Analysis of the 1986 Mt. Lewis, California, Earthquake: Preshock Sequence-Mainshock-Aftershock Sequence

Physics of the Earth and Planetary Interiors, 75 (1993) 267-288 267 Elsevier Science Publishers B.V., Amsterdam

Analysis of the 1986 Mt. Lewis, California, earthquake: preshock sequence-mainshock-aftershock sequence

Yi Zhou, Karen C. McNally and Thorne Lay Institute of Tectonics and C\ F. Richter Seismological Laboratory, Unil.,ersity of California, Santa Cruz, CA 95064, USA (Received 27 April 1992; revision accepted 7 July 1992)

ABSTRACT

Zhou, Y., McNally, K.C. and Lay, T., 1993. Analysis of the 1986 Mr. Lewis, California, earthquake: preshock sequence- mainshock-aftershock sequence. Phys. Earth Planet. Inter.. 75: 267-288.

The 1986 Mt. Lewis earthquake (M L = 5.7) occurred on a right-lateral fault northeast of and oblique to the Calaveras fault in a region that had not experienced significant seismicity since 1943. Data from the nearby Lawrence Livermore Seismic Network and selected U.S. Geological Survey stations are used to relocate events within 15 km of the mainshock epicenter during the period 1980-1987, using the master-event method. Beginning 17 months before the mainshock, 22 events ruptured in the depth range 5-9 km within 1.4 km and northwest of the mainshock epicenter, in an area subsequently almost devoid of aftershocks. This cluster of preshock activity is clearly separated both spatially and temporally from the background activity in the surrounding area. Composite focal mechanisms for the preshocks and for nearby aftershocks suggest that there are two slightly different focal mechanisms amongst the preshocks, one being similar to the mainshock and aftershocks and one being rotated in strike. Cross-correlations of digitally recorded short-period waveforms of 10 of the clustered preshocks (Mu. = 1.5-2.5) reveal that the average inter-event peak cross-correlation between seismograms is 0.62. Six nearby early aftershocks show an average inter-event peak cross-correlation between seismograms of 0.54. Only a few aftershocks have cross-correlations of 0.6 or higher, which implies that the events were slightly further apart or multiple mechanisms were active during the early aftershock period. No significant differences are observed between the spectra of the preshocks and aftershocks. The aftershock area expanded along the strike with time, ultimately defining a north-south fault plane, 11 km in length, extending from 3 to 10 km in depth. The mainshock appears to have originated at the base of the seismogenic zone and ruptured bilaterally along strike and updip. A forward modeling technique is used to model teleseismic body waveform data of the mainshock. The long-period data (P and SH) are consistent with a point source strike-slip earthquake with a teleseismic moment of 3.9 × 1017 N m. We infer that the mainshock involved the rupture of an asperity in the central portion of the aftershock zone.

I. Introduction variation of seismicity patterns. One of the most promising procedures for understanding this The study of seismicity patterns is one of many complexity is to analyze preshock, mainshock, approaches that can be used to understand the and aftershock behavior (e.g. Doser, 1989), but physical mechanisms of earthquake failure pro- relatively few events with clear preshock se- cesses. Unfortunately, seismic behavior exhibits quences have been studied. Many investigators substantial complexity and regional variation. The have studied spatio-temporal variation of seismic- complex structure and stress heterogeneity of ity before large earthquakes, in attempts to un- fault zones may be responsible for the observed derstand the physical processes leading to rupture (e.g. McNally, 1977; Ishida and Kanamori, Correspondence to: T. Lay, Institute of Tectonics and C.F. 1977, 1980; Ishida and Ohtake, 1984; Smith and Richter Seismological Laboratory, University of California, Priestley, 1988). These investigations have often Santa Cruz, CA 95064, USA. found that seismic activity before the mainshock

0~)31-9201/93/$06.00 © 1993 - Elsevier Science Publishers B.V. All rights re~rved 268 Y. ZIIOU E'I" AL. is located at or near the mainshock epicenter. plane that lacked both preshock and aftershock This can be interpreted as the result of stress activity. They inferred that before the mainshock increase in and around the rupture nucleation localized asperity regions wcrc locked and quies- region before the mainshock. It has also been cent whereas surrounding weaker areas on the observed that for some earthquakes with predom- fault slipped with small events or aseismically. inantly strike-slip focal mechanisms, the spatial Schwartz et al. (1989) examined the relationship distribution of well-located aftcrshocks has a dis- between aftershocks and mainshock fault slip for tinctive pattern. Reasenberg and Ellsworth (1982) other large intcrplatc events. They also found studied aftershocks of the Coyote Lake, Califor- that aftershock hypocenters tended to be absent nia, earthquake. They found that the larger after- from regions where mainshock slip was concen- shocks surround a quiet region whereas preshocks trated. Near-total relaxation of stress in the rup- concentrate within the central quiet portion of tured asperities is one possible explanation for the aftershock zone. The 1984 Morgan Hill earth- the sparsity of aftershocks in the asperity regions. quake and its aftershocks were studied by Cock- Hartzell et al. (1991) and Beroza (1991) found a erham and Eaton (1987). They showed that the similar tendency for aftershocks to lic outside most intriguing feature of the aftershock distribu- primary mainshock slip regions for the Loma tion is a central quiet zone surrounded and al- Pricta event. Thus, although exceptions certainly most completely outlined by aftershocks. The exist, it is often possible to infer where regions of mainshock hypocenter lies within this quiet zone principle coseismic slip arc located on the basis at its northwest end. of reduced aftershock activity. Tajima and Kanamori (1985a,b) investigated Although seismicity data are of great impor- the patterns of aftershock area expansion associ- tance, not all of the source physics information is ated with large subduction zone earthquakes on a included in hypocentral locations and time se- global scale. They argued that if the fault zone is quence. The actual seismograms contain more represented by relatively large asperities sepa- detailed information about the sources, but it is rated by small weak zones, then little expansion not always straightforward to extract that infor- of aftershock activity is expected. On the other mation. If the asperity model proposed by hand, if relatively small asperities are sparsely Kanamori (1981) is correct, foreshocks should bc distributed, significant expansion may occur. The concentrated along the edges of strong asperities. aftershock area expansion pattern may thus re- Foreshocks at stress concentrations may occur as flect the spatial variation of fault-zone properties. groups of events with very similar locations and Mendoza and Hartzeil (1988)and Hartzell and focal mechanisms and thus very similar wave- Iida (1990) analyzed aftershock patterns for sev- forms. Also, stress drops of preshocks should be eral moderate to large earthquakes. They found high on average if they represent the final stages that aftershocks tended to occur outside or near of stress accumulation around asperities (e.g. the edges of the regions of principal slip in the Zufiiga et al., 1987). mainshock. The pattern of aftershock activity Relatively few studies have investigated wave- concentrated around primary slip zones implies a forms of preshocks and aftershocks to test these redistribution of stress after the earthquake, as a ideas. Ishida and Kanamori (1978) observed five result of increased loading away from the area of events located in the epicentrai region of the greatest seismic moment release. Engdahl et al. 1971 San Fernando earthquake during the 2 years (1989) and Houston and Engdahl (1989) com- before the earthquake, and found that the wave- pared the spatial distribution of coseismic slip in forms were remarkably similar. Geller and the 1986 Andreanof Islands earthquake with relo- Mueiler (1980) studied four ME=2.7 earth- cated seismicity before and after the event. They quakes on the San Andreas fault in Central Cali- found that preshock and aftershock seismicity fornia. They hypothesized that earthquakes pro- coincided spatially, but the mainshock moment ducing nearly identical waveforms must have sim- release tended to occur in regions of the fault ilar focal mechanisms and hypocenters within THE 1986 MT. LEWIS, CALIFORNIA. EARTHOtJAKI~ 269 one-quarter of the shortest wavelength. To test Lawrence Livermore Seismic Network and a sub- this hypothesis, Thorbjarnardottir and Pechmann stantial number of preshocks and aftershocks (1987) studied cross-correlation of bandpass- were detected. Accurate relative locations can be filtered seismograms of open-pit mine blasts in obtained, and the digital recording capability of Utah with known locations. Their results sup- the network allows an examination of the wave- ported the one-quarter wavelength hypothesis and forms of many small preshocks and aftershocks. therefore increased the confidence in application This earthquake is somewhat unusual in that of this method to earthquake data. Pechmann most large California earthquakes have not had and Kanamori (1982) studied small-scale earth- pronounced foreshock activity. It occurred on a quake clustering before and after the (M~. --- 6.6) previously unmapped north-south-striking fault, 1979 Imperial Valley earthquake in California. located northeast of the Calaveras fault and south Evidence was found that the clustering of seismic- of Livermore, California. No surface rupture was ity is within source volumes less than 0.5 km in found. We use the master event method to obtain extent for time periods of up to 26 months before precise relative locations, which are then used to the mainshock and also during some of the after- understand better the patterns of preshock and shock sequences. Pechmann and Thorbjarnardot- aftershock activity. The maximum value of the tir (1990) used waveform cross-correlation to cross-correlation function calculated tor seismo- study clustering of small earthquakes. They found grams of different events recorded at the same an unusual cluster of four small (ML~< 1.7) station is used as a quantitative measurement of preshocks which occurred 10 months before the waveform similarity. The results show that 1982 Richfield earthquake and had nearly identi- preshocks have similar waveforms over the entire cal waveforms. Based on Kanamori's (1981) as- record length (approximately 15 s), and some perity model, the groups of similar earthquakes aftershocks also have similar waveforms. How- may represent the localized failure of small fault ever, the percentage of events with similar wave- asperities or clusters of asperities. forms decreases after the mainshock occurred. Frequency content of earthquakes depends on Analysis of the frequency content of these events a number of factors such as the stress level in the shows no systematic temporal change in spectral focal zone, the source dimension, near-source shape. The Mt. Lewis earthquake is a moderate velocity structure and the rupture process. To earthquake and produced a sparse data set of compare the frequency content of preshocks with short- and long-period observations on a global that of aftershocks, one should take events of scale. Wc model the teleseismic body waves in comparable magnitude from the same source area the time domain to determine the source proper- recorded by common sensors. Opening or closing ties. A simple point source can satisfy the long- of cracks or movement of pore fluids that change period teleseismic body waves. medium properties such as anelastic attenuation near the fault may complicate any comparison of source radiation. It may thus be difficult to detect 2. Spatial-temporal variation of seismicity temporal changes in stress drop if the mainshock strongly perturbs the medium. The data used here were obtained from the In this paper, we present a study of a moder- local short-period Lawrence Livermore (digital) ate California earthquake. This earthquake oc- Seismic Network (LLSN) and the U.S. Geological curred on 31 March 1986 near Mt. Lewis in Survey (USGS). P- and S-wave arrival times and central California. The local magnitude is 5.7 P-wave first-motion data were used. To perform (McKenzie and Uhrhammer, 1986). This earth- the relocation, we selected events from the USGS quake (11:55:39.39 UTC; 37°29.08'N; 121 ° catalog that were within 15 km of the epicenter of 41.63'W; depth 8.14 km) provides an opportunity the mainshock and were recorded by at least four for a detailed study of a preshock and aftershock stations. All of the events were relocated using sequence, as it occurred within the dense the master event method (e.g. Johnson and 270 Y. ZHOU ET AI..

