Bulletin of the Seismological Society of America, Vol. 77, No. 1, pp , February 1987

Transcription

1 Bulletin of the Seismological Society of America, Vol. 77, No. 1, pp , February 1987 APPARENT STRESSES, STRESS DROPS, AND AMPLITUDE RATIOS OF EARTHQUAKES PRECEDING AND FOLLOWING THE 1975 HAWAII Ms = 7.2 MAIN SHOCK BY F. R. ZUI~IGA, M. WYss, AND M. E. WILSON ABSTRACT S-wave spectra for approximately 300 earthquake signals (2.8 <_- ML -<-- 3.9) were obtained from hand-digitized Wood-Anderson records from Hilo, Hawaii. All sources were located within part of the rupture volume of the 1975 Hawaii Ms = 7.2 earthquake. The epicentral distance to Hilo was about km, the backazimuth was (south), and the sources were located at depths of km. The apparent stress and Brune stress drop, based on 3-sec windows of the S waves, were examined for possible temporal and spatial dependence. Most of the data showed a remarkable constancy of the average apparent stress and stress drop per event as a function of time. No measurable change in average signal content took place at the time of the main shock. A statistically significant difference in the average stress parameters per event was found between neighboring subvolumes (radius -~ 5 km) of the 1975 affershock zone: a volume previously proposed as asperity based on seismicity pattern interpretation contained earthquakes with higher apparent stresses of ~ ± 8.8 bars compared to the neighboring volume with ~ _ 6.1 bars. Seventy-eight spectra obtained from a second station showed the same difference between the two volumes as that found at station Hilo, suggesting that the signal differences should be interpreted in terms of source differences. In addition, amplitude ratios of SV/P waves measured at Hilo showed that a systematic difference of 10 in fault plane orientation existed during years before the 1975 main shock. This agrees with the rotation of the fault plane during the main rupture reported elsewhere. We hypothesize that heterogeneities along the fault plane with dimensions of about 5 km cause changes in fault plane dip, differences in average stress release per earthquake, strong contrasts in seismicity patterns, and differences in aseismic strain release by fault creep. These same heterogeneities in the fault plane apparently also control the location of major earthquakes and the complexity of their ruptures. INTRODUCTION The ratio of high- to low-frequency amplitudes in seismic signals is a function of the path and the source properties. If the effects of the path can be eliminated, or if the path is held constant, then changes in the ratio of high- to low-frequency amplitudes between individual earthquakes will be due to changes in source properties and errors. If the influence of errors is not too large, one may be able to identify differences in the relative high-frequency content as due to differences in source properties. One might expect that small earthquakes could have differences in source characteristics because the Earth's crust is heterogeneous in rock properties and concentration of stress. The multiple nature of signals of large earthquakes shows that stress concentrations exist. In many of these cases, the continuity of surface ruptures and aftershock distribution, as well as the long-period seismic wave excitation, require the conclusion that the rupture was continuous over a large area; however, the multiple high-frequency pulses show that high-frequency energy came primarily from small portions of the rupture area (e.g., Wyss and Brune, 1968; 69

2 70 F. R. ZUI~IGA, M. WYSS, AND M. E. WILSON Trifunac and Brune, 1970; Wu and Kanamori, 1973; Fukao and Furumoto, 1975; Kanamori and Stewart, 1978; Hartzell and Heaton, 1983; Zhou et al., 1983; Harvey and Wyss, 1986). For this reason, many rupture models have been constructed in which crustal heterogeneity plays an important role (e.g., Madariaga, 1976; Das and Aki, 1977; Aki, 1979, 1984). Highly stressed crustal volumes are generally expected to be the source of pulses enriched in high-frequency energy. Therefore, the relative high-frequency content in signals of small earthquakes may be a function of location. The ambient stress along faults varies with time. During the loading cycle for great earthquakes, it increases slowly, and during great ruptures, it decreases rapidly. Thus, it is reasonable to postulate that the high-frequency content of small earthquakes might vary as a function of time in addition to its variation with space. Several authors have therefore searched for changes in relative high-frequency content of seismic signals as a function of time and space {e.g., Wyss and Brune, 1968; Fedotov et al., 1972; Tsujiura, 1977; Archambeau, 1978; Zhu et al., 1978; Bakun and McEvilly, 1979, 1981; Ishida and Kanamori, 1980; Scherbaum and Kisslinger, 1984). The purpose of this paper is to test two hypotheses: (1) the stress tensor in the source region of a large earthquake can be a function of space during many years, causing differences in the relative high-frequency content of small earthquakes as a function of space and that (2) changes of ambient stress level or changes of fault properties may cause differences in the relative high-frequency content of small earthquakes as a function of time. The area of study is the source volume of the 1975 Hawaii Ms = 7.2 main shock. Independent evidence such as precursory seismicity patterns and strong motion characteristics shows clearly that the stress in this volume was distributed heterogeneously in space (Wyss et al., 1981; Harvey and Wyss, 1982) and that some crustal properties changed as a function of time (Wyss et al., 1981). The seismicity rate of this area is high, such that numerous signals could be analyzed, and the hypocenter location accuracy is excellent (about _+1 km), thanks to the dense seismograph network of the Hawaiian Volcano Observatory (HVO). Many questions have been raised on how to measure the relative high-frequency radiation of small earthquakes. The first requirement is that the effects of the paths are small or can demonstrably be corrected for. Second, one needs a parameter which expresses the ratio of high- to low-frequency energy radiated. In this study, we have characterized the S-wave signals by two parameters: the apparent stress 775 (Aki, 1966; Wyss and Brune, 1968) and the Brune (1970, 1971) stress drop Aa. We chose to use two parameters because both have advantages and disadvantages. The apparent stress is proportional to the integrated energy in the signal, normalized by the long-period displacement spectral density. Its advantage is that it does not depend on the choice of a particular corner or peak frequency; however, errors contained in the high-frequency part of the spectrum may affect it. The Brune stress drop is proportional to fc 3 (corner frequency of the displacement spectrum) multiplied by the long-period spectral amplitude level. Here, a choice has to be made for a particular fc, but errors which may be present in the high-frequency part of the spectrum do not influence the measurement. Recent discussions have clarified the usefulness of various stress estimates (Boatwright, 1984). The apparent stress was found to be a stable estimate of stress release, while the Brune stress drop was not well-correlated with other estimates. It is important to realize that, in this paper, we are not claiming to provide an accurate measurement of stress release 75 or Aa. Instead, the purpose is to

