Introduction

On Friday, October 30, 1998, at 1:53 AM local time (9:53:31 UT), an earthquake of magnitude M_w 4.9 (UNR) took place near the California-Nevada state border. The epicenter of the event is 7 km north-northwest of Incline Village, Nevada and the north shore of Lake Tahoe (39° 18.14'N, 119° 58.76'W) [see Figures 1 and 3]. A focal depth of 10.5 ± 0.3 km and strike-slip focal mechanism suggest it is similar to other recent eastern Sierra Nevada-Western Basin and Range Extensional Province earthquakes (Rogers et al., 1991; Ichinose et al. 1998a). The earthquake occurred 4 km northwest of the surface trace of the northeast-striking North Tahoe-Incline Village fault zone, which is the largest fault zone in the area (Figure 3). This is a system of southeast dipping normal faults, so the earthquake is well within the footwall of this fault zone. The Seismological Laboratory received felt reports, from which we assigned the following Modified Mercalli Intensities: from Kings Beach (VI), Incline Village (VI), Truckee (V), Reno (V), Sparks (V), and Carson City (V). A few felt reports were also received from Nevada cities as far away as Virginia City, Fernley (IV), Dayton (III), Minden (III), Glenbrook (V) and Cold Springs (III) and California cities of Placerville (IV), Vinton (IV), Twain Harte (IV), and San Andreas (IV).

The purpose of this paper is to report on the source parameters and seismotectonics of the Incline Village earthquake and its ground-motion effects in the city of Reno. A companion paper in this issue by Schweickert and others reports on the geology and recent fault mapping in the Lake Tahoe basin.

Historical and Modern Seismicity

The Incline Village earthquake occurred near the western edge of the Reno-Carson City urban corridor, which is the most seismically active part of Nevada. Thirteen earthquakes of M 6 or greater have occurred since 1850 in and near this area (dePolo et al., 1997). The 1914 M 6.7 and 1948 M 6 Verdi, Nevada earthquakes (Rogers et al., 1991) and the 1966 M 6 Truckee, California earthquake (Greensfelder, 1968) occurred approximately 30 km north and 10 km northwest of the 1998 Incline Village epicenter (Figure 2). The largest of the nearby historical earthquakes were the 1860 M 7 Pyramid Lake and the 1869 M 6.7 Virginia Range earthquakes (Toppozada et al., 1981; dePolo et al., 1997). Surface faulting from a recent earthquake was recognized by Sanders and Slemmons (1979) 50 km northeast from the 1998 epicenter on the Olinghouse fault. Although Sanders and Slemmons associated this with the 1869 earthquake, dePolo et al. (1997) suggest it may be associated with the 1860 earthquake instead.

Figure 2 shows the locations of earthquakes from the University of Nevada Reno (UNR) Seismological Laboratory catalog from 1852 through 1998. The Olinghouse lineament is associated with a weak northeast trend in seismicity, previously recognized by Martinelli (1989), which extends southwestward from the Olinghouse fault through south Reno and the Steamboat Springs geothermal area, and possibly to the North Tahoe-Incline Village fault zone (Figures 2 and 3). The location of modern seismicity near the North Tahoe-Incline Village fault zone is well constrained from coverage by several nearby permanent short period stations, with the nearest one from the Incline Village earthquake at Martis Peak (MPK), only 5 km epicentral distance, and 10 other sites within 50 km radius of the earthquake in operation by UNR since the 1970's.

Aftershocks

The aftershocks of the 1998 Incline Village earthquake are aligned on a linear trend northeast of the mainshock epicenter, consistent with the directivity analysis (Figure 3). There are also nearby northeast striking faults previously recognized by Lewis (1988) that parallel the aftershock activity. The actual aftershock decay rate dropped off quickly, with 11 located aftershocks after the first week and only one M >= 3.

An aftershock forecast by the USGS/UCB predicted 5 to 20 aftershocks with M >= 3 within the the first week. These aftershock forecasts are based on statistics of aftershocks typical for California. Previous aftershock sequences within the northern Sierra Nevada associated with the 1979 M 5.3 Fort Sage, CA, 1980 M 5.0 Donner Pass, CA, and 1995 M 4.5 Border Town, NV, earthquakes all behaved similarly, each with fewer than 25 aftershocks within 1 to 3 weeks and only one 3.0 < M < 3.8 aftershock occurring within the first week. The magnitude difference between the mainshock and largest aftershock was also larger than 2 magnitude units in these aftershock sequences rather than 1.0 ± 0.47 unit difference for most Western US sequences (Doser, 1989). The 1980 and 1995 aftershock sequences were well covered with at least 1 station under 20 km epicentral distance, resulting in an estimated detection threshold of about magnitude 2. This suggests that recalibration of the Omori aftershock parameters using local aftershock sequences might significantly improve forecasts of future aftershock probabilities.

