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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 112, B06315, doi:10.1029/2006JB004611, 2007

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Structure of the California Coast Ranges and San Andreas Fault at SAFOD from seismic waveform inversion and reflection imaging Florian Bleibinhaus,1,4 John A. Hole,1 Trond Ryberg,2 and Gary S. Fuis3 Received 3 July 2006; revised 23 December 2006; accepted 23 February 2007; published 28 June 2007.

[1] A seismic reflection and refraction survey across the San Andreas Fault (SAF) near

Parkfield provides a detailed characterization of crustal structure across the location of the San Andreas Fault Observatory at Depth (SAFOD). Steep-dip prestack migration and frequency domain acoustic waveform tomography were applied to obtain highly resolved images of the upper 5 km of the crust for 15 km on either side of the SAF. The resulting velocity model constrains the top of the Salinian granite with great detail. Steep-dip reflection seismic images show several strong-amplitude vertical reflectors in the uppermost crust near SAFOD that define an 2-km-wide zone comprising the main SAF and two or more local faults. Another prominent subvertical reflector at 2–4 km depth 9 km to the northeast of the SAF marks the boundary between the Franciscan terrane and the Great Valley Sequence. A deep seismic section of low resolution shows several reflectors in the Salinian crust west of the SAF. Two horizontal reflectors around 10 km depth correlate with strains of seismicity observed along-strike of the SAF. They represent midcrustal shear zones partially decoupling the ductile lower crust from the brittle upper crust. The deepest reflections from 25 km depth are interpreted as crust-mantle boundary. Citation: Bleibinhaus, F., J. A. Hole, T. Ryberg, and G. S. Fuis (2007), Structure of the California Coast Ranges and San Andreas Fault at SAFOD from seismic waveform inversion and reflection imaging, J. Geophys. Res., 112, B06315, doi:10.1029/ 2006JB004611.

1. Introduction [2] As a part of EarthScope, the San Andreas Fault Observatory at Depth (SAFOD) will sample and monitor the San Andreas Fault (SAF) at seismogenic depth in the source region of a repeating magnitude 2 earthquake near the town of Parkfield, CA [Hickman et al., 2004]. Several active and passive geophysical surveys were conducted in that region to obtain a detailed subsurface characterization. Thurber et al. [2003, 2004b], Nadeau et al. [2004], and Malin et al. [2006] used seismic observations to improve existing velocity models and to better constrain the target positions. From a shallow high-resolution seismic profile at the proposed drill site conducted in 1998, a compressional wave velocity model and a reflection seismic image were derived [Catchings et al., 2002; Hole et al., 2001]. The reflection image shows two steep reflectors in the vicinity of the SAF surface trace. However, the penetration depth of this survey was limited to 1 km, and the possible continuation of these reflectors to seismogenic depths remained speculative. In order to obtain seismic velocity

and steep-dip reflection images down to 5 km depth, a reflection and refraction seismic profile was acquired in 2003 across the SAF at the site of the SAFOD borehole. The profile trends northeast, from the Salinian terrane across the SAF through the Franciscan accretionary complex to the Great Valley Sequence (GVS; Figure 1). The Salinian block is composed dominantly of Cretaceous granites overlain by Tertiary marine sedimentary rocks. Juxtaposed against the granite by the SAF, the late Cretaceous to early Tertiary Franciscan subduction complex consists of a heterogeneous me´lange of mainly greywacke, shale, greenstone, and chert. It is separated from the quasi-coeval sedimentary rocks of the Great Valley Sequence (GVS) by the Waltham Canyon Fault (WCF). The WCF is a part of the larger terrane bounding Coast Range Fault, and it locally splays eastward relative to its regional trend [Dibblee, 1971; Page et al., 1998]. This paper presents highly resolved waveform tomography velocity models and steep-dip reflection seismic images of the upper crust from the SAFOD 2003 data and a low-resolution reflection image of the whole crust.

2. Data 1

Department of Geosciences, Virginia Tech, Blacksburg, Virginia, USA. 2 GeoForschungsZentrum, Potsdam, Germany. 3 U.S. Geological Survey, Menlo Park, California, USA. 4 Now at Earth Resources Laboratory, Massachusetts Institute of Technology, Cambridge, USA. Copyright 2007 by the American Geophysical Union. 0148-0227/07/2006JB004611$09.00

[3] Sixty-three inline explosive shots were recorded at 50-m receiver spacing on a 46-km-long line (Figure 1). A detailed description of the acquisition parameters is given by Hole et al. [2006]. The average CMP coverage of the survey is 30-fold. The line had two gaps of 2- and 4-km length where landowner permissions could not be secured.

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BLEIBINHAUS ET AL.: SAFOD 2003 WAVEFORM INVERSION AND IMAGING

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Figure 1. Map of the SAFOD site (white star) and the 2003 seismic survey near Parkfield, CA. Circles are shots; coincident black line is receivers. Model coordinates (distance along the line) are labeled in kilometers. Major faults (SAF, San Andreas; BCF, Buzzard Canyon; GHF, Gold Hill; SWFZ, Southwest Fracture Zone; WCF, Waltham Canyon) from the work of Hole et al. [2006] and references therein. From SW to NE the profile crosses the Salinian terrane, the Franciscan terrane, and the Great Valley Sequence.