3 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 71 demonstrate the existence or absence of variations in the relative high-frequency content of the S-wave signals radiated and investigate its probable correlatien with ambient stress. Whether such differences are caused by differences in the static or dynamic stress drop or in the complexity of the ruptures or other mechanisms cannot be answered by the data available. Although the chief concern is the ratio of high- to low-frequency content, we express this ratio in terms of apparent stress and stress drop because these quantities are well-known and clearly defined. The setup of the experiment is as follows. The single seismograph station for which digitizable records are available for the entire period of interest is Hilo (HIL), which is located at approximately 45 km to the north of the source volume of the 1975 Hawaii Ms = 7.2 main shock. Within an approximately 30 km (east-west) segment of the aftershock zone (Figure 1), there exist three subvolumes (radii range from 2.5 to 5 km) for which contrasting properties have been proposed (Wyss et al., 1981). First, the hypothesis is tested that the high-frequency content for signals emanated by earthquakes located in volumes 5 and 2 are larger compared to those in volume 4. {These volume numbers were defined by Wyss et al., 1981.) If this is the case, we must attempt to demonstrate that the signal contrasts are not caused by different path properties. This will be done by analyzing the signals of a limited number of recent events from the same source volumes, but recorded digitally at station Ahua (AHU, Figure 1) located 10 to 20 km from the sources in a different azimuth. Since rays from volume 4 to AHU appear to sample portions of both volumes 5 and 4, and both appear to travel at the same depth mostly along a horizontal path, similar observations in both data sets will help us to reject gross path effects. The second question addressed with the same data set is that of possible variation of frequency content as a function of time. Such a change might be expected in the 9 i i i i i i i LEGEND 6.0 bor bor //// ~f bor ~.~.'Z- - - / ~ ~ o./////~.-i ~ o ',, 0".//4/>.9 /.," o~~. ~ 4o.Y-", 10 km, I I I I I I I FIG. 1. The south flank of Kilauea volcano, island of Hawaii. Stress drops of earthquakes which occurred between 1962 and 1981, analyzed in this study, are plotted at their epicentral locations. The epicenters of the largest foreshock (ML = 5.9) and the 1975 main shock are plotted as a small and large star, respectively. The zone of early aftershocks is outlined by short dashes. Long dashes indicate the boundaries for the regions mentioned in the text, with numbers identifying them. Seismograph stations used in this study are shown as triangles. The South West Rift Zone (SWRZ) and East Rift Zone (ERZ) are also indicated.

4 72 F.R. ZUI~IGA, M. WYSS, AND M. E. WILSON whole data set, or in subvolumes, and it might be related to the stress release of the 1975 large main shock. Path effects are not a major concern in this experiment because the paths remain approximately the same with time. A third experiment was aimed at estimating the stability of the fault plane orientation or the principal stress directions. The ratio of S- to P-wave amplitudes at HIL could provide the data base for this test because the departure direction of rays to HIL are close to the fault plane of most earthquakes in the area, and therefore the amplitude ratio at that station is sensitive to small changes in fault plane dip. Several authors examined the ratio of compressional wave amplitudes to shear wave amplitudes generated by earthquakes. Engdahl and Kisslinger (1977) studied SV/P amplitude ratios of a group of earthquakes recorded at a station on Adak Island, Alaska. They noted that the value of SV/P changed at this station for four events preceding a magnitude 5 earthquake. Lindh et al. (1978) found that mean nodal plane orientations were different for foreshocks and aftershocks of three moderate California main shocks. In these cases, the difference between foreshock and aftershock fault plane orientations was small (5 to 10 ), approximately of the same size as the changes found in this study. The stress drops and apparent stresses of about 300 earthquakes (2.8 _-_ ML <-<- 3.9) which occurred over a period of 13 yr preceding and 6 yr following the 1975 Hawaii earthquake were estimated from records of Hilo. In addition, 78 signals recorded digitally at Ahua during January 1980 were analyzed also. S V/P amplitude ratios of about 54 earthquakes were measured at station Hilo. THE 1975 HAWAII EARTHQUAKE On 29 November 1975, a large earthquake (Ms = 7.2) occurred on the south flank of the volcano Kilauea after a sequence of foreshocks (Ando, 1979). Southward displacements of up to 10 m and subsidence of up to 3.4 m occurred along the south coast of Hawaii during the earthquake (Tilling et al., 1976; Ando, 1979). The nearhorizontal aftershock zone had dimensions of km (Figure 1). Ando (1979) concluded that aftershock distribution, tsunami, and crustal deformation data best fit a mechanism with strike N70 E, dip 20 SSE, slip pure normal, moment dyne-cm, and stress drop 43 to 93 bars. The main slip zone was located at about 8- to 10-kin depth, perhaps along the boundary between the buried seafloor and overlying lava flows (Furumoto and Kovach, 1979; Ando, 1979). Results of geodetic surveys conducted throughout the twentieth century indicate that NNW-SSE compression has accumulated in the south flank of Kilauea during eruptive activity. This compressive stress is possibly released about every 100 yr by large earthquakes like the 1868 and 1975 Hawaii events (Swanson et al., 1976). Various velocity models of the south flank are discussed by Crosson and Endo (1981). They conclude that solvable and consistent focal mechanism solutions are obtained by either including a low-velocity zone or thick crust in the model. Klein (1981) used a model with linear gradients that achieves similar results. Crosson and Endo (1981) plotted slip directions of a group of earthquakes which occurred in 1974 and noted that fault slip directions were consistently perpendicular to the east rift zone. Johnston et al. (1982) concluded that a precursory P-wave delay existed for nearly 4 yr in the volume of the foreshocks near the main hypocenter, while no velocity change occurred below five other seismographs located on the source volume. Precursory seismicity patterns also showed heterogeneity. Nearly 4 yr of seismic

5 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 73 quiescence existed in parts of the rupture volume before November 1975, while the foreshock volume and another segment of the rupture continued to produce small earthquakes at a nearly constant rate (Wyss et al., 1981). The signals of the main shock recorded by two strong motion seismographs clearly showed that the main shock consisted of several shocks somewhat larger in size than the ML = 5.9 foreshock (Rojahn and Morrill, 1977; Harvey and Wyss, 1982). The physical processes leading to the 1975 Hawaii earthquake and the main rupture are clearly complex functions of space. The ambient stress level and conditions on the fault plane may have controlled the heterogeneity of these processes. The aim of this study is to throw some light on the state of stress and its changes in the source area by estimating apparent stresses and stress drops for as many earthquakes as possible from that area. DATA SH-wave spectra. Seismograms of earthquakes which occurred within the source volume of the 1975 main shock during the years 1962 to 1981 were obtained from the two components of the Wood-Anderson torsion seismographs located at Hilo (Figure 1). All signals which were on scale, readable, and had an acceptable signalto-noise ratio were photographed and enlarged by a factor of 6 (some signals were enlarged by factors of 2 or 4, respectively). No signals which could be analyzed were omitted. The sizes of the earthquakes which were recorded on scale were 2.8 <_- ML --< 3.9. Thus, our sample of approximately 300 events (Table 1) consists of all earthquakes in this magnitude range which occurred in the area during the years 1962 to 1981, minus those events which were not recorded (or not recorded clearly) due to instrumental problems, and those events which were so rich in high frequencies that the trace could not be clearly seen, even when magnified. Of the 86 highfrequency events excluded, 33 were located in region 5, 33 in region 4, and 20 in region 2. Including all available signals, this translates into 23 per cent of exclusions for region 2, 19 per cent for region 4, and 28 per cent for region 5. This means that apparently more high-frequency events were located in regions 5 and 2 compared to region 4, and thus the average apparent stress (stress drop) in regions 5 and 2 may be underestimated by our sample. The hand-digitization of the signals must have introduced errors into the analyzed time series, and thus into the spectra and the results. These problems have been faced previously by other investigators (e.g., Hanks and Thatcher, 1972; Ishida and Kanamori, 1980). In this study, four individuals digitized randomly chosen signals. Two of these workers paid special attention to high-frequency signals, and signals which were not digitized correctly the first time. If this digitization process introduced an average bias in all spectral estimates, our analysis will not be affected because we are not aiming at estimating absolute stress release values but only at evaluating the possibility of differences in relative high-frequency content of the signals studied here and its connection with ambient stress. If the digitization process randomly introduces errors into the spectra of all events (regardless of origin time or hypocenter of the event), we are still able to make meaningful comparisons between the average parameters of events which were grouped as a function of space or time. There is only one parameter which may influence systematically the errors introduced by digitizing. This is the high-frequency content of the signal. Since it is more difficult to see and follow the traces of high-frequency signals, it is more likely that errors are made. It is not known whether these errors would artificially reduce or enhance the relative high-frequency content of signals