Waveform Analysis

Focal mechanisms for the Incline Village mainshock, a foreshock and an aftershock (Figure 4) were obtained by fitting waveforms using a nonlinear inversion process. The inversion process proves effective at regional distances and can be used for single or multiple station solutions with well-calibrated Green's functions (Ichinose et al., 1998b). The synthetics were generated using a fast f-K summation technique (Zeng and Anderson, 1995).

The mainshock is modeled from 100 to 2 seconds period, depending on the epicentral distance to the recording station. UNR digital broadband stations and University of California, Berkeley (UCB) Digital Seismic Network stations (Romanowicz et al., 1992; 1994) are used in the waveform optimization. UNR operates 3 digital broadband stations in northwestern Nevada and 23 in southern Nevada. Sites contain Guralp CMG40 sensors sampling at 100 sps. Figure 1 shows the locations of the stations and Figure 4 shows the waveform fits. We used a simple 32 km thick crust ( V_p = 5.9 km/s, V_s = 3.4 km/s, rho = 2.7 g/cm³ ) over mantle ( V_p = 7.8 km/s, V_s = 4.5 km/s, rho = 3.3 g/cm³ ) as reasonable P- and S-wave velocities and densities for this region. We also optimized for centroid depth using 2 km depth increments (Figure 4) and find the lowest misfit at 8 km. The centroid depth is 2.5 km shallower than the focal depth obtained from our UNR network but both of these locations do not take into account station elevation and so the both depths may be shallower. The focal depth may also reflect the point of rupture initiation while the centroid depth represents the overall center of the fault plane rupture. The east-west T-axis of the mainshock focal mechanism is in agreement with other studies of this region (Rogers et al., 1991; Ichinose et al., 1998a).

We forward modeled an Oct. 27 04:53(UT) foreshock and an Oct. 30 10:11(UT) aftershock which were only recorded at station WCN, the nearest 3-component station. Different velocity models were needed because of the smaller size of these events relative to the mainshock, requiring the addition of higher frequencies of up to 2 Hz. A simple two-layer-model cannot explain the complexity of the waveforms at these frequencies. We were able to fit the foreshock and aftershock initially with the same mechanism as the mainshock but obtained slightly better fits by slightly varying the fault geometry.

Directivity, Slip History Analysis, and Stress Drop

The focal mechanisms obtained from first motions, waveform analysis or moment tensor inversion cannot resolve the fault geometry and sense of slip without surface rupture information because of the fault plane and the auxiliary fault plane ambiguity. We therefore performed a directivity analysis using the empirical Green's function technique to deconvolve out the source time function (STF) by removing the site and path effects (e.g., Mueller, 1985). The technique takes advantage of a pair of colocated events where the seismogram from the small earthquake can be considered as the response of the medium to a point source. When this is deconvolved from the larger seismogram, one obtains an apparent STF of the larger earthquake. The STFs at different stations may vary depending on the azimuth of observation (Frankel et al., 1986). The principle of the Doppler effect is used to interpret the apparent STFs, where a shorter pulse duration means that the source is propagating toward the receiver and a broader pulse duration means that the source is propagating away from the receiver.

We used the M_w 3 aftershock which occurred at 10:11(UT) as the empirical Green's function and deconvolved it from the M_w 4.9 mainshock, which occurred at 9:53(UT), to estimate the apparent STF at 6 stations. The epicentral location and depth of this earthquake pair are each less than 1 kilometer apart, and the aftershock was large enough to be recorded at 6 sites with good azimuthal coverage. The P-waves were windowed and used in the time-domain deconvolution following Ichinose et al. (1997), and the resulting apparent STFs were band-passed within the frequency range of 0.5 to 10 Hz. We fitted a line along the steepest slopes of the apparent STFs and projected the lines to the time-axis. The source duration was estimated from the interval between the crossings of the projected lines and the time-axis, similar to the "narrow" pulse width picks of Hough and Dreger (1996).

The duration of the apparent STFs are listed in Table 1. They are narrower at stations PAH and WVY, which straddle the nodal plane that strikes N33°E. We interpret from the pulse widths of the STFs that the earthquake occurred along this fault plane and ruptured towards the northeast. A northeast oriented fault is consistent with the northeast trend of eleven aftershocks and several northeast striking fault segments originally mapped by Lewis (1988) which have average azimuths of N20-35°E (see Figure 3).