Hadley, 1976) during the period from 1980, when what deep focal depth for some of the master the LLSN was installed, to 1987. Table 1 lists ¢verfls. We were concerned about deficiencies in master events which include the mainshoek, Sta- the preliminary velocity model. To achieve im- tion delays for each station were calculated by proved master event locations, a velocity model taking the mean of the travel-time residuals (ob- (see Table 2), which is a slightly modified version served minus calculated) found for the master of the standard one-dimensional velocity model events. A restricted, or homogeneous, station ar- used for locations with the Livermore array, was ray was used to reduce the group variance intro- used to relocate the earthquakes. We tested our duced by any change in the number of stations model by omitting the nearest station when locat- recording large events vs. small events. Triangles ing the mainshock and we obtained almost the in Fig. 1 indicate the stations which were used for same depth as when the nearest station was in- the relocations in the present study. These sta- cluded. However, there are, of course, still uncer- tions are distributed over a wide azimuthal range tainties in the absolute source parameters, partic- around the aftershock zone of the Mt, Lewis ularly depth. A three-dimensional velocity model earthquake (rectangular area in Fig, 1), An aver- for the region is not available. S-wave velocities age crustal velocity model for the Liver((lore val- were calculated assuming a Vp/V~ ratio of 1.71. ley region (Hauk, 19851 was first used to locate The locations of the earthquakes were deter- the master events. This model yields larger than mined using the program HYPOINVERSE expected vertical errors in locations and some- (Klein, 19851.

TABLE 1 Hypocenter parameters of the master events Origin time Latitude Longitude Depth Local r.m.s. Error (N) (W) (kin) magnitude Horizontal Vertical 800408 1435 38.17 37-29.23 121-40.53 5,79 2.03 0.07 0.6 1.3 82(1127 0505 13.13 37-28.75 121-42.69 5.30 1.63 0.06 0.6 1.5 820128 1538 56.23 37-28.89 121-42.06 5.39 2.35 0.07 0.7 1.3 820323 1054 46.31 37-30.38 121-42.29 5.13 1.89 0.06 0.5 1.1 821126 0621 40.73 37-30.45 121-42.37 1~.38 1.82 0.03 0.6 1.1 830617 2250 33.15 37-29.64 121-42.34 5,38 2.31 (I.06 0.5 0.9 8311701 2302 28.44 37-29.75 121-42.71 5.14 2.02 0.09 0.9 1.5 830729 0003 07.5(I 37-28.32 121-39.69 7.60 2.55 (I.07 1.1 1.5 850223 1532 36.35 37-29.76 121-43.64 4.82 1.68 0.06 0.4 1.1 8511510 I1747 I13.34 37-27.82 121-40.42 6.51 1.57 0.05 0.4 (1.7 850605 2037 52.34 37-29.62 121-41.94 6.43 2.17 0.05 0.4 0.8 850712 2257 00.14 37-28.66 121-41.94 5.38 1.61 0.04 0.6 1.4 850924 0725 08.21 37-29.46 121-41.75 7.62 2.17 0.04 0.4 1.1 85(1924 21107 40.30 37-29.48 121-41.71 7.49 1.57 0.04 0.4 1.4 851005 1024 0().33 37-29.48 121-41.70 8.42 1.50 0.05 0.4 1.4 851010 0600 41.67 37-28.97 121-40.82 8.43 1.91 I).04 0.4 1.0 860109 1109 55.47 37-29.47 121-41.84 5.86 1.64 0.03 0.4 0.9 860112 0648 56.67 37-29.63 121-43.97 5.72 1.67 0.04 0.4 1.0 860211 0839 14.56 3%29.85 121-41.92 6.06 1.96 0.03 0.4 (I.8 860320 1542 56.66 37-29.03 121-41.85 8.29 1.50 0.04 0.5 1.0 860324 0154 41.04 37-29.22 121-41.80 8.67 2.75 0.02 0.4 0.8 860324 0226 21.29 37-29.27 121-41.75 8.63 2.31 0.02 0.4 0.8 860324 0506 0(I.06 37-29.27 121-41.72 8.39 1.70 0.03 0.5 0.9 860324 0629 07.46 37-29.12 121-41.79 8.31 1.84 0.05 0.5 1.0 860331 0404 59.43 37-29.11 121-41.67 8.18 2.45 0.01 0.4 1.0 860331 1155 39.39 37-29.08 121-41.63 8.14 5.50 0.01 0.4 1.3 THE 1986 MT. LEWIS, CALIFORNIA, EARTHQUAKE 271 \'.,i Hl,I..,.I,.,.I.H.l....I.,,,h,..t.,.,I,...l,,,,I,,,,h,,,I, '\ \ ,,I..,,I...,t ! preshock activity. Figure 2(c) shows the temporal SS' \ " pattern of this interesting sequence. The first swarm occurred on 24 September, 1985, about 6 months before the mainshock. This swarm in- s° \ I volves five events and includes the largest event ,,,,i! x , ',,,, (ML= 3.5) in the preshock sequence. The focal mechanisms of events in this swarm are the same as that of the mainshock, judging from the polari- ties at common stations. The rate of preshock

25' f activity increased beginning with swarm 1. The CMI' p, most intense foreshock activity occurred in another swarm on 24 March, 1986, 7 days before the mainshock. These events are all deeper than 10' the mainshock hypocenter and have slightly different focal mechanisms from the mainshock, as