6 74 F.R. ZUI~IGA, M. WYSS, AND M. E. WILSON TABLE 1 STRESS DROPS AND APPARENT STRESSES FOR EARTHQUAKES IN THE HAWAII 1975 MAIN SHOCK SOURCE VOLUME BASED ON HILO RECORDS Mo Date Hr ML R* Latitude Longitude Depth (dyne*cm F (yr m d) ( ) ( ) (kin) 102o ) (Hz) Drop Stress (bars) Apparent {

7 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 75 TABLE 1--Continued Date Latitude Longitude Depth Mo (yr m d) Hr M~ R* (dyne-cm ( ) ( ) (km) x 10 2 ) , , , , F, Stress (bare) (Hz) Drop , , , , , , , , ,21 3, , , , Apparent ,

8 76 F. R. ZUI~IIGA, M. WYSS, AND M. E. WILSON TABLE 1--Continued Da~ (yrmd) Hr O O O Latitude Longitude Depth Mo ML R* (dyne-cm ( ) ( ) (km) x 102 ) , l , F Stress (bars) (Hz) Drop Apparent

9 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK TABLE 1--Continued 77 Date (yr m d) Mo Latitude Longitude Depth Hr ML R* (.) ( ) (km) (dyne-era 10 ~) , , , , , , , , , , , , , , , , , , , t , , , , , , , Fc (Hz) Drop Stress (bars) Apparent , ,6 4, , , , , ~ ,

10 78 F. R. ZUlg/IGA, M. WYSS, AND M. E. WILSON TABLE 1--Continued Date (yr m d) Hr O ll Mo ML R* Latitude Longitude Depth (dyne-cm ( ) (0) (km) 102o ) Fc (Hz) Drop Stress (bars) Apparent , , , , , , , , , , , , , ~ , , , , , , ,

11 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 79 TABLE 1--Continued Mo Stress (bars) Date Hr ML R* Latitude Longitude Depth (dyne-cm Fc (yr in d) ( ) ( ) (km) (Hz) x 10 z ) Drop Apparent , , ,224 10, , , , , , , , , , * This represents the subvolume number (see text). difficult to digitize. However, if groups of shocks which we wish to compare radiate the same average frequency content, then any errors will affect them in the same way, and we are bound to find no difference in their average apparent stress. On the other hand if signals are different in the average, the error may enhance or reduce the difference. Therefore, if we do find a difference in the spectral content, we cannot be sure that its amount is estimated correctly, but we can be confident that the difference exists. Digital seismographs sample the incoming signal at even intervals, T, located randomly with respect to the signal's peak amplitudes. In such a case, it is usually assumed that, in the spectral analysis, the amplitudes of frequencies can be esti- mated reasonably well up to f = ~. In the case of hand-digitization, the signal is already known to the observer when it is being digitized, and the observer will

12 80 F. R. ZUI~IIGA, M. WYSS, AND M. E. WILSON carefully place digitization points at extreme values (peaks and troughs) as well as at inflections in between. In hand-digitizing a relatively high-frequency, highamplitude signal, it is not advisable to digitize at the highest possible rate because one introduces errors along nearly straight near-vertical segments of the trace. Instead, it is better to digitize extreme values by one point each and inflections by three points (beginning, center, and end of inflection). Thus, our sampling rate varied through each individual signal. Within the entire 15-sec-long S-wave digitization, the average sampling rate ended up to be 8 to 10 samples/sec. Within the S-window used for analysis, the sampling rate was higher, but did not exceed 24 samples/sec (a six-fold magnification of the Wood-Anderson records yields ~ sec = 1 mm, which is the smallest distance we could separate for digitization). We thus believe that the spectral amplitudes of the signals analyzed can be accepted up to 12 cycles/sec. The stability of the high-frequency fall-off of the observed spectra corroborates this assessment. The average corner frequency (and hence the peak power spectral frequency) is [c = cycles/sec. Only 1 per cent of the spectra have an fc > 5.4 cps (Table 1); thus, we could define and measure fc in our data. The hand-digitized time series were interpolated by a half-cosine and cubic (at inflections) interpolation routine with a sample interval of sec. In addition, spectra were also estimated for about 100 signals using linear interpolation. Some of the corner frequency estimates changed by small amounts, some remained the same. The apparent stresses and stress drops calculated in this paper were obtained after interpolation by the cosine-cubic method, but all basic conclusions remain the same if a linear interpolation is used. We conclude that the apparent stress and Brune stress drop can be estimated from the data at hand which are valid to 12 cps. Spectra were obtained from the first 3 sec of S-wave arrivals recorded on the east-west component seismograph (see Appendix, Figure A1). Since 95 per cent of the source to receiver azimuths are within _10 from the north, the east-west component recorded at Hilo represents true SH-wave motion with an error in displacement amplitude no greater than 10 per cent. Assuming that the structure of southern Hawaii consists of horizontal layers, the east-west component S-wave signals should not be significantly contaminated by P-wave-type motion. After applying a cosine taper to the SH phase, a Fast Fourier transform routine was used to obtain the spectra (Appendix, Figure A2). The raw spectra were corrected for a damped harmonic instrument response of period 1.0 sec, a damping coefficient of 0.45, and a static magnification of 2790, appropriate for station Hilo. The damping coefficient was determined from calibration curve amplitudes measured in 1980 and is assumed to have been constant during the time period 1962 to Spectral amplitudes were corrected for attenuation in a homogeneous medium according to the relation o exo[v] where t is the travel time from source to receiver, and fm is the frequency of the sample in Hertz and Um its displacement amplitude. The attenuation model used in this corection was obtained by Chouet (1976) from the measurements of seismic coda close to the summit of Kilauea. The Q values given are 260, 290, 270, and 520 for frequencies of 1.5, 3.0, 6.0, and 12.0 Hz, respectively.