The dense receiver spacing provides the basis for highresolution imaging and inversion methods. However, the shot spacing is irregular and varies from 0.2 to 3 km, which poses a challenge for waveform inversion and imaging. [4] Two sample shot gathers are displayed in Figure 2. The data show strong first arrivals, but only a few strongly coherent reflections. Highly variable topography and nearsurface geology make phase correlation difficult. The strongest reflection (B) is observed 2 s after the first breaks, and it has an abnormal reverse slope. It is recorded by most shots and receivers in the GVS at the northeast end of the line. A second strong, but significantly weaker reflection (A) in the shot gathers also has a reverse slope. Such reverse moveout curves are caused by steeply dipping structures near the receiver location at which the reflection intersects the first arrivals [Hole et al., 1996]. The location and moveout of the observed reflections (A, B) are appropriate for near-vertical structures at the WCF and SAF, respectively. [5] Typical reflections from shallowly dipping structures with normal (hyperbolic) moveout (NMO) are rare. A NMO reflection from the top of the Salinian crystalline basement

is observed on several shots southwest of the SAF. However, the survey was not designed for such a shallow (1 s TWT 1 km depth) reflector, and no attempt was made to image it. More NMO-reflections are observed at 4– 9 s TWT for a few shots, but much of the near-vertical data is blurred by strong S-waves and ground roll. In general, the scarcity of clear NMO reflections indicates that the main geologic formations are not strongly layered. [6] The dominant signal frequency is 8 Hz, and the signal ranges from 3 –40 Hz (Figure 3). Data quality varies markedly due to shot size variations from 10 – 200 kg, different shot coupling, and variations of the near-surface geology. Shots in the GVS and in the Salinian close to SAFOD generated the best signal, in contrast to shots in the Franciscan. This geologic trend is also exhibited in a plot of average amplitude at receivers (Figure 4). High amplitudes in the Salinian correspond to low attenuation in the crystalline basement. Low amplitudes in the Franciscan confirm the famously strong attenuation in this sedimentary me´lange, which is attributed to its strongly deformed and heterogeneous nature [Lendl et al., 1997; Romero et al., 1997]. In addition to the lateral attenuation heterogeneity, a

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Figure 2. (a) SP 8 in the Great Valley Sequence, and (b) SP 34 in the Salinian near SAFOD (white star). The SAF surface trace is marked by a vertical line. Shot locations are indicated on Figure 1. The data are bandpass filtered (3 – 24 Hz) and scaled with time squared. Time is plotted with a linear moveout of 5 km/s. A dashed line indicates the first arrivals computed in the traveltime tomography model of Figure 7. Pg and Sg denote the waves refracted in the upper crust. Reflected refractions correlated with the SAF (A) and the Waltham Canyon Fault (B) and a reflection from the top of the Salinian granite (PSP) are marked by arrows. Multiples (PgM) from the top of the Salinian granite are observed on most shots southwest of the SAF. Note the strong amplitudes of the direct wave Pch which is channeled in the sedimentary layer over bedrock. PgM, Pch , Sg, and the ground roll R are problematic noise during migration, and partly also during waveform incersion (Sg, R).

strong decrease of attenuation with penetration depth is observed (Figure 5). The spectral amplitude ratio of near field to far field (Figure 3) reveals an additional frequency dependency of attenuation. The observed significant vertical, lateral, and spectral attenuation-variations make the use of true amplitudes difficult for waveform inversion and prestack migration.

3. Waveform Tomography [7] Waveform tomography is an inversion method, which provides velocity models with much higher resolution than

traveltime tomography. It uses waveforms and band-limited ‘‘wave paths’’ [Woodward, 1992] in place of traveltimes and raypaths, that is, it is a fully dynamic method. The wave path kernels can be computed using the adjoint wavefield. Tarantola [1984] has shown for the acoustic approximation that the gradient of the waveform misfit function can be obtained from the convolution of the forward propagated source with the backpropagated residual (Figure 6). This approach reduces the cost of the inverse problem to twice the cost of the forward problem. The gradient is used for an iterative steepest-descent inversion, which requires another forward computation to derive a step length [Pratt, 1999].

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Figure 3. Mean first arrival amplitude spectra at 2 – 4 km offset (dashed), 10– 12 km (dash-dotted), and 20– 22 km (dotted). Their relative amplitude corresponds to the offset trend displayed in Figure 5. The solid line depicts the spectral ratio of near field (2 – 4 km) to far field (20 – 22 km). Its slope is proportional to f 2.1 indicative of a strong frequency dependency of signal attenuation. Tarantola [1984] also pointed out that this method can use all phases of the seismic wavefield (direct waves, reflections, multiples, surface waves) if they are taken into account by the forward modeling scheme. The strength of the gradient method lies nevertheless in the inversion of refractions and wide-angle reflections. To fully exploit the resolving power of near-vertical reflections or multiples, higher-order derivatives would have to be taken into account, which are, however, computationally much more expensive [Pratt et al., 1998].