I \ ~\ _ discussed below. Seven hours before the mainshock, the last foreshock (31 March 1986; 20' 10' 122 ° 50' q0' 3C' ~'0' 10' 04:04:59.43 UTC) occurred with magnitude 2.5, Fig. 1. Map showing the locations of the seismograph stations focal depth 8.2 km, and a similar fault mechanism (open tri:Insles) used in the relocation procedure. The star locates the epicenter of the mainshock, The rectangular area to that of the mainshock. These preshock events spans the zone of after~hock epicenters. (Table 3) occurred about 1.4 km to the northwest of the mainshock epicenter (sec Fig. 2(b)), and were clearly separated both spatially and tempo- High-quality relative locations are required to rally from the activity in the surrounding area. study the details of the preshock and aftershock Ishida and Kanamori (1980) showcd that the 1952 patterns. After relocation, we chose 1288 events Kern County, California, earthquake was also which occurred from January 1980 to August preceded by a concentration of events located 1987 and satisfied the following criteria: (1) r.m.s. near the mainshock epicenter, and a similar pat- error of the travel time residual less than or equal tern was observed for the Coyote Lake cvcnt to 0.15 s; (2) epicentral error less than or equal to (Reasenberg and Ellsworth, 1982). 1.4 kin; (3) error of the focal depth less than or It is of interest to examine the preshock sc- equal to 2.0 km; (4) local magnitude larger than quence in detail. We use the program FPFIT or equal to 1.0. (Rcasenberg and Oppenheimer, 1985) to dcter- The seismicity distribution before the Mr. Lewis event is shown in Fig. 2. Events during the period from January 1980 to 17 months before TABLE 2 the mainshock are plotted in Fig. 2(a). The clus- P-wave crustal velocity model ter of seismic activity north of the Mt. Lewis Velocity of layer (km s- t) Depth to top of layer (km) epicenter (star) is noteworthy. The trend of activ- 3.0 0.0 ity along the Calavcras fault is slightly offset from 3.5 1.0 the mapped surface location of the fault, but a 4.0 2.0 clear lineation is observed. Figure 2(b) shows the 4.6 3.0 epicentral distribution of events in the period 4.8 4.0 5.1 5.(1 from 17 months to 7 h before the mainshock. The 5.3 6,0 most notable feature of these distributions is the 5.6 7.0 concentration of preshock activity near the main- 5.9 12.0 shock epicenter beginning about 17 months be- 6.2 17.(1 fore the mainshock. There are two swarms in the 8.0 25.0 272 Y. ZHOU ET AL. mine the focal mechanisms of 22 preshocks, the cluster of preshock activity into two groups based mainshock, and nine aftershocks (Table 4) which on focal mechanism. The first group (Group A) have almost the same locations as the preshocks. includes 12 events represented by squares in Fig. As there is insufficient information to determine 3. For these events, the composite focal mecha- a well-constrained focal mechanism solution from nism has the same strike of the northwestern the local P-wave first-motion data alone for each nodal plane as that for the mainshock; N5°W. individual event (except for the mainshock), we This suggests that Group A preshocks have the combined first-motion data to generate compos- same focal mechanism as the mainshock. We also ite focal-mechanism solutions. We divided the made a composite mechanism for 10 different

, ...... , , , , , , i i i i L i i i

\ ""- ~ ~ * Magnitude

35' 35'

e. • @ • °N . * I °r * ', ~4k* I I ~ 3.o* 30' 30" ~,\ *. ** I 11r.t ," . 40- , J- * L._J." a, 5.0. 25' 25' •. ..

37o20, \ -\ 37020, x,..\ ' ' ' ' I ' ' ' ' I , ' ' , I ' ' ' ' I . ' ' ' ' I ' ' ' ' I , , , , I ' ' r , I ' 1 (a) 50' q5' q0' 121035 . (b) 50' q5' q0' 121o35 '

.~ IIIllllllrlllJll I Iilli Illllllllll~ JIiIIiJl~lllJ

main

~q ¢w z E ~3 x i 11 I IIl~Vl I IlitIJII llllll1 llill I11 II Nov 1985 Mar May Jul Sop Nov 1986 Mar May (c) Time Fig. 2. (a) Epicentral distribution of the relocated preshocks in the period from January 1980 until 17 months before the mainshock. The large star is the epicenter of the mainshock, and the rectangle spans the mainshock rupture zone. Symbol sizes are proportional to magnitude. (b) Epicentral distribution of the relocated preshocks in the period from 17 months to 7 h before the mainshock. Symbol sizes are proportional to magnitude. (c) Plot of magnitude vs. time for the preshocks. THE 1986 MT. LEWIS. CAI.IFORNIA. EARTHQUAKE 273

TABLE 3 Hypocenter parameters of the preshock events

Origin time Latitude Longitude Depth Local r.m.s. Error (N) (W) (km) magnitude Horizontal Vertical 841(131 0358 54.44 37-29.64 121-41.96 6.41 1.73 0.03 0.3 0.7 850605 2037 52.34 37-29.62 121-41.94 6.43 2.14 0.115 0.4 0.8 850924 0510 00.02 37-29.49 121-41.72 7.98 1.48 11.03 0.4 1.4 850924 0719 55.44 37-29.54 121-41.74 7.77 1.13 0.00 0.7 1.9 850924 0721 29.42 37-29.48 121-41.70 8.37 3.50 11.112 0.4 1.3 850924 0725 08.21 37-29.46 121-41.75 7.62 2.14 0.04 0.4 1.1 850924 2007 40.30 37-29.48 121-41.71 7.49 1.64 0.04 11.4 1.4 851005 11124 00.31 37-29.39 121-41.67 8.58 1.55 0.03 0.4 1.3 851009 0936 21.58 37-29.29 121-41.77 7.27 1.42 0.02 0.4 1.3 851120 1135 24.47 37-29.11 121-41.77 8.15 1.31 1/.02 (1.5 1.3 851123 1725 48.01 37-28.98 121-41.66 8.66 1.44 0.03 (I.4 I. 1 86011)9 1109 55.47 37-29.47 121-41.84 5.86 1.62 0.03 0.4 0.9 860211 0839 14.56 37-29.85 121-41.92 6.06 1.82 0.03 0.4 0.8 86032(I 1542 56.66 37-29.02 121-41.80 8.14 1.52 0.05 0.5 1.0 860324 0144 48.13 37-29.32 121-41.71 8.27 1.22 0.114 0.5 0.9 860324 0146 04.49 37-29.23 121-41.70 8.73 1.31 0.02 0.5 (1.9 860324 0146 21.62 37-29.22 121-41.67 8.42 1.19 0.02 0.6 1.0 860324 0154 41.04 37-29.22 121-41.80 8.67 2.50 11.02 0.4 0.8 860324 0226 21.29 37-29.27 121-41.75 8.63 2.24 0.02 0.4 0.8 860324 0506 00.06 37-29.27 121-41.72 8.39 1.50 0.113 0.5 0.9 860324 0629 07.46 37-29.12 121-41.79 8.31 1.67 0.05 0.5 1.0 86(1331 0404 59.43 37-29.11 121-41.67 8.18 2.50 0.01 0.4 1.(I

(Group B) events, indicated by circles in Fig. 3, To compare the focal mechanisms of these which are tightly clustered at slightly greater preshocks with those of the aftershocks in the depth than the mainshock, and found that these same vicinity, we made a composite mechanism events have a strike of N5°E for the northern for the few aftershocks located near the main- nodal plane, slightly different from that of the shock hypocenter (nine events, see Table 4), which mainshock. These subtle differences are con- are indicated by crosses in Fig. 3. It should be firmed by the P/S amplitude ratios at sensitive noted that the strike of the northwestern nodal stations along strike (see Section 3). plane of these aftershocks is the same as that of

TABLE 4 Hypocenter parameters of the aftershocks studied for focal mechanisms

Origin time Latitude Longitude Depth Local r.m.s. Error (N) (W) (km) magnitude Horizontal Vertical 860331 1224 14.71 37-29.24 121-41.72 8.80 1.64 0.03 0.4 1).9 860331 1411 17.32 37-29.40 121-41.65 9.27 1.84 0.02 0.5 1.3 860331 1422 40.00 37-29.06 121-41.68 9.46 1.31 0.04 (1.6 1.7 860331 1445 05.42 37-29.05 121-41.61 9.(]3 1.64 0.03 (J.5 1.3 860331 1517 17.09 37-28.97 121-41.57 9.53 1.38 0.02 (1.4 1.2 860331 1559 53.07 37-29.06 121-41.67 8.69 1.49 0.112 0.5 1.5 860331 1941 15.25 37-28.97 121-41.66 9.17 2.80 0.02 0.4 1.1 8611331 2130 08.89 37-29.28 121-41.74 9.06 2.80 (I.02 0.4 1.0 860331 2319 11.86 37-29.41 121-41.76 9.62 1.77 0.03 (I.4 1.3 274 Y. ZHOU ET AL.

N N N

GroupA:12eventl Gro-pB:llkvents ]dttin~:k Afmml~cl~ 9 ore,ram

30" -5 12 O -6 O % C~

D =. -7 o~ 0 +~ % D -B O & O Oo O ~O O O~ + O + + -9 + + D + + 37° 29' + + -I0 2 km s ' I q2' ~l' Distance Oan) Fig. 3. Map, cross-section and focal mechanisms of the mainshock (star1, preshocks which occurred about 17 months before the mainshock (squares and circles), and aftershocks (crosses) which have almost same locations as the preshocks. Squares show the Group A preshocks (12 events) with fault mechanism strike similar to that of the mainshock (N5°W), and circles represent Group B preshocks with rotated fault strike (N5°E).

the mainshock and Group A preshocks, but dif- fault orientation of the Group B events. It is ferent from that of the Group B preshocks interesting to examine the waveforms and spectra (circles) (see Fig. 3). The aftershocks are located of the preshocks and aftershocks (see Section 3). deeper than any of the preshocks; thus the Group The spatial distribution of the aftershocks and B mechanism is not different because of source surrounding seismicity in the 17 months after the depth alone. It is possible that a stress concentra- mainshock and the focal mechanism of the main- tion on a localized distortion of the lower edge of shock are shown in Fig. 4(a). The Mt. Lewis the mainshock asperity produced the anomalous aftershocks are the dense distribution of events

Fig. 4. (a) Focal mechanism of the mainshock and map view of the relocated aflershocks and surrounding seismicity during the first 17 months after the mainshock. Symbol sizes are proportional to magnitude. (b) North-south cross-section showing hypocenters of the aftershocks (crosses) in the mainshock rupture zone and cluster of preshock activity (circles). Distance is measured from the north, and only events in the box in Fig. 1 are shown. The events on the left side of the dashed line may be induced by the Del Valle reservoir (Wong, 1991). W o [~pth(kin)

o +

l A A i i I i i i A t I I I l I ; ÷ +

0 ...... ÷ 7. > # ÷~ 4. %

0 C

#%m "o ÷ %*

v . .