13 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 81 If the Q for the Hawaii structure is in fact more nearly constant over the bandwidth studied, then the corner frequencies could be shifted to lower values by the attenuation correction. We calculated apparent stresses and stress drops for a model with Q = 275 for all frequencies as well as for the Chouet {1976) model. The trends in these parameters are the same for both attenuation models, but the number of high-quality data is reduced for constant Q, and the trends have lower statistical significance. Since Chouet's model was obtained for the local structure in question, we believe that the data analysis should be based on it. The difficulty of estimating fc does not appear to be a serious problem. We have found that subjective judgments of several observers and estimates by computer fit usually agree on the value of [c within about 30 per cent, with the exception of some unacceptable estimates by the computer algorithm. Thus, the appropriate asymptotic fits of straight lines to low- and high-frequency parts of the spectra (Figure A2) were done subjectively by one observer for all spectra as consistently as possible. Although some errors are introduced here in the estimate of Aa, these errors should be random and they should approximately cancel out in the average. The seismic moment for a double-couple source is calculated from the lowfrequency amplitude ~2c using the equation (e.g., Aki and Richards, 1980) 47rpj33R Mo = - - ec 2FsH where ~ is the wave velocity near the source, R is the source to receiver ray distance, p is the average density of the crust, and FSH is the radiation pattern coefficient for SH waves. From the corner frequency and moment, the fault radius r and stress drop can be estimated (Brune, 1970) r = 0.37 fc Aa = ~ Mor -3. The distance R, take-off, and emergence angles from source to receiver are calculated from ray tracing and event location programs using the velocity model HG50 (Klein, 1981) and assuming a = J3/~. The average density of the south flank crust was assumed to be 2.9 gm/cm a. The radiation pattern of SH waves (FsH) was calculated from the strike, dip, and slip of focal mechanisms determined from first motion data recorded by the HVO seismic array. Only a small fraction of all earthquakes had adequate records for determination of a focal mechanism. Earthquakes without focal solutions were assumed to have mechanisms equivalent to the 1975 Hawaii main shock. This is justifiable because most of the focal mechanisms determined in this study have a low-angle fault plane solution, consistent with the main shock mechanism. And, in addition, the compression and tension axes of earthquakes in 1974 and 1975 clustered about the main shock principal axes (Klein, 1981), and slip directions were consistently perpendicular to the east rift zone, the direction of slip during the main shock (Crosson and Endo, 1981). The spectra were obtained from 3-sec time windows to reduce variation of the results by multiple refracted waves and surface waves arriving later in the coda. To estimate source parameters more accurately, it would be desirable to match synthetic

14 82 F. R. ZUI~IIGA, M. WYSS, AND M. E. WILSON signals to the observed ones. However, this would require a larger effort and would probably not yield a significant improvement because the source-receiver rays sample a fairly small portion of the crust. We therefore feel justified in assuming that contaminations of source spectra due to the structure remain approximately constant throughout our data set. The radiated seismic energy was estimated by Es = 1 + ~ f~(t) 2 dt Fsn oo where q is the ratio of S- to P-wave energy, assumed to be q = 20. From this value and the seismic moment, the apparent stress n5 is obtained by E~ n5= U Mo The shear modulus # is assumed to be # = N/m 2. The apparent stress is the product of the two unknowns, 7, the seismic efficiency, and ~, the average shear stress on the fault. Apparent stress and stress drop. One of the main purposes of this study was to test the hypothesis that some subvolumes of the 1975 main shock might produce earthquakes with relatively more high-frequency content in their signals. The volumes separated originally on the basis of seismicity patterns were grouped into "soft-volumes" numbered 1, 3, and 4 (Wyss et al., 1981) and "hard-volumes" 2 and 5. In the present study, we did not have enough events to examine the relative highfrequency content separately in all these volumes, thus the few events located in volumes 1 and 3 were merged with volume 4, except for those events which were located in the original volume 3 but along the boundary of volume 2. These events were merged with the data from volume 2. Also, a few events available from the aftershock area west of the original volume 5 (an area not analyzed by Wyss et al., 1981) were added to the volume 5 data (Figure 1). The average apparent stress and stress drop are examined separately for volumes 5, 4, and 2 in Figures 2, 3, and 4, respectively, which show the cumulative stress parameter (apparent stress and stress drop) as a function of event number. In this presentation of the data, the abscissa represents progressing time but the cumulative stress parameter is plotted at equal intervals for each subsequent event. This representation is chosen because the slope of the resulting curves is proportional to the average high- to low-frequency ratio per event, which is the parameter we are interested in. In Figures 2 to 4, the same scale applies for apparent stress and stress drop. In all three figures, the cumulative curves of apparent stress and stress drop track each other closely, which shows graphically that, in this experiment, the apparent stress equals approximately the Brune stress drop on the average and that the two parameters correlate fairly well. Fitting a linear regression through all data (Table 1), we find that ~ = 1.1Aa + 1, with a correlation coefficient of This correlation is better than that found for the data set analyzed in detail by Boatwright (1984). We do not attach general importance to this result, but we conclude that, for our specific experiment, the equivalence of ~5 and ha means that we have two valid estimates proportional to stress release which agree with each other. This

15 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK I I I I( I I! _ ASPERITY //r (/3 Q- Ld ry ry P- r~ - NOV 75 - (/3,.-- LO /J L~ OK - $ c' -- ry.< br) Q_ Q- o <' ~ /J/ 1962 to ODF-:~D<' 0 Yl'ggtlY'l" `781` 79 1" 81 1" 0 EVENT NUMBER (TIME) 140 FIG. 2. Cumulative stress release measured at station HIL as a function of event number, for events in volume 5. The solid line is the cumulative stress drop, and the broken line is the cumulative apparent stress. Since all events are plotted evenly spaced, the slope of the curve is equal to the average stress release per event. The single arrow marks the time of occurrence of the 1975 main shock. Arrows along the horizontal axis define periods of 1 yr i i i i i i o_ O c~- 6'3 0"3 tad ry t--- (,q SOFT REGION - NOV 75 O0 6O Ld Y 09 m t-- 3 C) / 963 to 1961 O 4 ~" 1~' 1~ 75 1` 76 1" 77 q' 78 1" 791" 81 0 EVENT NUMBER (TIME) 40 FIG. 3. Same as Figure 2 for events in volume 4. conclusion is based on the assumption that errors in estimating these two parameters are independent. We will use the expression "stress release" in the discussion of the results because all conclusions are equally valid whether they are based on ~5 or A~r. The cumulative stress release curves in Figure 2 are steeper than the curves in Figure 3, suggesting that the average stress release per event in volume 5 is approximately twice that in volume 4 over the entire study period. Before this can be accepted as a compelling result, one needs to examine the errors of the stress