Figure 4. Relative receiver amplitudes along the profile in the following three frequency bands: 4– 8 Hz (dashed), 8 – 12 Hz (solid), and 12– 16 Hz (dotted). RMS amplitudes were computed in a 1-s window around the first break of the t2-scaled data, and the average was calculated from sourcereceiver offsets of 5 – 25 km. The frequency-bands were normalized for a white source spectrum, such that differences between them reflect the subsurface structure rather than the source signature. The long wavelength trend of the curves is similar, and it is correlated to the main geologic units, exhibited as low amplitudes in the Franciscan versus high amplitudes in the Salinian and the GVS. The weak signal in the Salinian at km 18– 20 has no obvious surficial correlation. The short wavelength variations are caused by variations of the near-surface geology and coupling effects. Attenuation is strongest for the high frequencies, particularly in the Franciscan.

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Figure 5. Logarithmic plot of RMS first-break amplitudes after surface consistent shot and receiver amplitude corrections (solid) and after additional t2 scaling (dotted) versus offset. Synthetic RMS amplitudes were computed for different constant Q models and the two-dimensional velocity model of pFigure 7, and they were scaled subsequently with 1/ t to approximately correct for threedimensional spherical divergence (dashed lines). A comparison of the slopes yields a rough estimate of Q for the data. Up to 10 km offset (corresponding to 1 – 2 km depth), a Q of 10– 20 fits the data. This changes abruptly at greater offsets (greater penetration depths), where damping seems to become insignificant. Beyond 25 km offset, noise dominates the amplitudes. Note that signal can still be observed at long offsets by trace correlation despite a signalto-noise ratio near or less than 1.

Figure 6. Gradient method of waveform inversion. The gradient of the misfit function (bottom) is obtained by combining the wavefield of a forward propagated source (top) with the backpropagated residual (middle). The (frequency domain) wavefields were computed for the 4 Hz source and residual of shotpoint 34 and a receiver at km 42 (see Figure 1 for location) in the model of Figure 7. The total gradient is the stack of the individual gradients of all source-receiver pairs. In the actual inversion, the residual of all receivers is backpropagated simultaneously for each shot. Note the similarity of the gradient to Woodward’s [1992] wave paths. The main difference is that the wave path is a sensitivity kernel, while the gradient includes the residual. For more details see, for example, Pratt and Shipp [1999].

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Figure 7. Velocity model ‘‘a’’ of Hole et al., [2006, Figure 3a] derived by minimum-structure firstarrival traveltime tomography of the 2003 data. A dashed line is drawn below the surface trace of the SAF. The solid line marks the SAFOD drill hole after completion of the phase 2 drilling in 2005. The model shows strong vertical and lateral variations from sedimentary rock velocities of 2 – 4 km/s to granitic basement velocities of 5 – 6 km/s in the Salinian. The strong velocity contrast across the SAF reflects the juxtaposition of this granite against Franciscan sedimentary rocks. [8] In order to converge to the global minimum, gradient methods require the initial misfit function to be in the same valley. This means for real band-limited data that the long wavelengths of the starting model must be constrained by a more robust method. In most practical cases, traveltime tomography can provide an appropriate background model [Operto et al., 2004; Pratt, 1999; Pratt and Shipp, 1999; Pratt et al., 1996; Song et al., 1995]. Therefore waveform inversion is not an alternative to traveltime tomography, but a subsequent step to further increase model resolution. For this study, a frequency domain waveform inversion method of Pratt [1999] was used. The most important advantage of working in the frequency domain is the ability to further decompose the inversion process. The separate inversion of different frequencies allows the efficient reconstruction of the model from coarse to fine wavelengths, thus mitigating the inherent nonlinearities of waveform inversion for increasingly higher frequencies. It also eliminates a potential imprint of the source frequency spectrum on the model wave number spectrum. Furthermore, frequency domain inversion has the computational advantage of being able to adapt the grid spacing to the frequency to further decrease the computational cost, and the forward scheme from Pratt [1999] is very fast for multisource problems. [9] This method is field tested in crosswell configurations [Pratt and Shipp, 1999; Song and Williamson, 1995; Song et al., 1995], and its applicability to two-dimensional longoffset surface data was proven by several synthetic tests [Hole et al., 2005; Pratt et al., 1996; Sirgue and Pratt, 2004]. Operto et al. [2004] successfully applied it to a densely sampled surface seismic survey in the Apennine Mountains of Italy. 3.1. Background Model [10] A first-arrival traveltime tomography velocity model of Hole et al. [2006] (Figure 7) was used as background model. It has a RMS traveltime misfit of 40 ms. The largest misfits of 100 –150 ms pertain to the longest offsets. In order to decrease them, an additional traveltime inversion for velocities below 4 km depth was carried out using a weaker regularization. The resulting model was slightly smoothed to remove raypath artifacts and extrapolated to

mitigate edge effects. Low frequent seismograms (3 km). The accuracy of the final model can be locally assessed by comparing it with the sonic log from the SAFOD main hole (Figure 12). The model matches the transition from 3.5 km/s above sea level (sedimentary rock) to 5.5 km/s (granite) very well. The velocity decrease to 4.7 km/s at 1.2 km bsl (sedimentary rock) is recovered, although 0.2 km too shallow. The thin (100 m) low-velocity zones detected by the sonic log at 1.3, 1.6, and 2.0 km bsl are not imaged because of the limited model resolution. Attempts to invert Table 1. Groups of Simultaneously Inverted Frequenciesa Frequencies (Hz)

