over this area throughout the sequence. The stress drop distribution provides an additional line of evidence that the Round Valley earthquake was the result of a single asperity failure. In this interpretation, the wider range of stress drops that outline this region would be indicative of stress increases at the edge of the rupture surface.
158
K. D. SMITH AND K. F. PRIESTLEY
I f t h e h i g h e r s t r e s s - d r o p e v e n t s below t h e low s t r e s s - d r o p r e g i o n do define t h e edge of t h e r u p t u r e s u r f a c e , t h e R o u n d V a l l e y e v e n t did not r u p t u r e t h r o u g h t h e b a s e of t h e s e i s m o g e n i c zone. T h i s is also c o n s i s t e n t w i t h t h e a f t e r s h o c k distribution. T h e h i g h e s t s t r e s s d r o p s c o n f o r m to t h e b a s e of t h e N W - s t r i k i n g c o n j u g a t e p l a n e a p p r o x i m a t e l y 4 to 5 k m d i r e c t l y to t h e w e s t of t h e m a i n s h o c k r u p t u r e p l a n e ' s s t r e s s drop low. T h i s is p r e d i c t e d to be a zone of h i g h s h e a r s t r e s s (Das a n d Scholz, 1981) w i t h r e s p e c t to slip on a n a d j a c e n t s t r u c t u r e . "Off-fault" c l u s t e r i n g h a s b e e n o b s e r v e d in s e v e r a l e a r t h q u a k e s e q u e n c e s ( H a m i l t o n , 1972; W a r d et al., 1974; H u t t o n et al., 1980; M a g i s t r a l e et al., 1989). I n t h e e a r l i e r R o u n d V a l l e y study, t h i s a c t i v i t y w a s s h o w n to c o n f o r m to a p r o p o s e d s o u t h e r n e x t e n s i o n of t h e H i l t o n C r e e k F a u l t . T h e s i m p l i c i t y of t h e m o d i f i e d p u l s e w i d t h t e c h n i q u e is its r e a l v a l u e to s e i s m o t e c t o n i c studies. S h o w i n g t h a t p a t h c o r r e c t i o n s a r e not n e c e s s a r y a n d t h a t only i n d i v i d u a l s t a t i o n c o r r e c t i o n s a r e r e q u i r e d allows for t h e p r o c e s s i n g of m a n y e v e n t s in a p a r t i c u l a r region or w i t h i n a p a r t i c u l a r f a u l t zone w i t h r e l a t i v e ease. A l t h o u g h we w e r e able to s h o w t h a t p a t h effects w e r e m i n i m a l for t h e s m a l l s o u r c e - r e c e i v e r d i s t a n c e s for t h e R o u n d V a l l e y study, t h i s m a y n o t be t h e case for r e g i o n a l studies. Also, b y u s i n g m a n y s t a t i o n s c o v e r i n g a wide r a n g e of a z i m u t h s , t h e r a n d o m e r r o r s d u e to p i c k i n g e r r o r s a r e m i n i m i z e d a n d t h e s t a b i l i t y of t h e t e c h n i q u e is i n c r e a s e d . O u r confidence in t h e seismic m o m e n t d e t e r m i n a t i o n w a s also i n c r e a s e d in t h i s s t u d y b y h a v i n g a c o m m o n L g p a t h f r o m t h e l i m i t e d source region. T h e p u l s e w i d t h m e t h o d c a n be u s e d to look a t t h e s t r e s s p a t t e r n s a l o n g active f a u l t zones w h e r e n e a r - s o u r c e s h o r t - p e r i o d d a t a a r e available. T h e simplicity of t h e t e c h n i q u e p r o v i d e s a m e a n s for t h e c o n t i n u o u s m o n i t o r i n g of s t r e s s p a t t e r n s w i t h i n active f a u l t zones w h e r e h i g h s a m p l e r a t e low d y n a m i c r a n g e d a t a a r e a v a i l a b l e b u t w h e r e b r o a d b a n d h i g h - q u a l i t y i n s t r u m e n t a t i o n is n o t in place. ACKNOWLEDGMENTS This research was supported by the United States Geological Survey under contract 14-08-001G1326. We thank John Anderson, William Peppin and William Walter for their critical reviews of the manuscript. Wally Nicks and Austin Wilson of the University of Nevada Seismological Laboratory maintained the short period network throughout the Round Valley sequence and are directly responsible for the complete data set acquired from the Round Valley sequence. REFERENCES Archuleta, R. J., E. Cranswick, C. Mueller, and P. Spudich (1982). Source parameters of the 1980 Mammoth Lakes, California, earthquake sequence, J. Geophys. Res. 87, 4595-4607. Barker, J. S. and T. C. Wallace (1985). A note on the teleseismic body waves from the 23 November 1984 Round Valley, California, earthquake, Bull. Seism. Soc. Am. 76, 883-888. Benz, H. M., R. B. Smith, and W. Mooney (1990). Crustal structure of the northwestern Basin and Range Province from the 1986 Program for Array Seismic Studies of the Continental Lithosphere seismic experiment. J. Geophys. Res. 95, 8123. Boatwright, J. (1980). A spectral theory for circular seismic sources: simple estimates of source dimension, dynamic stress drop and radiated energy, Bull. Seism. Soc. Am. 70, 1-28. Brune, J. N. (1970). Tectonic stress and the spectra of seismic shear waves from earthquakes, J. Geophys. Res. 75, 4997-5009. Brune, J. N. (1971). Correction, J. Geophys. Res. 76, 5002. Chavez, D. E.,and K. F. Priestley (1985). M L observations in the Great Basin and Mo of the 1980 Mammoth Lakes, California, earthquake sequence, Bull. Seism. Soc. Am. 75, 1583-1598. Das, S. and C. H. Scholz (1981). Off-fault aftershock clusters caused by shear stress increase, Bull. Seism. Soc. Am. 71, 1669-1675.