16 84 F.R. ZUNIGA, M. WYSS, AND M. E. WILSON I I I I I I - /,7 FORESHOCK VOLUME ~// a_ /// ~z / / - F~,., ry jr.,-...*'"....- > NOV 75./'.,z" rfj < f_e i**** ',D M" ~ ( 1964 to u O]/.~rt~,~l'ttt ~ 1" 771" 7a 7~ s~ 0 EVENT NUMBER (TIME) 66 FIG. 4. Same as Figures 2 and 3 for events in volume 2. A line representing the average slope of volume 4 {short dashes) is added for comparison. release estimates and the possibility that differences in the path might cause differences in the signals recorded. We may compare the statistical significance of the difference between two means M1 and M2 which are known with standard deviations of $1 and $2 by the standard deviate z test Z ~ M1 - M2 "~/ n2 where nl and n2 are the respective number of samples. We will assume, for the moment, that the average stress release per event did not change as a function of time in volumes 5 and 4. Then the average apparent stress for earthquakes in volume 5 is ~-~(5) = 9.01 bars (n5 = 88) and in volume 4 is ~-~(4) = 4.16 bars (n4 = 139). Thus, we find that z = 4.55 which means that ~-~(5) ~ ~-~(4) at a confidence level exceeding 99.9 per cent. However, it may be argued that the error estimates for the two means are not correctly representing the true errors because the errors equal approximately the apparent mean stress values themselves. So we asked the question: for what size of real errors S(real) would the observed difference in the means AM = M1 - M2 = 4.85 bars still be significant at a level of 99 per cent. We found that would be the case if S(real) ~ 1.7 S(observed). If it was assumed that S(real) = 2 S(observed), the confidence level is 98 per cent. Therefore, we conclude that a difference does exist between the seismic signals emanated from volume 5 compared to volume 4, with the volume 5 signals showing relatively enriched high-frequency contents. Alternatively, one may estimate the difference in frequency content by Afc = fc (5) - fc(4), the difference in corner frequency or peak frequency as done by Ishida and Kanamori (1980). The disadvantage of this approach is that f~, the parameter used,

17 m EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 85 will not be normalized for event size. One might argue that in a first approximation this may not be necessary because the magnitudes are approximately constant on the average (ML = 3.3 _+ 0.5). We find that [~(5) = _ 1.00 cps (n5 = 88) and /,.(4) = 2.73 _ 0.94 cps (n, = 139). In this case, the standard deviations are smaller than the values themselves, and they can therefore be accepted as valid estimates of the real distribution of f~. By the standard deviate z test, we thus find that f~(5) >/~(4), with a confidence level exceeding 99.9 per cent. The difference between the average stress release from volumes 5 and 4 are also clearly evident in Figures 2 and 3. The variance of the cumulative stress release curves are small compared to their pronounced difference in slope. We propose that the observed difference between the data sets from volumes 5 and 4 is significant. If absorption of high-frequency waves is larger along the path between volume 4 and volume 5, the observed signal difference might not be due to differences in the radiated signals. If a difference in attenuation existed between the two paths (volume 5 to HIL and volume 4 to HIL), it would be more likely that the paths closer to the volcano summit (volume 5 to HIL) contained larger absorption, and hence one would rather expect that the frequency content of signals from volume 5 would be reduced compared to those of volume 4. An additional experiment was carried out to shed light on the possible role of the paths in the question of spectral differences. For the years after 1979, digital data are available for many small seismic signals. We thus obtained spectra of recent earthquakes from volumes 5 and 4 derived from records of station AHU which is located in an azimuth which differs by 90 from that to HIL (Figure 1). The average magnitude was ML = 1.5 for these recent events recorded by the U.S. Geological Survey short-period local seismograph at AHU, since signals for earthquakes with ML > 2.0 saturate the recording system. The spectra were obtained from the transverse component of the signal which was derived from the NS and EW components of station AHU. The results again show that, on the average, the apparent stresses and stress drops calculated for volume 5 earthquakes are twice as large as those for volume 4 earthquakes (Figure 5). Ray tracing using the HG50 8O ce C] 0 I I I I 80 ~9 :D -d Q_ a:: O b I I I SOFT REGION 09 LIJ m LG L/~ r~ Ld <~: EL F-- EL LCJ <:~ I-d w >- Ld a_ G_ j ~J D ~ (D LD ~ : n 1980 (..9 0 I I I I 0 -~-~ 0 EVENT NUMBER (TIME) 50 0 EVENT NUMBER (TIME) 50 Fro. 5. Cumulative stress release versus event number as measured at station AHU. All earthquakes occurred in January (a) Events in volume 5. (b) Events in volume 4. Other features are the same as in prior figures.

18 86 F.R. ZUNIGA, M. WYSS, AND M. E. WILSON model indicates that, for the most part, the path from events in volume 4 to both HIL and AHU is almost horizontal. Rays from volume 5 also have an initial horizontal path. Therefore, since all hypocenters lie at approximately the same depth ( km) and the end of the path is shared by all rays, we think that these observations indicate that the rays sample mostly the same layer. Although the period of time covered by this experiment ( January 1980) is small, we take these results to show that it is not probable that the observed spectral differences between signals from volumes 5 and 4 (Figures 2, 3, and 5) were caused by differences along the paths. Instead, we propose that earthquakes in volume 5 generate more high-frequency energy because of fundamental differences in source properties, which seemed to have existed over the entire observation period of nearly two decades. The average stress release per event in volume 2 was expected to be similar to that of volume 5 because of the similarity in seismicity patterns. However, the data show a mixture between similarities to both volume 5 and volume 4 slopes (Figure 4). The first 10 and last 35 consecutive events show an average apparent stress of approximately 7 bars similar to the volume 5 average, but around the time of the 1975 main shock, 23 consecutive events give an average apparent stress of 2.5 bars (less than the volume 4 average). This suggests that the signal properties from volume 2 (the foreshock volume) might have changed as a function of time. If we would assume that the observed standard deviations estimate correctly the error of the means, we calculate a confidence level of 99 per cent. However, it is more reasonable to assume that the real standard deviations are equal to or larger than approximately 10 bars (not 2.5 bars as observed). Under this assumption, the two samples are different at the 92 per cent confidence level only. Thus, we believe that it is not clearly resolved whether or not the average stress release per event changed as a function of time in volume 2. The possibility of a temporal change of radiation characteristics was a prime source of motivation for this study. However, there exists no clear change of the spectral characteristics at the time of the main shock when the ambient stress was reduced substantially. In November 1975, none of the slopes in Figures 2, 3, and 4 change, clearly rejecting the hypothesis that the main shock may have influenced the average stress release per event by reducing the ambient stress in the entire aftershock area. At the beginning of each of the cumulative stress release curves (Figures 2, 3, and 4), the slopes are somewhat higher than average. We joined all data and tested the significance of the difference between the 39 first data points (1962 through November 1970) and the rest of the data. The average apparent stress decreased from 11.6 bars to 5.3 bars, with standard deviations of similar size. With the observed standard deviations, the confidence level of the change would be 99 per cent; however, if we assume that the real standard deviation should be about double the observed one, we find a confidence level of 96 per cent. Thus, the data suggest that before 1971 the average stress release per event may have been higher than afterwards. Amplitude ratio of SV/P waves. Consistency of focal mechanisms with the main shock source can be tested by measurement of radial component S V to P amplitude ratios. These ratios measured at Hilo were taken from the north-south component records which represent the true radial SV and P motion with probably less than 10 per cent error. The amplitudes of the first cycle of each SV and P phase were