Grid Spacing (m)

Nodes

Wavelength Filter (km)

Residual Reduction (%)

3.2, 3.6, 4.0 4.4, 4.8, 5.2 5.6, 6.0, 6.4 6.8, 7.2, 7.6 8.0, 8.8, 9.6 10.4, 11.2, 12.0 12.8, 13.6, 14.4

100 50 50 50 25 25 25

401*101 801*201 801*201 801*201 1525*361 1525*361 1525*361

1.250 1.000 0.800 0.700 0.500 0.400 0.250

21.3 28.6 32.3 33.8 24.6 23.7 25.9

a The wavelength filter cuts smaller features and was applied to the velocity gradient in horizontal direction to remove artifacts at length scales that cannot be resolved. Half of that value was used for filtering in vertical direction. To save CPU time, the model grid was adapted during the inversion, increasing sampling as the inverted frequency increases.

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higher frequencies failed. Above 15 Hz, an increasing number of apparently artificial velocity anomalies appeared (similar to the ones near km 27 at sea level), and the maximum velocity exceeded 7 km/s. Since the signal is still good at these frequencies, spatial aliasing and the lack of an

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appropriate attenuation model are considered to be the main limiting factors for the inversion. [18] The final waveform model resolves variations of the top of the Salinian basement in great detail, and it also shows prominent structures within the granite. It confirms

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Figure 10. (a) Phases (top), phase residuals (middle), and weights (bottom) for SP 34 (Figure 8, Figure 9) at 4.0 Hz and (b) for SP 3 at 6.8 Hz. Stars mark the position of the shots. The phase at each receiver (top panel) is shown for the observations (grey solid line), and for the synthetics before (red dashed line) and after the inversion (blue dotted line) of the corresponding frequency group (see Table 1). In the vicinity of the shotpoints, the trend of the phase is relatively simple, and it corresponds essentially to the first arrival. Toward greater offsets, wide-angle reflections intermix with the first arrival, and the phase becomes more complicated. The middle panel shows the residual phases before (red dashed line) and after the inversion (blue dotted line). For better visibility, both curves were smoothed by a running average. In terms of traveltimes, the misfit reduction in (a) near km 33 and 37 corresponds to 40 ms, and in km 18– 25 (b) to 60 ms. The small initial misfit at 4 Hz (and, hence, relatively small misfit reduction) reflects the high accuracy of the traveltime starting model with respect to this frequency. Weighting factors are displayed at the bottom of each panel. They were computed from the derivative of the unwrapped phase. Small weights indicate noisy data. A taper has been applied at 2 –3 km distance.

some higher-resolution details of an alternate velocity model ‘‘b’’ of Hole et al. [2006, Figure 3b] that were not strongly constrained by the traveltimes, such as a depression in the 5 km/s contour from km 18– 20, and two distinct high-velocity bodies at km 16 and 23 –24. A joint gravity and traveltime tomography model by Thurber et al. [2004b] showed the latter anomaly in the same location with similar velocities. It is also contained in the model of Malin et al. [2006] based on downhole recordings of active and passive data, although smaller in amplitude and located 1 km further to the southwest.

4. Reflection Seismic Imaging [19] Kirchhoff prestack depth migration was used for imaging seismic reflectivity, because it is able to image steep dips, and to cope with strong lateral velocity variations [Schneider, 1978]. The migration algorithm used in

this study involves a finite difference solution of the eikonal equation [Hole and Zelt, 1995] to compute the diffraction traveltime contours. Amplitude effects are not accounted for. Since the SAF and related strike-slip faults are the main target of the survey, the data processing focused on steeply dipping reflections. Low dips were processed separately to allow for a more rigorous steep-dip processing. 4.1. Steep-Dip Image [20] Reflections from steeply dipping or vertical reflectors need an incident refracted wavefield traveling horizontally away from the source. Upon being reflected, it propagates back toward the source, and the corresponding traveltime curve has a negative apparent velocity [Hole et al., 1996]. These ‘‘reflected refractions’’ can be enhanced, and even recovered where they are masked by often much stronger first arrivals and S waves. An FK filter was designed to rigorously mute all apparent velocities corresponding to

Figure 9. The phase at 4 (top) and 8 Hz (bottom) is plotted for each seismogram as a function of receiver and shot position. The vertical and horizontal white bars represent a gap in the line. The diagonal white bar marks the omitted nearoffset traces. Other white areas indicate killed traces. The varying width of the rows corresponds to the varying shot spacing. Random patterns in the upper right and lower left corner of the plots (i.e., at long offsets) indicate noise. Short cycles indicate low velocities; broad cycles indicate high velocities. At 4 Hz, the profile is resolved by 40 cycles (corresponding to an average velocity of 4 km/s), although the signal of most shots does not propagate beyond 10 km. The cycles are mostly continuous in the densely sampled receiver domain (x direction), but show large steps in the shot domain (y direction) indicative of spatial aliasing. Such phase plots are an essential tool to assess data quality for frequency domain waveform inversion. An arrow points to the phase of SP 34, which is also displayed in Figure 10. 8 of 15