•

o + ÷ ÷ o ~ o o 9 ~ ;7 / O

I 2"76 Y./-HOU ET AI.

3/31/~ 11:53.12:08 (13M)

,2,0 , *. I,,,,, I ' ...... I ' ' '' ...... I ' ~'I .... ' .... ":'>'""' ...... ' ...... 3/31/36 12:09.13:59 (130M)

+ L +*%', Z

i i

(3/31/86") 14:00-(411/~5)8:00 (20H)

. 0.0-

°+ ,

,,,,,,''' I'' , , , , , , , , , , ,, i .... I""l''"l" I~ 5.0+ (4/I/~)8:01-(4#~/86)1 I:33 (gD)

lll[llllt l'lllll'l|Ltlllllll|lll II'IIIII~ I I I I I l l , I l I I I I I I , I l ~ I

•.,~ ;(,,...

.... '''" ' ' i r ,|, , , . . , ' ] ' ' >~.,, ...... , ......

(419/8~ I I :~,-("711/36")8:40 (90D)

°. -C

-e+ .~ .+ ,

...... ' ' ' I ..... ~' ' ' I ' " ...... ,. >.,.",~, ...... , ..... J (?/I/S6)S:dO-~d (S/'2A/S7) A' / ~6' \,}. 30'

37o~5. ' ...... ' ...... ").x\~ -I0 i ' .... -,-~,, i ...... , , I0 20 "1 ...... ~':'~'~"' .... ' .... ' ..... "~ (a) I2~p° 50' 148' 30' (b) Distance(kin) THE 1986 MT. LEWIS, CALIFORNIA, EARTHQUAKE 277

TABLE 5 tlypocenter parameters of the events selected for waveform analysis

No. Origin time Latitude Longitude Depth Local r.m.s. Error (N) (W) (km) magnitude Horizontal Vertical 1 841031 0358 54.44 37-29.64 121-41.96 6.41 1.73 0.1)3 11.3 0.7 2 850605 2037 52.34 37.29.62 121-41.94 6.43 2.14 0.05 11.4 (1.8 3 850924 11511) (X1.03 37-29.49 121-41.73 7.94 1.48 0.03 1).4 1.4 4 850924 0725 08.21 37-29.46 121-41.75 7.62 2.14 0.04 (1.4 1.1 5 850924 2007 40.30 37-29.48 12 I-41.71 7.49 1.64 0.04 0.4 1.4 6 86,0109 1109 55.47 37-29.47 121-41.84 5.86 1.62 0.03 0.4 0.9 7 860321) 1542 56.66 37-29.02 121-41.811 8.14 1.52 0.115 0.5 1.11 8 860324 0226 21.29 37-29.27 121-41.75 8.63 2.24 0.02 0.4 0.8 9 860324 0506 00.06 37-29.27 121-41.72 8.39 1.511 0.03 0.5 0.9 10 860331 0404 59.43 37-29.1 l 121-41.67 8.18 2.50 0.01 0.4 1.0 11 860331 1206 09.90 37-29.89 121-41.70 6.75 2.01 0.03 11.4 0.6 12 860331 1231 46.57 37-29.78 121-41.80 6.37 1.35 0.06 11.5 0.8 13 860331 1247 04.76 37-29.89 121-41.73 7.28 1.88 11.03 0.3 1.0 14 860331 1251 28.93 37-28.95 121-41.62 8.75 1.62 11.112 11.5 1.0 15 860331 1330 18.18 37-29.81 121-41.58 7.56 1.69 0.04 11.4 I. ] 16 860401 0630114.44 37-29.59 121-41.68 6.26 1.52 11.03 11.4 0.7 with a north-south trend. The mainshock oc- with the trend of the aftershock distribution. The curred at the center of the aftershock lineation. fault orientation is also consistent with the cen- There is no surface expression of this fault. The troid-moment tensor (CMT) solution (Dziewonski mieroseismieity thus indicates that the Mt. Lewis ct al., 1987). The CMT inversion gives a best event ruptured a north-south striking fault doublc couple (plane 1: strike, ~b = 353 °, dip, 6 = northeast of the Calaveras fault. It should be 79 ° , rake, A=170°; plane 2:4~=261 ° , t5=81 ° , noted that there is a gap and possible offset in A = -11 °) for the Mt. Lewis event. Figure 4(b) north-south-trending seismicity. The seismicity in shows the hypocentral distribution of the after- the north includes many events after the Mt. shocks (crosses) in the mainshock rupture zone Lewis event, but also has high background levels and the cluster of preshock activity (circles). It (see Fig. 2(a)), possibly induced by the Del Valle should bc noted that the cluster of preshock reservoir (Wong, 1991). A distinct southwest- activity is concentrated within the central quiet northeast trend of seismicity branches from the portion of the aftcrshock zone. It is interesting to Calavcras fault and intersects the rupture zone see that the preshocks occurred only in a limited near the Mt. Lewis mainshock epicenter. Figure portion of the quiet zone, along what appears to 4(a) reveals that, omitting the activity clustered in be a transition in the aftershock distribution. The the north, the aftershocks were distributed be- central region of the aftershock zone has very fcw tween 5 km north and 5.8 km south of the main- events over an area of about 12 km 2. Focal depths shock epicenter, implying bilateral rupture in the in the northern region of the mainshock were mainshock. The N5°W right lateral strike-slip between 5 and 8.5 km, whereas events arc in a nodal plane in the P-wave first motion solution slightly deeper range (6.5-9.5 km) in the southern for the mainshock can be confidently identified as region. The mainshock hypocenter (h = 8.14 km) the actual fault plane because it is very consistent occurrcd ncar the base of the zone of aftcrshock

Fig. 5. (a) Map view of aftershock activity fl)r six nonuniform time intervals after the mainshock. Rectangles and labeled end-lmints correspond to vertical cross-sections. (b) North-South cross-sections akmg strike of the fault for the aftershocks in the time windows in (a). Only earthquakes whose epicenters are located within the rectangular boxes shown in (a) are plotted and distance is measured from the north end of the box. 278 Y. ZHOU ET AL. activity, but aftershocks do extend to as deep as for the second, and the final length of the after- 10 km. shock area is about 10,8 km, which is 3.6 times as Another interesting characteristic of this se- large as in the first time period. Fig. 5(b) shows quence is that the area of aftershock activity vertical cross-sections along the rupture zone for expands with time. To illustrate the expansion these six time periods. We calculated the after- with time, the aftershock zone is plotted in Fig. 5 shock area from these cross-sections by assuming for six discrete timc periods: (a) 13 min after the a circular rupture area. The aftershock area is mainshock; (b) between 13 and 130 min; (c) be- about 12 km 2 for the first time period and about tween 130 min and 20 h; (d) between 20 h and 9 28 km 2 for the second. The aftershock area within days; (e) between 9 days and 3 months; (f) bc- 9 days increases to 50 km 2, which is 4.2 times as tween 3 and 17 months. It can be seen in Fig. 5(a) large as in the first time period. It should be that the length of the aftershock zone increased noted that after 13 rain there are no aftershocks gradually and systematically in the north-south within the patch of about 4 km in diameter near direction. The length of the aftershock zone is the mainshock hypocenl~er (see Fig. 5(b)). The about 3 km for the first time period, about 7 km last time interval showa only slight expansion of

CMN (E-W) CDV (E-W)

__...... _ JLIi_l.t|r'v ...... t ......

..90 ..... ~llh ,.,.~ ...... 1180109 -- --~pr ......