19 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 87 chosen to avoid contamination by later arrivals. The P-wave travel-time curve for a source at a depth of 10 km shows that two phases arrive at Hilo for epicentral distances greater than 35 km (Wilson, 1981}. In addition, if the crust mantle [,oundary transition is rapid enough, refracted waves could arrive just after the P and S phases. Since at least one phase usually arrives i to ¼ sec after the P and S phases for earthquakes recorded at Hilo, the amplitude ratios must be picked carefully to obtain accurate results. Observed amplitude ratios of 54 earthquakes recorded at Hilo (triangles) and theoretical main shock SV/P ratios (circles) are plotted in Figure 6. Theoretical SV/P ratios are calculated by placing a main shock fault mechanism (Ando, 1979) at each south flank earthquake hypocenter, and they are compared with observed SV/P ratios. The value of SV/P ratios recorded at Hilo is strongly dependent on fault dip for earthquakes, with slip directions perpendicular to the rift zones. For example, if the fault mechanism varies between dips of 10 and 45, the SV/P ratios vary between 55 and 0. However, the fault azimuth can vary as much as 30 from due east without affecting the amplitude ratios significantly. For events in regions 2 and 5, the observed ratios are primarily less than 18 which agrees with the theoretical ratios generated by a main shock fault mechanism with dip = 20 SSE. Earthquakes located in region 4 have SV/P ratios mostly larger than 18, suggesting that these faults have a shallower dip of 10 to 15. The magnitude of the ratios are remarkably constant over time which may indicate that the earthquake mechanisms are also fairly constant over time. The main shock mechanism then is probably the best average representation of the fault geometry of south flank earthquakes with undeterminable focal mechanisms. Instead of finding a temporal variation in the SV/P-wave amplitudes, we found a spatial one. Earthquakes located in region 4 again have different properties than those in regions 2 and 5. During the main shock rupture, the fault plane strike and dip also changed as the rupture progressed from region 2 through regions 4 and 5. Two strong motion signals recorded at Hilo and Panaluu (near the SW end of the main shock rupture} could not be modeled satisfactorily unless the dip of the fault plane was varied between 0 and 30 (Harvey and Wyss, 1986). DISCUSSION The search for precursory changes of high-frequency S-wave signal content and ratio of SV/P waves was negative in the source volume of the 1975 Hawaii Ms = 7.2 main shock. Amplitude ratios of about 50 events could be obtained for the 7 yr before the main shock. These data did not show a change with time. Instead, they showed a spatial variation by about 10 of the dip of the fault planes of small earthquakes (ML ~ 3.3 _+ 0.5). This result agrees with the observation that the fault plane of the main shock also varied by approximately 30 along the rupture, first in one direction, and then back again (Harvey and Wyss, 1986). Both of these observations show that small structural differences exist within the source volume of the 1975 main shock on a scale of 5 to 10 kin. The regionalization of the source volume based on the precursory seismicity pattern (Wyss et al., 1981) was such that it separated the shocks with the anomalous S V/P-wave ratios from those with normal ones. This is a significant result, because it implies that the structural difference which gives rise to different dips of the fault planes also controls the contrasts in seismicity patterns.

20 88 F. R. ZUNIGA, M. WYSS, AND M. E. WILSON REGION 5 0 i 20- & 10 A oo o 6 oo li o 2 c~ A& I e' I I I A! REGION 4 I I I I I I I Q.. :> Kn ~o~ ,~' o o o o o oo0 o- I I I I I I I REGION 2 I I I I I I I I! 40 m o I I I I i I I time (years) FIG. 6. S V/P amplitude ratios of earthquakes located in volumes 2, 4, and 5 recorded at station HIL. Triangles represent observed amplitude ratios, and open circles represent theoretical amplitude ratios generated from an equivalent main shock source placed at each earthquake hypocenter. The spatial variations in fault plane orientation found here and by Harvey and Wyss (1986) is not surprising because the surface expressions of well-mapped faults like those of the San Andreas system also change strike on a scale of 10 km (e.g., Bilham and Williams, 1985). These variations of fault plane orientation pose a

21 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 89 problem if one attempts to identify precursory rotations of focal mechanisms: if different segments of the fault are active at different times, the wrong impression of a temporal pattern may emerge. A time dependence of the relative high-frequency content of seismic signals could also not be demonstrated conclusively. There is no change of the average stress release per earthquake noticeable at the time of the main shock in any of the data sets. At this time, a reduction in ambient stress exceeding 100 bars must have taken place based on co-seismic surface deformations (Tilling et al., 1976). On the average, the earthquakes before the main shock have larger stress release than those after it, but all the earthquakes with high stress release occurred before 1971, which is a period poorly represented by the data. Thus, we feel that this apparent change may not be real (the confidence level based on the z test is 96 per cent if the real standard deviations are assumed to be twice the observed ones). The apparent temporal changes of average stress release seen in volume 2 (Figure 4) are significant at the 92 per cent level (assuming standard deviations are twice the observed ones). Thus, again these changes cannot be accepted as established. We conclude that no precursory changes of average stress release could be demonstrated, although the data suggest a reduction in late 1970, just at the time when other precursors appeared (Wyss et al., 1981). At some seismograph stations, local effects on recorded spectra are so strong that earthquakes and explosions have the same corner frequency (Frankel, 1981). We do not believe that such an effect caused the temporal constancy of the observed average stress release in the case studied here because a significant difference is observed in spectra recorded by the same station from sources in volumes 5 and 4 located at backazimuths which differ by less than 10. Thus, we reach the important conclusion that the average stress release per event (as measured by apparent stress or Brune stress drop) is not a measurable function of ambient stress in the volume studied. The most significant positive result is the identification of differences in average stress release per earthquake in small subvolumes of the 1975 main shock. Although the differences in slope of the curves in Figures 2 and 3 are quite pronounced, we performed a number of tests to see whether these data sets are really different. The window length for spectral analysis, the attenuation structure, and the method to define the corner frequency were varied. The resulting systematic and random changes in the apparent stress and stress drop values could be of importance if we wanted to base some conclusions on the comparison of the values obtained in this study area with values derived for earthquakes in another tectonic area, recorded by different instruments, and reduced by different methods. However, since we are only interested in the difference between the averages of volumes defined in the 1975 aftershock volume, we need not estimate the error of the absolute stress release values. For our purposes, it was more important to know that none of the method changes altered the trend of the data. The signals from volume 5 showed, on the average, a larger high-frequency content regardless of the data reduction technique. Also, the fact that most of the events (in percentage) rejected due to very high frequencies occurred for events from volume 5 helps support the hypothesis regarding that volume as prone to generate higher frequencies. Two methods of choosing the corner frequency were used: the Appendix contains examples of the subjectively chosen asymptotes to the long-period and short-period trends of the spectra. In addition, we also fitted such asymptotes by computer algorithm. The computer choice of the corner frequency would have been preferable