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Figure 11. (a) Starting velocity model modified after the work of Hole et al. [2006], (b) waveform inversion model after 4 Hz, (c) 7.6 Hz, and (d) 14.4 Hz. A thick black line marks the SAFOD main hole. Contour lines are annotated in kilometers per second. After each group of frequencies, smaller wavelength structures are added to the smooth starting model. Note the sharp gradient from sedimentary rock to Salinian granite between km 20– 25 near sea level in panel (d). Gray arrows denote smile-shaped artifacts due to the gap in the survey near km 30. energy travelling horizontally away from the shot (positive slope in the shot gather). This filter suppresses much of the first arrivals, S waves and most NMO P reflections. The ground roll and 0.1 s of the first arrivals still had to be muted because of their strong amplitudes. Static corrections to a floating datum were applied before filtering and removed afterward, to improve the alignment of the phases, and thereby the filter efficiency. In the face of a few, but visible, steep-dip reflections, the application of an AGC was considered detrimental to the image. Instead, surface consistent amplitude corrections for shot charges and receiver coupling were applied, and spherical divergence and attenuation were roughly approximated by scaling the data with t1.5. [21] The steep-dip images were restricted to 6 km bsl, and a 4-s-long reduced-time window (corresponding to Figure 2) was selected for migration. Including later times did not improve the image near the SAF.

[22] Single shot gathers were migrated separately and subsequently stacked. Using only a selection of shots produced a better image than using all shots, primarily due to the imaging geometry and velocity structure. The best results were obtained from the shots to the northeast (Figure 13). Because of reciprocity, shots closer to the SAF observed on the northeastern end of the line should be suited just as well. However, this configuration would require FK filtering in receiver gathers, which is not possible because of the large shot spacing. Because of symmetry, shots to the southwest of the SAF should be equally qualified for imaging the SAF. However, these shots show a strong secondary direct phase Pch (Figure 2b). It is not entirely eliminated by the FK filters, and it would blur the image. Defocusing effects from the high velocities of the Salinian granite additionally attenuate potential reflected refractions from the SAF. [23] Steep-dip migrations were computed for the two traveltime tomography models of Hole et al. [2006]. To

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Figure 12. Close-up from the final waveform tomographic image at the SAFOD main hole (black, sedimentary rock; white, granite). A vertical dashed line marks the position of the SAF surface trace. At right, sonic log (grey dotted) from the work of Hickman et al. [2005], starting model (blue dashed), and final model (red solid). Velocities from the work of Thurber et al. [2003, 2004a] are shown for comparison (black dashed). The sonic log shows several thin low-velocity zones within and below the granite, which are indicated by pink segments on the drill hole trajectory (left). The size of these zones is below the resolving power of the final model. Differences at 0.7 and 1.1 km bsl between the sonic log and the final model could indicate artifacts in the model, but they could also be the result of the very different lateral resolution of the velocity functions. mitigate the impact of small velocity errors, an additional section was produced by forming the trace envelope in the migrated shot gather before stacking. If the reflector phase is incoherent between different shots, which may be the case because of the large shot spacing, the envelope prevents destructive interference. This ‘‘envelope migration’’ [Simon et al., 1996] is a more robust image at the expense of spatial resolution. Migrations for the final waveform model of Figure 11 are not displayed because they appeared more blurred and showed less structure, indicating a breakdown of the kinematic approximation of Kirchhoff migration in the face of the high resolved waveform model. Operto et al. [2004] made a similar observation. [24] All migrated sections of Figure 13 show a prominent vertical reflector (A) near the SAF. Ray-tracing modeling confirms that it corresponds to the reflection A in SP 8 (Figure 2a). The reflector appears to be divided into an upper segment at 0– 0.5 km bsl at 0.2 km distance northeast of the SAF surface trace and a lower segment down to 2 km bsl at 0.5 km distance. The shallowest part of the reflector A was previously imaged by Hole et al. [2001] and can be confirmed and traced to greater depth by this data. An apparent northeastern dip of the lower part of A is regarded as artifact, because it follows the diffraction traveltime contour, its amplitude fades out rapidly, and it is not confirmed by the envelope migration. [25] A second prominent reflector B can be identified in all sections near the surface trace of the WCF (km 35). Ray tracing confirms that it is the image of the corresponding reflection B in Figure 2a. It is vertical in its upper part and dips progressively toward 80° northeast in its lower part in Figures 13a and 13b. The curvature of this reflector B in Figure 13c, with its upper part dipping slightly to the southwest, emphasizes the connection to the WCF.