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Elias, J. and A. Lindh (1976). Linearity of VCO-discriminator system with respect to zero crossing times, U.S. Geol. Surv. Open-File Rept. 76-873. Frankel, A., J. Fletcher, F. Vernon, L. Harr, J. Berger, T. Hanks, and J. Brune (1986). Rupture characteristics and tomographic source imaging of M L = 3 earthquakes near Anza, southern California, J. Geophys. Res. 91, 12,633-12,650. Frankel, A. and H. Kanamori (1983). Determination of rupture duration and stress drop for earthquakes in southern California, Bull. Seism. Soc. Am. 73, 1527-1551. Fuchs, K. and G. Muller (1971). Computation of synthetic seismograms with the reflectivity method and comparison of observations, Geophys. J. R. Astr. Soc. 23, 417-433. Gomberg, J. and T. G. Masters (1988). Waveform modeling using locked-mode synthetic and differential seismograms: application to the determination of the structure of Mexico, Geophys. J. R. Astr. Soc. 94, 193-218. Hamilton, R. B. (1972). Aftershocks of the Borrego Mountain earthquake from April 12 to June 12, 1968, in The Borrego Mountain Earthquake of April 9, 1968, U.S. Geol. Surv. Profess. Pap. 787, 31-54. Hanks, T. (1982). fmax, Bull. Seism. Soc. Am. 72, 1867-1879. Hill, D. P., R. A. Bailey, and A. S. Ryall (1985). Active tectonic and magmatic processes beneath Long Valley caldera, eastern California: on overview, J. Geophys. Res. 90, 11,111-11,120. Hill, D. P., R. E. Wallace, and R. S. Cockerham (1985). Review of evidence on the potential for major earthquakes and volcanism in the Long Valley-Mono Craters-White Mountains regions of eastern California, Earthquake Predict. Res. 3, 571. Hutton, L. K., C. E. Johnson, J. C. Pechmann, J. E. Ebel, J. W. Gwen, D. H. Cole, and P. T. German (1980). Epicentral locations for the Homestead Valley earthquake sequence, March 15, 1979, Calif. Geol. May, 110-114. Li, Y. and C. H. Thurber (1988). Source properties of two microearthquakes at Kilauea volcano, Hawaii, Bull. Seism. Soc. Am. 78, 1123-1132. Madariaga, R. (1973). Dynamics of an expanding crack, Bull. Seism. Soc. Am. 66, 639-666. Magistrale, H., L. Jones, and H. Kanamori (1989). The Superstition Hills California, earthquakes of 24 November, Bull. Seism. Soc. Am. 79, 239-251. Mueller, C. (1985). Source pulse enhancement by deconvolution of an empirical Green's function, Geophys. Res. Lett. 12, 33-36. O'Neill, M. E. (1984). Source dimensions and stress drops of small earthquakes near Parkfield, California, Bull. Seism. Soc. Am. 74, 27-40. O'Neill, M. E. and J. H. Healy (1973). Determination of source parameters of small earthquakes from P-wave rise time, Bull. Seism. Soc. Am. 63, 599-614. Papageorgio, A. S. and K. Aki (1983a). A specific barrier model for the quantitative description of inhomogeneous faulting and the prediction of strong ground motion. Part I. Description of the model, Bull. Seism. Soc. Am. 73, 693-722. Papageorgio, A. S. and K. Aki (1983b). A specific barrier model for the quantitative description of inhomogeneous faulting and the prediction of strong ground motion. Part II. Description of the model, Bull. Seism. Soc. Am. 73, 953-978. Patton, H. J. and S. R. Taylor (1984). Q structure of the Basin and Range from surface waves, J. Geophys. Res. 89, 6929-6940. Priestley, K. F. and J. N. Brune (1978). Surface waves and the structure of the Great Basin of Nevada and western Utah, J. Geophys. Res. 83, 2265-2272. Priestley, K. F., K. D. Smith, and R. S. Cockerham (1988). The 1984 Round Valley, California earthquake sequence, Geophys. J. R. Astr. Soc. 95, 215-235. Savage, J. C. (1974). Relation of corner frequency to fault dimension, J. Geophys. Res. 77, 3788-3795. Smith, K. D. and K. F. Priestley (1988). The foreshock sequence of the 1986 Chalfant, California earthquake, Bull. Seism. Soc. Am. 78, 172-187. Ward, P. L., J. Gibbs, D. Harlow, and A. Aburto (1974). Aftershocks of the Managua, Nicaragua earthquake and the tectonic significance of the Tiscapa fault, Bull. Seism. Soc. Am., 64, 1017-1029. SEISMO. LAB./ 168 UNIVERSITYOF NEVADA RENO, NEVADA89557-1041 Manuscript received 27 March 1991