22 90 F.R. ZUI~IGA, M. WYSS, AND M. E. WILSON from the point of view of objectivity; however, some poor choices of corner frequency resulted because the assumed average high-frequency fall-off of f-2.5 was not always approximating the data well. Because of this shortcoming, we used the subjectively chosen corner frequencies in the final analysis. The average results based on the two methods were the same, and the apparent stresses, which are independent of a choice of corner frequency, confirm the stress drop results. Thus, we are sure that the choice of/~ did not critically influence the results. The average difference in the fault plane clips between volumes 5 and 4 could lead to a systematic error in the moment estimates. However, such errors could also not be the cause of the observed difference because: (a) the corner frequencies alone show a statistically highly significant difference between volume 5 and 4, and (b) the radiation pattern correction enters in different ways into the Aa and 715 estimates (an error leading to overestimation of one of the two would cause an underestimation of the other). The question of path effects is the most serious unknown. For this reason, we compared the spectra of recent small earthquakes recorded by a different station, and we found that signals from volume 5 contained more high-frequency energy than those from volume 4 in this data set also. This observation strongly suggests that differences in the paths are not likely to be the cause of the observed differences in the spectra, although we still cannot rule out that some structural peculiarity within volume 4 or close to it, caused the difference. The difference between 76(5) = 9.01 bars and 75(4) = 4.16 bars is statistically significant at the 99 per cent level even if the errors in both data sets are assumed to be about 14 bars, instead of the calculated standard deviations of 8.8 and 6.1 bars. Based on the above considerations, we propose that it is not probable that the observed signal differences between regions 5 and 4 were produced by random or systematic errors, and that the most likely explanation is that the earthquakes in volume 5 produce more high-frequency energy on the average. Volume 5, which showed the higher frequency content of earthquakes for the years 1962 through 1981, was also the volume which was proposed as an asperity volume (Wyss et al., 1981). The seismicity rate for earthquakes (ML ->- 1.7) within volume 4 decreased abruptly by about 50 per cent approximately 3.8 yr before the November 1975 main shock. The rates in volumes 5 and 2 remained at normal levels, while in volume 4 seismic quiescence lasted up to the time of the main shock. Volume 2 was defined by the hypocenter of foreshocks which occurred 70 min before the main rupture and by the main shock hypocenter. The location of foreshocks (for the largest ML ) and of the first two multiple events of the main shock in volume 2 (Harvey and Wyss, 1986) suggests that volume 2 was under high stresses at the time of the main shock. For this reason, it was proposed as an asperity volume (Wyss et al., 1981). Since volume 2 also showed constant seismicity rate, but was surrounded by precursory quiescence, it was hypothesized that subvolumes of large earthquake ruptures which are major asperities may be identified by this seismicity pattern. Thus, volume 5 was also proposed as a major asperity (Wyss et al., 1981). In addition to the seismicity pattern of constant rate within surrounding quiescence, volume 5 also accumulated compressive tectonic strain during a time when the strain was released or remained constant in the quiescent volumes (Wyss et al., 1981). A more recent main shock in Hawaii (1983, Ms = 6.6 Kaoiki) confirmed the hypothesis that major asperities can be identified by precursory seismicity patterns: the main shock hypocenter was located within a central volume of constant seismicity rate surrounded by a volume of seismic quiescence (Wyss, 1986). The

23 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 91 results of the present paper which show that small earthquakes in volume 5 have comparatively more high frequencies in their signals further support the idea that this is an asperity volume, within which comparatively high strength and high stresses exist at all times. We assume that volume 2 is also an asperity with high strength and stresses, although in the present analysis the data did not allow a conclusive classification of volume 2. The contrast between volumes may be caused by volcanic plugs which may cut through the oceanic sediments in volumes 5 and 2, or by geometrical discontinuities in the sediment layer, or by differences in pore pressure in the sediments. The higher strength leads to larger stress concentrations along the asperity segment of the fault and in the volume surrounding it, such that the ruptures of higher strength material will lead to seismic signals richer in high-frequency energy. The precursory seismic quiescence will take place only in the soft volumes within which precursory fault creep takes place. This model is not unique, although the independent data sets which constrain it are unusually numerous for a single earthquake. Also, it may be questioned whether this case history is pertinent for other tectonic areas. Although details of the tectonic strain accumulation will be different elsewhere, it seems evident that many of the major features of the 1975 Hawaii rupture are present in other large earthquakes: Multiplicity of the rupture is clearly indicated by complex signals, and precursory seismic quiescence exists elsewhere as well. The asperity model proposed by Wesson and Ellsworth (1973) for the San Andreas fault is the same model we have in mind for the 1975 Hawaii main shock and its precursors. Thus, we propose that the "hard patches" (Wesson and Ellsworth, 1973) along the San Andreas fault may be identified by mapping apparent stresses and stress drops of small earthquakes. Contrasts in average earthquake signal properties exist along the San Andreas system with high apparent stresses apparently present near fault bends (Wyss and Brune, 1971; O'Neill, 1984). More case histories are needed to establish or disprove the hypothesis that high average apparent stresses can identify fault asperities (Wyss and Brune, 1971), and the idea that precursory seismicity patterns together with apparent stress patterns could define likely locations for hypocenters of large earthquakes. The estimate of occurrence time of a future earthquake, however, can evidently not be improved by apparent stress studies because the average stress release seems not to be a function of time. In particular, the spectra of two foreshocks did not stand out as unusual, a result which agrees with the observations of Bakun and McEvilly (1979, 1981). CONCLUSIONS We conclude that variations in strength and stress were present in the source volume of the 1975 Hawaii Ms = 7.2 earthquake during years before and after the main shock. This conclusion is based on the contrast in the relative high-frequency content of small earthquake signals which occurred in the area during 1967 through The signals enriched in high-frequency content were emanated from a volume which was previously proposed as an asperity volume around a "hard patch" on the fault surface, and from which high-frequency signals were probably radiated during the main rupture. The lowering of the ambient stress by the main rupture did not affect the average stress release per small earthquake, as measured by apparent stress and Brune stress drop. This suggests that the higher average stress release in the asperity is due more to different material properties than a different stress level.