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[26] A series of strong reflectors (C) to the southwest and below the SAFOD main hole approaches the SAF, but does not cross it. Apart from the weak reflector D, the wedge between SAFOD and the SAF appears transparent. The reflector image (D) is relatively sharp in Figure 13c, but it is only partly confirmed using the smoother velocity model (Figure 13a). This may indicate that D is an artifact, or that it does not represent a single reflector but a reflector pattern that is too complex to be unraveled. Possibly intersecting or merging reflectors spaced at only a few hundred meters may not be resolved by the low dominant wavelength of 250 m (at 2 km/s). Therefore the transparent appearance of the wedge between the SAFOD main hole and the SAF surface trace could also be a result of a lack of resolution. Further to the southwest of the SAFOD site, several steep-dip reflectors are observed at shallow depth. They are particularly strong around km 18 (E in Figure 13). 4.2. Low-Dip Image [27] Horizontal to moderately dipping reflectivity was imaged separately using a similar processing sequence. The strong Sg wave was again suppressed with an FK filter for apparent velocities. In this case, however, the filter was designed to suppress all traveltime curves with velocities slower than 8 km/s, leaving only reflections with high apparent velocities, such as the typical NMO reflections. More extensive first-breaks mutes (0.5 s) were applied, and the amplitudes were normalized with an AGC of 2 s. For low dips, the recording length provides a section of the whole crust. However, the velocities below 5 km depth are not constrained by this survey. The velocity model used for this migration is an interpolation between model ‘‘a’’ of Hole et al. [2006] (Figure 7) and a uniform velocity of 7 km/s at 30 km depth. In the Salinian middle crust, the model is roughly consistent with a refraction seismic velocity model derived by Trehu and Wheeler [1987] for a parallel, deep seismic profile located 40 km to the southeast of SAFOD. [28] Because of the poorly constrained velocities, imaging problems are to be expected. Indeed, a simple stack of the depth-migrated single shots does not show much detail. However, forming the trace envelopes before stacking reveals several crustal scale features (Figure 14). The most prominent reflectivity (J) at 21– 26 km bsl (8 –10 s TWT in the data) dips 20° northeast toward the SAF. Because of its dip, this reflector is well constrained by several shots despite a decrease of the nominal CMP coverage toward the edges of the section. The apparent broad width of that zone (J), however, may be the result of velocity errors or of merging two or more separate reflectors into one. The horizontal part of reflector J between km 15– 25 cannot be traced back to single shots. Another reflector H in the Salinian at 15 –18 km bsl dips at 25° toward the southwest. Two features F and G are barely visible in the complete envelope stack, but can be unraveled by stacking only two shots in the vicinity of the SAFOD site (Figure 15). They are horizontal reflectors at 8.5 and 12.5 km bsl and extend at least 4 km along the profile. The gap between reflectors F and G and the SAF could be a result of the difficulty of imaging beneath the strong velocity contrasts in the upper crust (the dip of reflector H allows it to be imaged from

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Figure 13. Steep-dip prestack depth migrations. The shots used to produce the images are indicated on top of the panels. Panels (a) and (b) are based on the velocity model shown in Figure 7. For panel (b), trace envelopes have been formed after single-shot migration and before stacking. Panel (c) was produced with the more complex velocity model of Hole et al. [2006, Figure 3b]. All sections were normalized for subsurface coverage. Hollow black arrows indicate the position of two reflectors imaged by Hole et al. [2001] using the shallow 1998 data. Green arrows point to a weaker reflection D. Reflectors A– E is referred to in the text. Slightly elevated amplitudes at 3 – 4 km in the Franciscan are migration artifacts ‘‘leaking’’ from the strong reflectors C below SAFOD. High amplitudes near the surface northeast of km 35 probably indicate steep faults and folds within the sediments of the Great Valley Sequence that are not well resolved due to shot spacing and frequency content of the signal. surface points further southwest). Errors in the velocity model may bias the depth and dip of the deepest reflectors by up to 10%.

5. Joint Interpretation and Discussion 5.1. Middle and Lower Crust [29] An interpretation of the deep low-dip reflection image is presented in Figure 16. The reprocessed seismic COCORP (Consortium for Continental Reflection Profiling) Parkfield data [McBride and Brown, 1986] also observe deep (9 s TWT) reflections in the Salinian terrane at 5 – 15 km distance from the SAF. This aligns approximately with the northeastern part of the reflector J. On the basis of reports of Moho depth by Walter and Mooney [1982] and others of 25 km for the Salinian in the Parkfield area, McBride and Brown [1986] interpret this reflector as crustmantle transition. More recent data and models for Central California [Fuis, 1998; Page et al., 1998] corroborate this Moho depth. It is therefore unlikely that J represents a single continuous Moho-reflector, because neither the steep dip nor the shallow depth at its southwestern end fit the regional models. Trehu and Wheeler [1987] observed a deep reflective band on a SAF-perpendicular profile located 40 km to the southeast. Because of a coincident low-velocity zone, they interpreted these reflectors as a wedge of subducted sediments. In their model, this wedge termi-