The dynamic stress drop $dsd$ is estimated for the mainshock using the relationship of Eshelby (1957) for a circular rupture

dynamic stress drop = 7 M_o / 16 r³

where M_o is the seismic moment, which is 2.78x10²³ dyne cm from the waveform analysis, and r is the source radius. The apparent STF duration tau, adjusted for directivity, can be used to determine the the source length $L$ following Boatwright (1980)

L = { tau c } over { 1 - [ v/c cos ( theta ) ] }

where v is the rupture velocity, c is the wave velocity, and theta is the angle between the fault plane and the outgoing ray. We set c equal to the P-wave velocity of 6 km/s and assume a constant rupture velocity of 3.06 km/s, 0.9 times the shear wave velocity at the source. To determine $theta$, we assume the source ruptured along a strike of N33°E and use tau values listed in Table 1 to determine the source length. We then divided the individual source lengths by 2 for the source radii. This results in a dynamic stress drops for each STF listed in Table 1 with a mean stress drop of 723 bars and a ± 1 sigma scatter of 468 bars. Static and dynamic stress drops of Basin and Range earthquakes mainly fall within the 10 to 100 bars range (e.g., Ichinose et al., 1997) so the the $mw$ 3 empirical Greens function used in this analysis may of been affected by attenuation which would tend to increase the dynamic stress drop. For an unilateral rupture, the source radius can also vary by a factor of 2 from directivity, causing a factor of 10 uncertainty in stress drop. For this earthquake, stations WCN and BEK sees two peaks on the STF (Figure 5.) and may be resolving a subevent or a 2D propagation effect but these details in the rupture process do not change our interpretation of the overall rupture direction.

Ground Motion Effects

The mainshock was recorded at station SF (Figure 1) on a Kinemetrics FBA11 accelerometer located 29.3 km from the epicenter. The sensor is located on the foundation of the historical UNR Mackay School of Mines building which sits atop Quaternary basin sediments over andesite bedrock. At this location, the basin thickness is at least 177 m based on a well (Garside and Schilling, 1979) that did not penetrate bedrock. From gravity modeling assuming a 0.46 g/cm³ density contrast, Abbott and Louie (1998) estimated a basin thickness of 300 m. A horizontal peak ground acceleration (PGA) of 0.036g was recorded with a S-wave coda duration of approximately 3 seconds (Figure 6) at station SF. The S-wave coda acceleration spectrum is flat from 1 to 8 Hz. Empirical attenuation relationships between horizontal PGA and distance by Campbell (1997) and Sadigh et al. (1997) estimate that for a $mw$ 4.9 strike-slip earthquake at 30.3 km hypocentral distance, the horizontal PGA will be about 0.027g and 0.028g respectively. Spudich et al. (1997) derived an empirical relation for extensional tectonic regimes which predicts a 0.037g PGA.

Conclusions

From waveform modeling, rupture directivity analysis, and the aftershock distribution, the October 30, 1998 M_w 4.9 Incline Village, Nevada earthquake occurred on a N33°E striking high-angle strike-slip fault, probably rupturing to the northeast toward station PAH. This result is also supported by geologic mapping of the area and places the earthquake 3 km northwest of the active northeast-striking North Tahoe-Incline Village fault zone (Schweickert et al., this issue). The mainshock stress drop, estimated from P-wave source durations, is 723 bar with a ± 1sigma of 468 bars. A horizontal PGA of 0.036 g was recorded at one location in Reno.

Little can be inferred with confidence from this earthquake sequence about the seismic hazard for the Reno-Carson City urban area. However, the combination of the results of this and other studies suggests an hypotheses that may be worth some consideration. The Oct. 30 earthquake is part of a diffuse seismicity trend (Martinelli, 1989) that extends northeast from Lake Tahoe into the southern Truckee Meadows, which includes the Reno urban area. Schweickert and others (companion paper) show an active zone of faulting more or less coinciding with this trend. Abbott and Louie (1998) found a negative gravity anomaly striking northeast across the southern Truckee Meadows, also more or less coincident with the seismicity trend. East of the Truckee Meadows and offset to the north, the left-lateral Olinghouse system strikes northeast for approximately 30 km. Thus, there is a possibility for an active zone of deformation striking northeast through the southern Truckee Meadows and possibly even further northeast. This hypothesis is different from the usual model of the Truckee Meadows as a graben controlled by normal faults on both the east and west sides (e.g., dePolo et al, 1997) although both modes of faulting are consistent with the east-west extensional stresses that previous studies have proposed for this area.

Acknowledgments

We thank the Berkeley Seismological Laboratory, University of California, Berkeley, for contributing waveform data from their BDSN Broadband Network. This work was funded by the U.S. Geological Survey-National Earthquake Hazard Reduction Program grant, contract number 1434-94-G-2479.

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Data