24 92 F. R. ZUI~IGA, M. WYSS, AND M. E. WILSON The apparent stress and stress drop data contained suggestions of changes with time; however, these changes were not significant enough to be accepted as real. It seems that temporal changes of spectral content may not be established easily if they exist, unless extremely numerous high-quality signals recorded digitally are available. Temporal changes of the amplitude ratios of SV/P waves may also be very difficult to identify because we found that the fault plane dips of small earthquakes change systematically by about 10 on a scale of 5 km. This change agrees with the inferred rotations of the fault plane during the main rupture (Harvey and Wyss, 1986). If more cases can be found where high average stress release per earthquake correlates with a seismicity pattern of normal rate surrounded by seismic quiescence, one may have a useful tool to identify likely rupture initiation points for future main shocks. ACKNOWLEDGMENTS We express our gratitude to Robert Koyanagi and Fred Klein of the Hawaiian Volcano Observatory, who supplied seismograms and earthquake locations used in this study. Carl Kisslinger pointed out some of the problems and techniques of measuring amplitude ratios and made many helpful comments to improve the manuscript. We also thank Steve Ihnen and Danny Harvey for developing computer programs useful to this research. Joyce Kruger assisted in digitizing some of the Wood-Anderson seismograms. We regret very much the untimely death of our co-author Mark Wilson, who died in a climbing accident. This research was supported by the National Science Foundation Grant EAR REFERENCES Aki, K. (1966). Generation and propagation of G waves from the Niigata earthquake of June 16, Estimation of earthquake moment, released energy and stress-strain drop from G-wave spectrum, Bull. Earthquake Res. Inst., Tokyo Univ. 44, Aki, K. (1979). Characterization of barriers on an earthquake fault, J. Geophys. Res. 84, Aki, K. (1984). Asperities, barriers, characteristic earthquakes and strong motion prediction, J. Geophys. Res. 89, Aki, K. and P. G. Richards (1980). Quantitative Seismology, Theory and Methods, W. H. Freeman and Co., San Francisco, California. Ando, M. (1979). The Hawaii earthquake of November 29, 1975: low angle normal fault due to forceful magma injection, J. Geophys. Res. 84, Archambeau, C. B. (1978). Estimation of non-hydrostatic stress in the earth by seismic methods: lithospheric stress levels along Pacific and Nazca plate subduction zones, in Proceedings of Conference II, Methodology for identifying seismic gaps and soon-to-break gaps, U.S. Geol. Surv., Open-File Rept Bakun, W. H. and T. V. McEvilly (1979). Are foreshocks distinctive? Evidence from the 1966 Parkfield and the 1975 Oroville, California sequences, Bull. Seism. Soc. Am. 69, Bakun, W. H. and T. V. McEvilly (1981). P-wave spectra for ML 5 foreshocks, aftershocks, and isolated earthquakes near Parkfield, California, Bull. Seism. Soc. Am. 71, Bilham, R. and P. Williams (1985). Sawtooth segmentation and deformation processes on the Southern San Andreas Fault, California, Geophys. Res. Letters 12, Boatwright, J. (1984). Seismic estimates of stress release, J. Geophys. Res. 89, Brune, J. N. (1970). Tectonic stress and the spectra of seismic shear waves from earthquakes, J. Geophys. Res. 75, Brune, J. N. (1971). Correction, J. Geophys. Res. 76, Chouet, B. (1976). Source, scattering and attenuation effects on high frequency seismic waves, Ph.D. Dissertation, Massachusetts Institute of Technology, Cambridge, Massachusetts. Crosson, R. S. and E. T. Endo (1981). Focal mechanisms of earthquakes related to the 29 November 1975 Kalapana, Hawaii earthquake: the effect of structure models, Bull. Seism. Soc. Am. 71, Das, S. and K. Aki (1977). Fault plane with barriers: a versatile earthquake model, J. Geophys. Res. 82,

25 EARTHQUAKE SIGNALS OF THE 1975 HAWAII MAIN SHOCK 93 Engdahl, E. R. and C. Kisslinger (1977). Seismological precursors to a magnitude 5 earthquake in the Aleutian Islands, J. Phys. Earth 25 (suppl.), Fedotov, S. A., A. A. Gusev, and S. A. Boldyrev (1972). Progress of earthquake prediction in Kamchatka, Tectonophysics 14, Frankel, A. (1981). Source parameters and scaling relationships of small earthquakes in the northeastern Caribbean, Bull. Seism. Soc. Am. 71, Fukao, Y. and M. Furumoto (1975). Foreshocks and multiple shocks of large earthquakes, Phys. Earth Planet. Interiors 7, Furumoto, A. S. and R. L. Kovach (1979). The Kalapana earthquake of November 29, 1975; an intraplate earthquake and its relation to geothermal processes, Phys. Earth Planet. Interiors 18, Hanks, T. C and W. Thatcher (1972). A graphical representation of seismic source parameters, J. Geophys. Res. 77, Hartzell, S. H. and T. H. Heaton (1983). Inversion of strong ground motion and the teleseismic waveform data for the fault rupture history of the 1979 Imperial Valley, California, Earthquake, Bull. Seism. Soc. Am. 73, Harvey, D. and M. Wyss {1982). A detailed source model for the multiple 1975 Hawaii rupture, M , Earthquake Notes 53, 56. Harvey, D. and M. Wyss {1986). Comparison of a complex rupture model with the precursor asperities of the 1975 Hawaii Ms = 7.2 earthquake, PAGEOPH (in press). Ishida, M. and H. Kanamori (1980). Temporal variation of seismicity and spectrum of small earthquakes preceding the 1952 Kern County, California earthquake, Bull. Seism. Soc. Am. 70, Johnston, A. C., M. Wyss, R. Koyanagi, and R. E. Habermann (1982). P-wave travel times: stability and change in the source volume of the M = 7.2 Hawaii earthquake of 1975, J. Geophys. Res. 87, Kanamori, H. and G. S. Stewart (1978). Seismological aspects of the Guatemala earthquake of February 4, 1976, J. Geophys. Res. 83, Klein, F. W. {1981). A linear gradient model for South Hawaii, Bull. Seism. Soc. Am. 71, Lindh, A. G., G. Fuis, and C. Mantis {1978). Seismic amplitude measurements suggest foreshocks have different mechanisms than aftershocks, Science 201, Madariaga, R. (1976). Dynamics of an expanding circular fault, Bull. Seism. Soc. Am. 66, O'Neill, M. E. (1984). Source dimensions and stress drops of small earthquakes near Parkfield, California, Bull. Seism. Soc. Am. 74, Rojahn, C. and B. J. Morrill (1977). The island of Hawaii earthquakes of November 29, 1975: strong motion data and damage reconnaissance report, Bull. Seism. Soc. Am. 67, 493. Scherbaum, F. and C. Kisslinger (1984). Variations of apparent stresses and stress drops prior to the earthquake of 6 May 1984 (mb = 5.8) in the Adak seismic zone, Bull. Seism. Soc. Am. 74, Swanson, D. A., W. A. Duffield, and R. S. Fiske (1976). Displacement of the south flank of Kilauea volcano: the results of forceful intrusion of magma into the rift zones, U.S. Geol. Surv. Profess. Paper 963. Tilling, R. I., R. Y. Koyanagi, P. W. Lipman, J. P. Lockwood, J. G. Moore, and D. A. Swanson (1976). Earthquakes and related catastrophic events, island of Hawaii, November 29, 1975: a preliminary report, U.S. Geol. Surv. Circ Trifunac, M. D. and J. N. Brune (1970). Complexity of energy release during the Imperial Valley, California, earthquake of 1940, Bull. Seism. Soc. Am. 60, Tsujiura, M. (1977). Spectral features of foreshocks, Bull. Earthquake Res. Inst., Tokyo Univ. 52, Wesson, R. L. and W. L. Ellsworth (1973). Seismicity preceding moderate earthquakes in California, J. Geophys. Res. 78, Wilson, M. E. (1981). Stress drops and amplitude ratios of small earthquakes preceding the 1975 Hawaii M = 7.2 mainshock, M.S. Thesis, University of Colorado, Boulder, Colorado. Wu, F. T. and H. Kanamori (1973). Source mechanism of February 4, 1965, Rat Island earthquake, J. Geophys. Res. 78, Wyss, M. (1986). Seismic quiescence precursor to the 1983 Kaoiki (Ms = 6.6), Hawaii, earthquake, Bull. Seism. Soc. Am. 76, Wyss, M. and J. N. Brune (1968). Seismic moment, stress, and source dimensions for earthquakes in the California-Nevada region, J. Geophys. Res. 73, Wyss, M. and J. N. Brune (1971). Regional variations of source properties in Southern California estimated from the ratio of short-to-long period amplitudes, Bull. Seism. Soc. Am. 61, Wyss, M., F. W. Klein, and A. C. Johnston (1981). Precursors to the Kalapana M = 7.2 earthquake, J. Geophys. Res. 86,