nates 30 km southwest of the SAF at 20 km depth. The steep southwestern part of reflector J in 25 km distance from the SAF at 21 km bsl matches the along-strike projection of Trehu’s low-velocity zone quite well. Zone J is interpreted to represent this wedge of subducted sediments merging into the Moho 15 km southwest of the SAF (Figure 16). [30] Reflectors F, G, and H lie within the Salinian terrane. Reflector H is probably related to transpression across the SAF, but there is no additional evidence to support this interpretation. The reflector F was previously imaged by the COCORP data [McBride and Brown, 1986] extending from 2 km to at least 15 km southwest of the SAF. It could indicate the base of the Salinian batholith, which is at 8 km bsl according to recent magnetotelluric surveys [Becken et al., 2005]. Alternatively, the reflector F may represent a fluid-rich horizon within layers of metasedimentary rocks. The brittle-ductile transition is constrained by seismicity, geodesy, and heat flow at 11 – 14 km bsl [Michelini and McEvilly, 1991; Murray et al., 2001; Waldhauser et al., 2004; Williams et al., 2004]. The depth of reflectors F and G near the base of the brittle crust suggests that they could represent (probably fluid lubricated) shear zones, which potentially decouple upper and lower crust. The study of Waldhauser et al. [2004] revealed two moderately northwest-plunging streaks of seismicity at 7 – 8 and 11 – 12 km depth southeast of SAFOD on the locked part of the SAF. The correspondence of these depths with reflectors F and

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Figure 14. Low-dip prestack depth migration of the AGC-scaled reflection data. Trace envelopes were formed after single-shot migration and before stacking. A 10-km wavelength low-pass filter was applied to balance the image. Coherency was enhanced by raising the k space amplitudes to the power of 1.5. In the Salinian terrane, prominent reflectors (F – J) can be seen at all crustal levels. Above 8 km depth, the image is blurred by remnant direct phases (mainly Sg, Pch) and affected by first-break mutes. G suggests a relationship. If these reflectors extend to (and terminate at) the fault, the asperity could be responsible for the observed localization of seismic slip. 5.2. Upper Crust [31] To further the interpretation, a manual line drawing of the steep-dip reflection section was overlaid onto the waveform inversion velocity model (Figure 17). Most of the steep reflections originate from within the granite (C, E). The reflectors E to the southwest are related to a steep rise of the sediment-granite contact. The low-average amplitudes recorded in the area of the adjacent trough from km 18 – 20 (Figure 4) indicate that it is filled with clastic deposits, which attenuate the signal by diffractions. The reflectors C and the associated velocity variations from 5 – 6.5 km/s show a pronounced zoning of the granite, perhaps indicative of multiple intrusion cycles. Magnetic susceptibility variations recorded in the pilot hole [McPhee et al., 2004] also indicate strong heterogeneity within the granite. The northeasternmost reflectors of the series C and the upper part of reflector D correlate roughly with the outline of the granite as inferred from the waveform tomography velocity model. [32] Between 0 and 2.5 km bsl, low-velocity separates the Salinian granite from the SAF (Figure 17). This low velocity zone coincides with low resistivities [Unsworth and Bedrosian, 2004; Unsworth et al., 1997], low densities [McPhee et al., 2004], and drilled sedimentary rock [Barton et al., 2005; Zoback et al., 2005]. There is evidence from surface geology that this sedimentary wedge involves at

Figure 15. Portion of low-dip prestack depth migrated section using only SP 30 and 32. A star indicates the closely spaced shots positions near SAFOD. Processing is the same as for Figure 14. The image of the two horizontal reflectors at 8.5 and 12.5 km bsl (arrows) labeled (F, G) in Figure 14 has improved considerably. The apparent bending at the edges of the reflectors and a weak southwest-dipping structure at km 22– 25 are artifacts caused by migrating two shots only.

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Figure 16. Structural interpretation of the crustal sections displayed in Figure 15. Solid lines mark reflectors or velocity contours; dashed lines are interpretation. Arrows indicate hypothetical fault movements. Dots on the SAF correspond to the depths of subhorizontal streaks of seismicity that extend from the northwest (5-km depth) or southeast (8 and 12 km) [Waldhauser et al., 2004]. According to Williams et al. [2004] the average depth of the 350° isotherm in the SAFOD area varies from 11 to 14 km along strike from the creeping to the locked segment (northwest to southeast). least three major faults [Hole et al., 2006]. Shallow reflection seismic images indicate a multitude of steep faults in its uppermost kilometer [Catchings et al., 2006]. The mapped trace of the steeply southwest dipping active Buzzard Canyon Fault (BCF) [Hole et al., 2006] can be correlated to the reflector D, suggesting a possible connection to the SAF at depth. Such an interpretation is supported by observations of fault zone guided waves partitioning on the BCF [Alvarez et al., 2005]. Probing the sediments at depth, the SAFOD logs [Hickman et al., 2005; Zoback et

al., 2005] indicate several shear zones, at least two of which are active. Microseismicity also reveals at least two strands of activity at depth [Nadeau et al., 2004]. [33] The strong, isolated reflectors A lie under the Gold Hill Fault (GHF) 0.2 –0.5 km northeast of the active trace of the SAF. A few kilometers to the southwest, Li et al. [2004] detected guided waves in a northern fault strand at 300-m distance to the SAF surface trace. Geodynamic models of Simpson et al. [2006] show that the stress field due to the transition from the locked to the creeping segment could

Figure 17. Manual line drawing of the major reflectors from the steep-dip migration of Figure 13 (thick white lines) overlaid onto the final waveform tomography velocity model of Figure 11. The top of the Salinian granite, as inferred from the 5 km/s contour, is marked by a thin white line. Local earthquake hypocenters (white circles) from the work of Roecker et al. [2005] and Thurber et al. [2004a] constrain the position of the SAF at depth. Shear zones (pink bars) penetrated by the SAFOD main hole (gray) were inferred from low-velocity zones (Figure 12). Solid black lines are interpretation. The following surface position of faults is indicated by dotted lines: BCF, Buzzard Canyon; SAF, San Andreas; GHF, Gold Hill; WCF, Waltham Canyon; for location, see Figure 1. 13 of 15

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Compression normal to the SAF is responsible for the uplift of the Franciscan block and for the backthrust.