.1~ . ~_~-~

It" ...... ~I lJ~,~,~...... ~"~_

.86 860331.1206 ./,,~~a. " 8~k~1.1231

o ~1

I .... I .... I , , , , I 0 5 Sec 10 15

Fig. 6. Seismograms for some preshocks and aftershocks recorded at two three-component stations (CDV, A = 9.1 km: CMN, A = 15.9 km) after filtering the data from 1.0 to 20 Hz. They are in chronological order. Horizontal dashed lines separate the preshocks from the aftershocks. Seismograms are positioned horizontally according to the recalculated origin times, and are plotted with the same maximum amplitude. The maximum correlation between each seismogram and the seismogram above it is shown at the left of each seismogram. THE 1986 MT. LEWIS, CAI.IFORNIA, EARTHQUAKE 279 the aftershock zone beyond its 3 month size, ing naturally rotated SH arrivals in a stable part although activity to the north near the reservoir of the radiation pattern. gives an apparent expansion. Because the character of seismograms of local seismic events is strongly controlled by the details of the local crustal velocity structure, similarity of the waveforms implies events situated close to- 3. Waveform analysis gether with similar source mechanisms. Some similarities are observed in Fig. 6. To quantify the To test for any temporal changes, we compare similarity of the seismograms, we cross-correlated the ~vaveforms and spectra of the preshocks with the filtered seismograms recorded at the same aftershocks which have roughly the same size and station, following the procedure of Pechmann similar source-receiver geometry. We analyze and Kanamori (1982). The record length used for events with local magnitudes ranging from 1.5 to the cross-correlations was 15 s, which includes 2.5 for the preshocks and from 1.35 to 2.1 for the the entire window of P and S arrivals and their aftershocks. Waveforms of events larger than 2.5 are clipped and those less than 1.4 have a low Preshocks signal-to-noise ratio. Table 5 lists the earthquakes selected for analysis, which include 10 preshocks 1-2 2-4 l'Ix 1'0 I and six aftershocks. o.e "':.".-:-.._. "..~ .'. 0,g0'11 ~. ~ .i..." • Good recordings for most of the events in 0.20'40'6t "" "' "" " :' 0.4 Table 5 were available at three three-component 0.2 i i i i i i i i short-period stations (CDV, CMN, and CVL) and 0.0 1.0 Z.0 3.0 0.0 1.0 210 i 3.0 two vertical-component stations (CPL and CMJ) 4-8 F.z 8-18 nz of the LLSN array (see Fig. 1). These stations 1.0 [.:..; ...... have very similar instrumentation, and span a 90 ° 04..,.'.,. :. azimuth range from the source. We were particularly interested in three-component data along ~i~l~,°'"/. ,-..,-,,;" "i 0,0 1.0 2.0)", 3.0 0.0 t.O ~'.0 strike of the fault, and none of the other regional 310 Otstanca (kin) USGS stations had three-component data at close stations, so we limited ourselves to the LLSN Aftershocks recordings. The digital seismograms from these 1-2 ~ 2-4 Hz stations were used in the studies of waveforms and spectra of prcshocks and aftershocks dis- 0.8 % * * • • 0.8 /~ ° ° • • . 0.8°I t" 0.0°I t " ' cussed in the following sections. 0.4 0.4 Figure 6 shows some records from two stations 0.2 0.2 along the strike of the fault. Horizontal dashed 0.0 |.0 20 30 0.0 10 2.0 30

lines separate the preshocks from the after- 4-8 [-Lt 8-16 Hz shocks. Seismograms are aligned with actual travel 0.8 : • o~ 0.8 0.s ; t, , ' " ~'o.e times according to the recalculated origin times. " " N ~° .. ". Electronic noise is high above about 20 Hz and 0.4 ~j 0.4 • . 0.2 0.2 the signal-to-noise ratio is low below about 2 Hz | I I i. I I .... [ I I 0.0 10 20 30 0.0 1.0 Z0 30 (Zhou, 1992), so we have bandpass filtered the Vlltence (kin) data between 1.0 and 20 Hz. This also reduces Fig. 7. Mean of the maximum cross-correlations calculated for source size effects. The maximum correlation be- preshocks and aftershocks for filtered seismograms from five tween each seismogram and the seismogram stations (three three-component stations: CDV, CMN and CVL: two vertical component stations: CPL and CMJ). Values above it is shown at the left of each seismogram. are plotted vs. separation distances between events. The hori- Clear direct P and S waves are apparent, with zontal lines (at 0.6) indicate a possible separation between east-west components at thesc two stations hav- 'well-correlated' and 'poorly correlated' events. 280 Y. ZHOU ET AI.. codas. This relatively long window is used to 1-20 Hz No. Event 1 10/31/8,1 exploit the full complexity of the crustal transfer 2 06/05/85 3 09/'/A/850510 function, heightening sensitivity to the relative 4 09/7,A/850725 IO0 O00IOl~IIIII locations and radiation patterns. The maximum l~O0 O010l~Illl• 5 09/'~A/852007 00000 OIO_O~OI 0~0 6 01/09/86 gO0000 OOCRDI O00I value of the cross-correlation function was deter- IlIllIO O I~IIIIl 7 O3/2O/86 NOOOOOOO~OI~I 8 03r~,4/86 0226 mined for all possible event pairs recorded at the immmmmmoni~ iOmImle 9 03/24/86 0506 NNNImOOIO=ILOONOm NONIIONNO~ •00N I0 03/31/86O4O4 five stations. I mmm••ommm~m mIo II 03/31/86 1206 NONINOOION~O° m• 12 03;31/86 1231 The degree of waveform similarity depends on mommmoOmOmpOmI • •I•mIOIiI•~Immi 13 03/31/86 1247 the characteristics of the random medium; differ- 14 03/31/116 1251 15 03/31/1161330 ent paths clearly give different correlations, which • o 0 0 16 04/01/860630 may involve both source and medium effects 0.2 0.6 1.0 (Zhou, 1992). Smith et al. (1982) evaluated the possibility that the average response of a large 1-2 Hz 2,4 Hz rigid foundation may be less than the peak values o8••o°=o=~o0ooo[~ 0 O000000l~OIO00]2 observed at a single free-field station. Summing IO 000000~000000 io OOOOOO0~ooooo13 IO000000NOOOOOO I00 000001100000014 • 000 OOOOqOOOOOO I000 O000~QO000013 correlations for several stations may average out o0000 000o~0°00o o0oo0 0o0olo0ooool3 =o0000 ooqo0000o BOO000 OOqO0oooo17 both random fluctuations and wave transit ef- IO00000.OqlOlOOI IOOOOOO.O~iOIOOll8 I0000000 qOOOOOO 0ooooo0o qooooool0 fects, isolating source similarity. Hence, we de- IIIIlIOOIO llOlO~l IlllmOOOOlIOlOOlllO !o000000lOII OOOOO O•IO00QIOIIOO000111 O000000000~ lOOl oooooooooolo ooooln cided to average the correlations for all five sta- IlOOOIOIOI~OI IO0 lIOOOOOlOi~O OO0113 IO0000000qOOI,Oi 00000000~00 0II14 tions to obtain more robust event-to-event corre- • ooooo0ooqO0oo o oooooo0oooKx)o0 o is OOOOOOOIOm~NOIO OOOOOOOlOl~OOlO 18 lations. As the similarity of the waveforms places a 4-8 Hz 8-16 Hz strong constraint on the maximum distance between the hypocenters, we cross-correlated all • lllllll lililOI 1 possible event pairs recorded at the five stations Ii iiiii after bandpass filtering them in four one-octave 0 passbands (1-2, 2-4, 4-8, and 8-16 Hz). Figure 7 eo° 0 shows the mean maximum cross-correlation for _,~:g:g~:g~o~g: ,, IIIIIIIII IIIIIl I0 iiiiiii!!liiiii12 ,," the five stations calculated separately for the °°IIN•OBOO~O OOO 13 • 13 OOOOOOOOO~OO OR 14 ...... o.o.l ! o • ,, prcshocks and aftcrshocks for the four passbands, ioimooO°O~OOO • ~S • o,:::,,N, : : 13 °OOOoOOOO~OOni le IIl iiI iII I~ plotted against event separation distance. The mean values involve equally weighted sums of all Fig. 8. The mean of the maximum cross-correlations calcu- 11 component correlations, although not every lated for filtered seismograms from the five stations for all possible pairs of preshocks and aftershocks. Each symbol combination is available. The horizontal lines are represents the peak correlation for the event pair correslmmd- drawn at 0.6, which is the arbitrary cutoff value ing to its position on the matrix. The symbol size is propor- used by Pechmann and Kanamori (1982) to sepa- tional to the correlation value. Open circles indicate values rate 'well-correlated' events from 'poorly corre- greater than or equal to 0.6 and solid squares indicate smaller lated' events. This figure shows that the highest values. Event numbers are keyed to upper-right corner and Table 5. cross-correlation values arc for nearby events. The cross-correlations level off to a relatively low constant value with increasing distance between This is not surprising, as higher-frequency signals events. The maximum source separation dis- should be less likely to yield high cross-correla- tances between events which have correlation co- tions by chance. The preshocks and aftershocks efficients of 0.6 or greater decreases as the center show similar behavior. Corresponding trends for frequency of the filter passbands increases. The individual stations are similar, but vary in base- average correlation coefficients for the more dis- line. tantly separated events also decrease as the cen- For the filtered data, Fig. 8 shows the mean ter frequency of the filter passbands increases. maximum cross-correlations calculated for all THE 1986 MT. I.EWIS, ('AI.IFORNIA, EAR'I'IqOtJAKE 281 possible event pairs from five stations, plotted as preshocks are somewhat more consistent. For the cross-correlation matrices. Within each matrix, slightly filtered data (1-20 Hz), the average peak the events are arranged chronologically. Each correlation for the preshocks is 0.62 (upper left symbol represents the mean peak cross-correla- box) and that for the aftershocks is 0.54 (lower tion between an event pair. The size of each right box). The percentage of the preshocks with symbol is proportional to the peak correlation similar waveforms (the mean of the maximum value. Open circles indicate values greater than cross-correlation is greater than 0.6) is 60%, which or equal to 0.6 and solid squares indicate smaller is larger than that of aftershocks, 33.3%. values. This figure indicates that the degree of The average (five stations, 11 total compo- similarity between preshocks and that between nents) peak cross-correlations for preshocks and aftershocks are not dramatically different, but the aftershocks in each octave passband are shown in

