Results associated with systematic analysis of such signals at several large strike-slip fault zones can be summarized as follows: The observed fault zone trapped waves at several fault and rupture zones are generated by relatively shallow structures that extend generally only over the top ~3-4 km of the crust [ Ben-Zion et al., GJI, 2003 ; Peng et al., GJI, 2003 ; Lewis et al., GJI, 2005 ]. The shallow trapping structures are characterized by about 100m-wide fault zone layers with strong reduction of seismic velocity (e.g., 30-40%) and strong attenuation (e.g., Q values of 10-30). The shallow trapping structure in the North Anatolian fault is surrounded by broader anisotropic and scattering zones that are also confined primarily to the top 3 km [ Peng and Ben-Zion, GJI, 2004 ]. The shallow structure responsible for generating both fault zone trapped waves and anisotropic effects may be related to the top part of a flower-type structure of the fault zone. It contains highly damaged materials with intense microcracks and resides above the active portion of seismogenic crust where earthquakes nucleate. Such structure significantly increases the seismic shaking hazard near the fault zone, since seismic sources external to the fault zone are capable of generating strong motion amplification in the shallow trapping structure.

My recent work in this direction focuses on using fault zone head waves to image the fault interface at seismogenic depth. Major plate boundary faults such as the San Andreas fault often juxtapose blocks with different elastic properties. A sharp material contrast is expected to refract seismic energy along the interface. Fault zone head waves (FZHWs) are generated by slip along the material interface and recorded by near-fault stations on the slower block as emergent first arriving waves with opposite polarity from that of the following direct P waves [Ben-Zion and Malin, 1991; McGuire and Ben-Zion, 2005 ]. Since FZHWs spend most of their times propagating along the fault interface, waveform analysis based on such phases provides high-resolution information of fault zone structure at seismogenic depth.

Together with my Ph.D. advisor Yehuda Ben-Zion , we have systematic analyzed spatio-temporal changes in anisotropy [ Peng and Ben-Zion, GJI, 2005 ] and seismic velocities [ Peng and Ben-Zion, PAGEOPH, 2006 ] around the North Anatolian fault using repeating earthquake clusters in the aftershock zones of the 1999 Mw7.4 Izmit, and Mw7.1 Duzce, Turkey, earthquake sequences. Our results do not show precursory temporal evolution of properties before the Duzce mainshock. The anisotropy results show small co-seismic changes. However, the scattering results show clear co-seismic changes and post-seismic logarithmic recovery. The observed co- and post-seismic effects are likely to reflect mostly changes in properties of the top most section of the crust in response to strong ground motions of the nearby major earthquakes. These results provide strong support that damage, or nonlinearity, is widespread in the shallow crust during strong shakings, and may help to explain fault interaction and earthquake triggering by seismic waves.

Together with my postdoctoral advisor John Vidale and Professor Heidi Houston (formerly at University of California, Los Angeles, now at University of Washington, Seattle), Professor Miaki Ishii at the Harvard University, and Dr. Agnes Helmstetter at LGIT University Joseph Fourier, France, we have conducted two projects on deciphering the seismicity rate around main shocks from high-frequency waveforms. Our targets include 82 M3-5 earthquakes in Japan recorded by the Hi-Net array [ Peng et al., JGR, 2007, in press ], and the 2004 Mw6.0 Parkfield earthquake [ Peng and Vidale, GRL, 2006 ]. In summary, we observe that the seismicity rate immediately before and after the main shock is less than predicted from the Omori's law with c = 0 and the long-term p value. Our results can be explained by the epidemic-type aftershock sequence model, and the rate-and-state model of Dieterich [1994]. Alternatively, non-seismic stress changes near the source region, such as transient aseismic slip or pore fluid pressure fluctuations, may share responsibility for the non-Omori behavior immediately before and after the main shock.