6. Conclusions

Figure 18. Three-dimensional sketch of the WCF reflector. The fault-parallel axis refers to the SAF. The solid thick line indicates the reflector, and the dashed line indicates the surface trace of the WCF. Approximately 10 km to the southeast, the terrane boundary tears and jumps to the southwest to above its deep location. deflect the main SAF toward the northeast. In this context, the reflectors A could represent a paleo-SAF, which is now no longer active. [34] In summary, the evidence suggests that the deep sedimentary wedge between the edge of Salinia to the southwest and the GHF to the northeast composes a flower structure of three or more major and several minor faults, which together form the San Andreas Fault Zone (SAFZ). From guided wave observations, Li et al. [2004] assessed the width of fault strands within the SAFZ at 50– 150 m. Constructive and destructive interference of reflections from such faults with varying spacing and thickness is strongly frequency dependent. The spatial resolution of the images and of the velocity models is, however, not sufficient to unravel more details of the internal structure of the SAFZ. The lack of hypocenters above 2 km depth, and the low velocities indicate low fault strength in this zone. [35] The steep reflector B is interpreted to represent the boundary between the Franciscan terrane and the Great Valley Sequence at depth. The deviation of 1 km between its location and the WCF surface trace can be reconciled taking the regional trend of the terrane boundary into account (Figure 1). The WCF is generally parallel to the SAF, but locally splays to the east. On the basis of the assumption that this local deviation affects only the surface, a three-dimensional reconstruction can be attempted (Figure 18). Page et al. [1998] argued that the strong topography and the low resistance of the Franciscan me´lange to tectonics and erosion indicate extraordinary uplift rates. On that basis, the reflector is interpreted as a northeast-dipping fault which was locally tilted in the nearsurface by backthrusting in the upper sedimentary layers.

[36] This manuscript presents an early example of waveform tomography applied to surface refraction data. Several issues were encountered with complex surface topography and geology, spatial aliasing in the shot domain, and strong variation in attenuation. The heterogeneous attenuation prevented the inversion of true amplitudes and led to the decision to model phase only. In order to minimize the impact of noise, weighting factors were computed at each frequency from local phase coherency. Large shot spacing and noise issues limited the inversion to frequencies below 14 Hz, even though good signal exists at higher frequency. The waveform velocity model stops short of resolving single faults. More details of the inversion strategy will be presented in another paper. Further development of these methods is merited by the observed improvement in velocity resolution. Noise issues and shot aliasing also affected the reflection seismic images. Forming the trace envelopes in the migrated shot gathers before stacking solved this problem, albeit at the expense of resolution. Despite strong topography, near-surface geology variation, and aliased shots, direct images of vertical reflectors were obtained from the long-offset data. [37] The waveform inversion and steep-dip prestack migration methods produced images of the SAF at SAFOD with unprecedented spatial resolution. The top of Salinian granitic rocks was sharply resolved, revealing a complex erosional and faulted topography. The boundary between the Franciscan and the GVS is imaged as a curved reflector suggesting a thrust wedge geometry. A very strong vertical reflector 0.2–0.5 km northeast of the surface trace of the SAF was imaged between 0 and 2 km bsl. Together with reflectors and velocity boundaries associated with the complex edge of Salinia, it unveils a 2-km-wide and 3-km-deep wedge of lowvelocity sedimentary rock containing several faults between Salinia and the SAF. These seismic images will aid the interpretation of features in the SAFOD hole. [38] Acknowledgments. This work was funded by the U.S. National Science Foundation grant EAR-0106534. The data collection was supported by NSF, the Deutsche Forschungsgemeinschaft, the GeoForschungsZentrum Potsdam, and the U.S. Geological Survey. Steve Hickman aided in coordination of the field logistics. Thanks to Gerhard Pratt for making his waveform inversion code available and to David Okaya, who provided the migration programs. Comments by George Spence, Stuart Henrys, and an anonymous reviewer helped to improve this manuscript.

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F. Bleibinhaus and J. A. Hole, Department of Geosciences, Virginia Tech, 4044 Derring Hall, Blacksburg, VA 24061, USA. ([email protected]; [email protected]) G. S. Fuis, U.S. Geological Survey, 345 Middlefield Rd., MS977, Menlo Park, CA 94025, USA. ([email protected]) T. Ryberg, GeoForschungsZentrum, Telegrafenberg E322, Potsdam 14471, Germany. ([email protected])

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