Preshocks Aftershocks 1 , [ I I 1 I I 1 l I I I I I l 1

,- 0.8 I- 0.8 O

"~ 0.6 0.6 et. ~

0.4 ~ 0.4 .<

1 1 t I I I I 0.2 I I I I I I I I 0.2 0 3 6 9 12 3 6 9 12 Center of passband (a) Center of passband

f I'. 0,8 0.8 1

l,( i. 0.1 I 0.6 ""--4

~0.4 • 0.4 < mainshockl 0.2 J ! 0.2 1984 1985 ltl 8~ M~.Ig, 19U Time AIx"l (b) time (C)

Fig. 9. (a) The average (five stations) peak cross-correlation between filtered records (open symbols) for preshoeks and aftcrshocks for the event pairs in Fig. 8. For the filtered (1-20 Hz) seismograms, the average peak cross-correlations also are shown for reference by the solid symbols at the left of each graph. (b) Average (five stations) peak cross-correlation values for each pair of consecutive events for filtered data (I-20 tlz) are plotted as a function of time. The arrow marks the time of the mainshock. (c) Enlarged plot of the section around the mainshock. 282 Y. ZHOU ET AI..

Fig. 9(a) for the event pairs in Fig. 8. The average record is available in most cases, and P and S peak cross-correlations for preshocks and after- waves and their coda may be affected differently shocks of the lightly filtered (1-20 Hz) seismo- by changes in the attenuation or scattering prop- grams are shown by the solid symbols at the left erties of the medium, we have windowed the of each graph. Both preshocks and aftershocks individual phases for the spectral analysis. We are well correlated (average peak correlations find that the S-wave spectra of the events appear generally are greater than 0.6) up through at least to have basically the same shape, with corner the 4-8 Hz frequency band (see Fig. 9(a)). We frequencies at about 6 Hz and similar rates of would like to use these correlations to obtain fall-off for frequencies above the corner. No sys- some estimate of the spatial clustering of events. tematic differences are apparent between pre- Assuming that the near-source P-wave velocity is shocks and aftershocks. The observation of stable about 5.6 km s- I, at 6 Hz the wavelength is about spectral shape is also apparent in the P-wave 933 m for the P wave. In this case, the similarity spectra. between preshock waveforms and that between aftershock waveforms imply a maximum event separation of 1/4 wavelength, or approximately 4. Teleseismic modeling 233 m. As the correlations are for band limited data and the level of correlation is not 1.0, we Long- and short-period body wave data estimate that the well-correlated events all lie rccorded at World-Wide Standard Scismograph within 1.5 km or so, which is consistent with the Network (WWSSN) stations were collected for master event locations. the Mt. Lewis mainshock and analyzed in this Average peak cross-correlation values from the study. The event is small, with unfavorable tele- five stations for each pair of consecutive events seismic P-wave radiation typical of vertical strike for the lightly filtered data (1-20 Hz) are plotted slip events, so we rely on the S waves, which are vs. time in Fig. 9(b). The arrow marks the time of somewhat larger. The S-wave data were hand-dig- the mainshock. To see the details better, we itized and rotated to isolate the SH component. expand the interval around the mainshock (sec Few digital data were available for this event Fig. 9(c)). We can see that some differences exist other than the Global Digital Seismographic Net- between the preshock and aftershock time peri- work (GDSN) long-period body wave and surface ods. The preshock swarms involve groups of very wave signals previously analyzed in the CMT similar events, giving spikes in the plot. Early inversion of Dziewonski et al. (1987). Synthetic aftershocks are different from the last prcshock, seismograms were calculated using the forward and variable from onc another. We have consid- teleseismic modeling technique of Langston and ered aftershocks only within a day after the main- Helmberger (1975), and the structural parameters shock, whereas we have considered preshocks used in the modeling are given in Tables 6 and 7). which span a time period of about 17 months. The velocity model in Table 7 has an average The overall aftershock expansion caused there to velocity the same as the structure we used for be few later aftershocks very close to the preshock relocation (Table 2). The focal mechanism deter- locations, so we did not analyze more aftershocks. mined from first-motion data was used as our An analysis of the signal spectra was under- starting point. It has ~b = 355 °, ,5 = 85 °, and A = taken in an attempt to identify any changes in the source parameters of the earthquakes not readily apparent in the correlation procedure (e.g. Ishida TABLE 6 and Kanamori, 1980; Ishida and Ohtakc, 1984). Parameters in the modeling In this study, spectral analysis was performed for Parameter P wave SH wave the events in Table 5, to investigate objectively the difference in thc frequency content between Attenuation t * 1.0 s 4.0 s the preshocks and aftershocks. As at least 15 s of Basic rays P, pP and sP S and sS "l'|tl~ 1986 MT. LEWIS. CAI,IFORNIA. EARTHQUAKE 283

TABLE 7 trapezoid (t, = 1 s, t 2 = 1 s, t 3 = 1 s) can also be Source velocity model used, and the overall duration uncertainty is + 2 s. The best fitting focal mechanism has 4, = 172 °, P-wave velocity S-wave velocity Density Thickness (km s ,) (km s -I) (gcm -3) (km) 8 = 83 °, and A = 181 °. This mechanism fits the teleseismic signals significantly better than the 5.6 3.27 2.6 20 near-source solution, perhaps indicating a small 8.0 4.62 3.2 - change in dip of the fault during rupture. From this analysis, it is apparent that the available long-period teleseismic body wave signals of the -160 °. At both short and teleseismic distances Mt. Lewis earthquake may be explained by a there is the problem of earth structure uncer- simple model. tainty, which could easily give some inconsistency We have used our best point source solution to in focal mechanism determinations, so we allow calculate the synthetic amplitudes. By comparing this orientation to be perturbed in the modeling. the observed peak body wave amplitudes with the Here, we concentrate on modeling waveshape synthetic amplitudes, the moment determined characteristics produced by the source and ac- from SH modeling is 3.9 × 1017 N m. The tele- companying surface reflections to find a best seismic P waves indicate a higher moment, but we mechanism. do not have confidence in this because the P The analysis of the long-period body waves consists of matching the observed teleseismic P waveforms and SH waveforms with synthetic seis- Long Period P and SH waves mograms. We found four stations which have fair long-period P-wave records and six stations which SH-waves SH-waves mikc.172o dip.830 I~e~181o have good SH-wave records. Fortunately, these stations are located near both nodal planes of the "°"""d"" A focal sphere, so their waveforms are sensitive to datl I ~/ o D. /..,.,source / / -4"/ / ~\\ [ changes in the fault geometry; however, this makes the P-wave moments particularly unstable. Figure 10 summarizes the analysis including data, synthetics and station distribution. The time function is a symmetric triangle with a 3 s duration. Several observations can be made from this figure. First, the reflected phases are separated from the direct arrivals at most of the stations; a single point source is sufficient to model both the P waves and SH waves. The pulse width and the interval between the two pulses are the major parameters controlling the fit of the synthetics to the data. These are functions of the source time P-WaVeS function duration and the depth of the event. A depth of 9 km is the best-fitting depth using our average crustal model. This source depth com- 3 ~.8 .5 ~..~V~ l pares favorably with the 8.14 km depth determined by our relocation, which was found with a structure that has some low velocity near surface Fig. 10. Summars' of the long-period P and SH bode-wave layers. The source duration of 3 s is not unique analysis. Teleseismic P and SH data are upper traces and because the long-period waves are not very sensi- synthetics are bottom traces. The ratio of the station momen tive to small changes in the time function. A to the SH average moment is shown on each trace. 284 v. ZHOU ET AL.

Short Period P-waves zoidal source time function (t, = 0.25 s, t 2 = 0.5 s, 13 = 0.25 S). Figure 11 shows the short-period synthetics for this model. The figure reveals that our short-period solution can roughly match some aspects of the data, but more complex propagation and source effects are not resolvable with this information. We found that the short-period solution does Syn N~jG~A not fit the long-period data as well as the long- period solution. The problem is probably the result of trade-offs in the source time function, crustal velocity structure, and data quality. As the short-period data quality is not high, we do not pursue further modeling.

5. Discussion Fig. 11. Summary of the teleseismic short-period P-wave analysis. Upper traces show data; bottom traces show synthetics. This investigation of the preshock-mainshock-aftershock sequence of the Mt. Lewis earthquake has revealed several interesting fea- tures. Microearthquake activity in the impending waves are in such unstable portions of the radia- rupture zone increased beginning 17 months be- tion pattern. fore the mainshock. The preshocks had very simi- The short-period teleseismic data set consists lar waveforms, and were clustered tightly near of only four records in which the signal-to-noise the mainshock hypocenter and within the central ratio is relativeb favorable. The short-period quiet portion of the subsequent aftershock zone. modeling procedure is analogous to that for the First motions suggest that two slightly different long periods. Initially, synthetics are computed by fault mechanisms were activated during the 17 using the long-period solution with the following months before the mainshock. The percentage of parameters: (1) qb = 172 °, 6 = 83 °, and A = 181°; preshocks with similar waveforms in the same (2) a 3 s symmetric triangle source time function; region was higher than that of the aftershocks, (3) focal depth 9 km; (4) t* = 1 s, as used in the and the variability of waveforms of consecutive long-period analysis; (5) a WWSSN short-period preshocks was lower than that of consecutive instrument response. The short-period data are aftershocks. There is no systematic change in not totally compatible with the long-period solu- spectral shape between preshocks and after- tion. Figure 11 shows the short-period records, shocks near the hypocenters. which are much more complex than the long- These observations suggest that the preshocks period waves. When we use the long-period solu- of the Mt. Lewis earthquake originated from a tion, station COL is on the nodal plane and the relatively small number of highly localized re- first pulse of COL cannot be produced by this gions in comparison with the aftershocks. The focal mechanism. Also, a shorter source duration mainshock hypocenter occurred near thc base of seems to be needed. To improve the fit of the the zone of aftershock activity and a simple point short-period synthetics, we tried a large number source can produce good results for long-period of solutions including t * values less than 1.0. The P and SH waves. The aftershock activity sur- best synthetics were obtained when we used a rounds a quiet region of about 12 km 2, and the slightly different focal mechanism (qb = 174 °, 3 = aftershock area expands with time, reaching a 87.5 ° , and A = 180°), t*= 1 s, and a 1 s trape- total area of about 65 km 2. THE 1986 MT. I.EWIS, CAI.IFORNIA, t-AR'I'IIQUAKI ( 285

An advantage of studying tclcseismic signals of 2.5 km and its associated stress drop estimate of shallow earthquakes is that one can obtain a clear 9.8 MPa spans the region of aftershock activity in picture of the source mechanism by separating the first 13 rain, but still implies very limited the reflected phases, such as pP, sP, and sS, from overlap of the mainshock rupture and aftershock the direct arrivals (e.g. Bent and Helmberger, distribution. Kanamori and Anderson (1975) 1989). The local array solution does not match demonstrated that most largc intraplate earth- the observed teleseismic waveforms exactly, al- quakes have stress drops between !.0 and 10 though it does give the correct first motions. This MPa. However, Kanamori et al. (1990) studied is probably because the teleseismic long-period the Pasadena earthquake and found that it in- signals represent seismic radiation from the en- volved rupture of very strong localized asperities tire fault length, which may give a different aver- with stress drops near 100 MPa. If we comparc age mechanism from the first motions (e.g. Wal- the moment and source duration of the Mt. Lewis lace et al., 1981). There are some trade-offs in- earthquake with the moment vs. source duration volved in the modeling procedure, but our fit to plot of Cohn et al. (1982), we obtain a similar the long-period signals is very sensitive to the (poorly constrained) stress drop of 1.0-10.0 MPa. mechanism used. Unfortunately, the teleseismic This estimate is indcpendent of the choice of data do not tightly constrain the rupture dura- fault area. On the basis of the aftershock and tion, and are inadequate to attempt any finite preshock patterns and mainshock modeling, we fault inversion for comparison with the preshock infer that the Mt. Lewis event involved thc failurc and aftershock patterns. of a single strong asperity of radius from 1.5 to Our teleseismic modeling indicates a short du- 2.5 km surrounded by weaker regions of the fault ration of rupture, 3 + 2 s, but the uncertainty is which caused subsequent cxpansion of the after- rather large for constraining the rupture length. shock zone. Assuming a bilateral rupture with rupture veloc- An alternative model of fault plane hetero- ity of 3 km s- i, we estimate rupture lengths from geneity is provided by the barrier model (e.g. Aki, 6 to 30 km, with our preference being for the 1984). In this model wc would expect the main- shorter length. Applying the empirical moment- shock to have occurred in the weaker area of the magnitude relationship of Wyss and Brune (1968) fault, with the slip zone being surrounded by and Thatcher and Hanks (1973), log M 0 = 1.5Mt, unbroken strong patches. In this case, the + 16.0, wc estimate the moment of the main- preshock activity would also have occurred in the shock to be about 3.5 x 1017 N m, which is con- weaker area. However, the high stress drop esti- sistent with the CMT solution M 0 = 3.2 x 1017 N mates are inconsistent with this model unless m (Dziewonski et al., 1987) and our body wave stresses in the surrounding aftershock regions are moment of 3.9 x 1017 N m. The stress drop Ao- very high. The temporal expansion of the after- for an earthquake of moment M 0 on a circular shock zonc is also not easily reconciled with the fault of radius a is given by Ao-= 7Mo/16a 3 barrier model. One possibility is that the absence (Eshelby, 1957; Keilis-Borok, 1959). Applying this of aftershocks in the central region is not indica- expression, we obtain estimates of the stress drop tive of enhanced slip. Yet another model is that for the mainshock of approximately 46 or 9.8 nonlinear frictional behavior causes faulting com- MPa if we use rupture arca radius values of plexity without any physical variations on the a = 1.5 km or a = 2.5 km, respectively. The surface. In this case, there may be no clear rela- smaller rupture radius of 1.5 km and the 46 MPa tionship between aftershocks and mainshock slip. stress drop estimate are appropriate only if no The preshock cluster was concentrated in a aftershocks occurred in the mainshock rupture region where aftershocks are nearly absent (Fig. area. In this case, the aftershocks in the first 13 4(b)). This may suggest a complex internal struc- min (Fig. 5(b)) occur on the edges of the rupture. ture of the main asperity in which somc points Most of the preshocks would have occurred within are weaker than others and preslip weakening the mainshock rupture zone. The larger radius of can occur. The rotations of some of the preshock 286 V. ZHOU ET AI.. fault mechanism nodal planes (Fig. 3) also sug- differ significantly from that of nearby after- gest that the local stress field in and around the shocks. This may imply that waveform variability main asperity has a complex internal structure, is more sensitive to the presence of asperities perhaps associated with localized undulation of than is the spectral content. The mainshock origi- the fault plane. The cross-correlation technique nated near the base of the seismogenic zone does not resolve the small rotation in strike very defined by preshock activity and ruptured upward well, although direct inspection of the waveforms and bilaterally along strike. The cluster of confirms the first-motion results. This is because preshock activity falls within a central quiet por- we choose the maximum value of the cross-corre- tion of the aftershock zone with a 1.5-2.5 km lation function over a range of lags corresponding radius. Tcleseismic data indicate that the primary to plus or minus one-quarter of the record length energy release in the Mt. Lewis earthquake oc- in our cross-correlation analysis. Most of the en- curred at about 9 km depth. A simple point ergy in the 15 s records is contained in the S-wave source model produced good results for long- train, desensitizing the correlations to small S/P period P and SH waves, and the best-fitting fault amplitude ratio variations. plane solution has a strike of 172 °, a dip of 83 °, Our data do not support the second prediction and a rake of 181 °. The teleseismic long-period of the asperity model: higher stress drops for SH moment is 3.9 x l017 N m. The aftershock preshocks. We did not find evidence for the fre- area expands with time along strike. We explain quency content of preshocks being significantly the observations by the failure of a strong asper- different from that of nearby aftershocks. How- ity on a fault surrounded by weaker regions. This ever, the similarity of the spectra observed for provides a qualitative explanation for the concen- preshocks and aftershocks in the vicinity of the tration of preshock activity, the gap in the after- Mt. Lewis mainshock does not necessarily imply shock distribution, and the inferred high stress that the source spectra of these events are identi- drop of the mainshock. cal if propagation effects homogenize the signals. Broader band data may yet reveal some differences. Pechmann and Kanamori (1982) also found Acknowledgments no systematic temporal change in spectral shape. We thank Terri Hauk for her extensive help with collection of the waveform data. Dave Op- 6. Conclusion penheimer provided helpful comments on a pre- vious draft of the manuscript. This research was We have used the master event relocation supported by a Regents' Fellowship of UCSC, technique to study the preshock and aftershock NSF grant EAR 8904707, 9002704 (to K.C.M.), sequence of the 1986 Mt. Lewis earthquake. and 9104764 (to T.L.). This paper is Contribution Preshock activity increased beginning 17 months No. 165, Institute of Tectonics and the C. F. before the mainshock. The waveforms of these Richter Seismological Laboratory, University of preshocks strongly suggest that these events origi- California, Santa Cruz. nated from a very small area in the immediate proximity of the hypocenter of the mainshock. Although similar events were found during the early aftershock period (about 1 day after the References mainshock), the aftershock waveforms are more variable from one event to the next. These obser- Aki, K., 1984. Asperities, harriers, characteristic earthquakes vations are plausibly explained by the asperity and strong motion prediction. J. Geophys. Res., 89: 5867- 5872. model in which the preshocks presumably reflect Bent, A.L. and Helmberger, D.V., 1989. Source complexity of a build-up of stress on the mainshock asperity. the October 1, 1987, Whittier Narrows Earthquake. J. The frequency content of the preshocks does not Geophys. Res., 94: 9548-9556. THE 1986 MT. I.EWIS, CALIFORNIA. EARTHOUAKE 287

Beroza, G.C., 1991. Near-~urce modeling of the Loma Prieta large earthquakes. Am. Geophys. Union, Maurice Ewing earthquake: evidence for heterogeneous slip and implica- Series, 4: 1-19. tions for earthquake hazard. Bull. Seismol. Soc. Am., 81: Kanamori, It. and Anderson, D.L., 1975. Theoretical basis of 1603-1621. some empirical relations in seismology. Bull. Seismol. Soc. Cockerham, R.S. and Eaton, J.P., 1987. The earthquake and Am., 65: 1(173-1095. its aftershocks, April 24 through September 30, 1984. U.S. Kanamori, H., Mori, J. and Heaton, T.H., 1990. The 3 De- Geol. Surv. Bull., 1639: 15-28. cember 1988, Pasadena earthquake (M t = 4.9) recorded Cohn, S.N., Hong T.-L. and Helmberger, D.V., 1982. The with the very broadband system in Pasadena. Bull. Seis- Orcwille earthquakes: a study of .source characteristics and tool. Soc. Am., 80: 483-487. site effects. J. Geophys. Res., 87: 4585-4594. Keilis-Borok, V., 1959. On estimation of the displacement in Doser, D.I., 1989. Foreshocks and aftershocks of large (M 5.5) an earthquake source and of source dimensions. Ann. earthquakes within the western cordillera of the United Geofis., 12: 205-214. States. Bull. Seismol. Soc. Am., 80: 110-128. Klein, F.W., 1985. User's Guide to HYPOINVERSE, a pro- Dzicwonski, A.M., Ekstrom, G., Franzen J.E. and Wood- gram for VAX computers to solve for earthquake location house, J.H., 1987. Centroid-moment tensor solutions for (6/88 version). U.S. Geol. Surv. Open-File Rep., 88-000, January-March 1986, Phys. Earth Planet. Inter., 45: 1-10. 1-50. Engdahl, E.R., Billington S. and Kisslinger, C., 1989. Teleseis- Langston, C.A. and Helmberger, D.V., 1975. A procedure for mically recorded seismicity before and after the May 7. modelling shallow dislocation sources. Geophys. J.R. As- 1986, Andreanof Islands, Alaska, earthquake. J. Geophys. tron. Soc., 42: 117-130. Res., 94: 15481-15498. McKenzie, M.R. and Uhrhammer, R.A., 1986. Bulletin of the Eshelby, J.D., 1957. The determination of the elastic field of Seismographic Stations of the University of California, an ellipsoidal inclusion and related problems. Proc. R. University of California, Berkeley, CA, 56, pp. 1-105. Soc. London, Ser. A, 241: 376-396. McNally, K.C., 1977. Patterns of earthquake clustering pre- Geller, R.J. and Mueller, C.S., 1980. Four similar earthquakes ceding moderate earthquake, Central and southern Cali- in central California. Geophys. Res. Left., 7: 821-824. fornia (abstract). EOS, 58:1195. Hartzell, S. and lida, M., 1990. Source complexity of the 1987 Mendoza, C. and Hartzell, S.H., 1988. Aftershock patterns Whittier Narrows, California, earthquake from the inver- and mainshock faulting. Bull. Seismol. Soc. Am., 78: 1439- sion of strong motion records, J. Geophys. Res., 95: 1449. 12475-12485. Pechmann, J.C. and Kanamori, H., 1982. Waveforms and Hartzell, S.H., Stewart G.S. and Mendoza, C., 1991. Compari- spectra of preshocks and aftershocks of the 1979 Imperial son of LI and L2 norms in a teleseismic waveform inver- Valley, California, Earthquake: evidence for fault hetero- sion for the slip history, of the Loma Prieta, California, geneity. J. Geophys. Res., 87: 10579-10597. earthquake. Bull. Seismol. Soc. Am., 81, 1518-1539. Pechmann. J.C. and Thorbjarnardottir, B.S., 1990. Waveform tlauk, F.T., 1985. 1984 earthquake catalog for the Livermore analysis of two preshock-mainshock-aftershock sequences Valley region, Lawrence Livermore National Laboratory, in Utah. Bull. Seismol. Soc. Am., 80: 519-550. Livermore, CA, pp. 1-39. Reasenberg, P. and Ellsworth, W.L., 1982. Aftershocks of the Houston, H. and Engdahl, E.R., 1989. A comparison of the Coyote Lake, California, earthquake of August 6: 1979: a spatio-temporal distribution of moment release for the detailed study. J. Geophys. Res., 87: 10637-10655. 1986 Andreanof Islands earthquake with relocated seis- Reasenberg, P.A. and Oppenheimer, D., 1985. FPF1T, FP- micity. Geophys. Res. Lett., 16: 1421-1424. PLOT and FPPAGE: FORTRAN computer programs for Ishida, M. and Kanamori, H., 1977. The spatio-temporal calculating and displaying earthquake fault-plane solu- variation of seismicity before the 1971 San Fernando tions. U.S. Geol. Surv. Open-File Rep., 85-739:109 pp. earthquake. Geophys. Res. Lett., 4: 345-346. Schwartz, S.Y., Dewey J.W. and Lay, T., 1989. Influence of Ishida, M. and Kanamori, H., 1978. The foreshock activity of fault plane heterogeneity on the seismic behavior in the the 1971 San Fernando earthquake, California. Bull. Seis- Southern Kuriles Islands arc. J. Geophys. Res., 94, 5637- mol. Soc. Am., 68: 1265-1279. 5649. Ishida, M. and Kanamori, H., 1980. Temporal variation of Smith, K.D. and Priestley, K.F., 1988. The foreshock sequence seismicity and spectrum of small earthquakes preceding of the 1986 Chalfant, California, earthquake. Bull. Seis- the 1952 Kern County, California, earthquake. Bull. Seis- mol. Soc. Am., 78: 172-187. mol. Soc. Am., 70: 509-527. Smith, S.W., Ehrenberg J.E. and Hernandez, E., 1982. Analy- Ishida, M. and Ohtake, M., 1984. Seismicity and waveforms of sis of the El Centro differential array for the 1979 Imperial the microearthquakes before and after the Shizuoka-Seibu Valley earthquake. Bull. Seismol. Soc. Am., 72, 237-258. earthquake, central Japan. Bull. Seismol. Soc. Am., 74: Tajima, F. and Kanamori, H., 1985a. Aftershock area expan- 605-620. sion and mechanical heterogeneity of fault zone with sub- Jobnson, C.E. and Hadley, D.M., 1976. Tectonic implications duction zones. Geophys. Res. Lett., 12(6): 345-348. of the Brawley earthquake swarm, Imperial valley, Califl~r- Tajima, F. and Kanamori, H., 1985b. Global survey of after- nia. Bull. Seismol. Soc. Am., 66: 1133-1144. shock area expansion patterns. Phys. Earth Planet. Inter., Kanamori, H., 1981. The nature of seismicity patterns before 40: 77-134. 288 Y. ZHOU ET AL.

Thatcher, W. and Hanks, T.C., 1973. Source parameters of Wyss, M. and Brune. J.N., 1968. Seismic moment stress and southern California earthquakes. J. Geophys. Res., 78: source dimensions for earthquakes in the California- 8547-8576. Nevada Region. J. Geophys. Res., 73: 4681-4694. Thorbjarnardottir, B.S. and Pechmann, J.C., 1987. Constraints Zhou, Y., 1992. A comprehensive study of the 1986 Mt. Lewis, on relative earthquake locations from cross-correlation of California earthquake: preshock sequence-mainshock- waveforms, Bull. Seismol. Soc. Am., 77: 1626-1634. aftershock sequence. M.S. Thesis, University of California Wallace, T.C., Helmberger D.V. and Ebel, J.E., 1981. A at Santa Cruz, 80 pp. broadband study of the 13 August 1978 Santa Barbara Zufiiga, F.R., Wys.s M. and Wilson, M.E., 1987. Apparent earthquake. Bull. Seismol. Soc. Am., 71: 1701-1718. stresses, stress drops, and amplitude ratios of earthquakes Wong, I.G., 1991. An evaluation of reservoir-induced seismic- preceding and following the 1975 Hawaii M~ = 7.2 main ity in the eastern San Francisco Bay region, California. shock. Bull. Seismol. Soc. Am., 77: 69-96. Seismol. Res. Lett., 62, 52 pp.