Ultracataclasite structure and friction processes of the Punchbowl fault, San Andreas system, California

ELSEVIER

Tectonophysics 295 (1998) 199–221

Ultracataclasite structure and friction processes of the Punchbowl fault, San Andreas system, California Frederick M. Chester Ł , Judith S. Chester Department of Geology and Geophysics, Texas A&M University, College Station, TX 77843, USA Received 10 October 1996; accepted 28 May 1998

Abstract The Punchbowl fault is an exhumed, 40C km displacement fault of the San Andreas system. In the Devil’s Punchbowl, the fault contains a continuous ultracataclasite layer along which the Punchbowl Formation sandstone and an igneous and metamorphic basement complex are juxtaposed. The fabric of the ultracataclasite layer and surrounding rock indicate that nearly all of the fault displacement occurred in the layer. By analogy with nearby active faults, we assume that the Punchbowl fault was seismogenic and that the ultracataclasite structure records the passage of numerous earthquake ruptures. We have mapped the ultracataclasite layer at 1 : 1 and 1 : 10 to determine the mode of failure and to constrain the processes of seismic slip. On the basis of color, cohesion, fracture and vein fabric, and porphyroclast lithology, two main types of ultracataclasite are distinguished in the layer: an olive-black ultracataclasite in contact with the basement, and a dark yellowish brown ultracataclasite in contact with the sandstone. The two are juxtaposed along a continuous contact that is often coincident with a single, continuous, nearly planar, prominent fracture surface (pfs) that extends the length of the ultracataclasite layer in all exposures. No significant mixing of the brown and black ultracataclasites occurred by offset on anastomosing shear surfaces that cut the contact or by mobilization and injection of one ultracataclasite into the other. The ultracataclasites are cohesive throughout except for thin accumulations of less cohesive, reworked ultracataclasite along the pfs. Structural relations suggest that: (1) the black and brown ultracataclasite were derived from the basement and sandstone, respectively; (2) the black and brown ultracataclasites were juxtaposed along the pfs; (3) the subsequent, final several kilometers of slip on the Punchbowl fault occurred along the pfs; and (4) earthquake ruptures followed the pfs without significant branching or jumping to other locations in the ultracataclasite. By comparison with rock friction experiments, the slip localization along the pfs in the ultracataclasite implies rate weakening behavior with a critical slip distance similar to laboratory values, and thus relatively small nucleation and breakdown dimensions for earthquake ruptures. Of the various mechanisms proposed to explain the low strength of the San Andreas and to produce dynamic weakening of faults, those that require or assume extreme localization of slip are most compatible with our observations.  1998 Elsevier Science B.V. All rights reserved. Keywords: structure; faulting; friction; cataclasis; earthquakes

Ł Corresponding

author. E-mail: [email protected]

0040-1951/98/$19.00  1998 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 9 8 ) 0 0 1 2 1 - 8

200

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

1. Introduction Large deformations of the upper portion of the Earth’s crust are primarily achieved through seismic faulting. Plate boundary faults, such as the San Andreas fault, achieve large displacements via numerous seismic slip events. The San Andreas fault and possibly other large displacement faults slip under a resolved shear stress that is low compared to simple quasi-static models of faulting based on experimental rock mechanics (e.g., Hickman, 1991). In spite of the extensive study of seismic faulting on the San Andreas system, the physical and chemical processes responsible for the nucleation, propagation and arrest of earthquake ruptures are poorly understood (e.g., Sibson, 1989; Scholz, 1990; Brune, 1991; Segall, 1991). In fact, the geophysical literature contains many underconstrained models and untested hypotheses for the mechanics of seismic faulting (e.g., Melosh, 1979, 1996; Byerlee, 1990, 1993; Brune, 1991; Rice, 1992; Sleep and Blanpied, 1992; Brune et al., 1993). Experimental and field studies of the formation and growth of brittle faults lead to the conclusion that a fault does not form and grow as an isolated shear crack (e.g., Gay and Ortlepp, 1979; Cox and Scholz, 1988; Cowie and Scholz, 1992; Reches and Lockner, 1994). Rather, a fault grows through a complex breakdown process at the fault tip that involves coalescence of fractures and shears to form throughgoing principal slip surfaces (Sibson, 1986). Concentration of shear stress at a fault tip produces arrays of isolated cracks that extend, interact, and link to form fracture networks and throughgoing shears. This growth process is associated with a breakdown in strength from the intrinsic yield strength of the intact rock to the residual friction on the throughgoing principal slip surfaces (Cowie and Scholz, 1992). Continued slip on the newly formed principal slip surfaces is associated with various wear processes that produce layers of breccia, cataclasite and gouge bounded by tabular zones of damaged rock (Sibson, 1986). Subsequent slip could modify the internal structure of a fault zone in a number of ways (e.g., Hull, 1988; Wojtal and Mitra, 1988; Means, 1995). Growth of an individual earthquake rupture involves a breakdown process at the rupture tip similar to the process of fault growth (e.g., Rudnicki,

1980; Swanson, 1992; Scholz et al., 1993). The dimensions of the breakdown zone for an earthquake rupture on an existing fault are probably less than those for new faults because an existing fault already has a damaged zone (Cowie and Scholz, 1992). The structural signature from the passage of a seismic rupture probably is a narrow zone or zones of concentrated shear demarcating the rupture surface within a broader zone of distributed fracturing. The occurrence of repeated earthquakes on an existing principal slip surface implies that the surface is restrengthened during interseismic periods (e.g., Rice, 1983). The degree of restrengthening may influence the path of subsequent earthquake ruptures. If an existing principal slip surface remains weak relative to the surrounding rock, then a subsequent earthquake rupture may occur along that surface rather than branch or jump to a new location. In this case, a principal slip surface could accommodate numerous slip events and ultimately accumulate very large displacement. In contrast, if a principal slip surface is greatly restrengthened by some process, such as neomineralization, subsequent ruptures may occur elsewhere within the fault zone. In general, the degree to which an earthquake rupture follows a preexisting principal slip surface and the thickness of an earthquake damage zone relative to the thickness of an entire fault zone are unknown. Exhumed faults display a static view of the cumulative structure produced during fault growth and subsequent episodes of fault activity at various environmental conditions. Although deciphering the kinematics and separating out the alteration resulting from near-surface weathering can be difficult, these faults provide an important view of the physical and chemical processes operating in the seismogenic regime (Sibson, 1977; Bruhn et al., 1990; Chester et al., 1993). Exhumed faults show considerable variation in internal structure, but most large displacement faults are tabular zones of concentrated shear bordered by damage zones of fractured and brecciated rock (Flinn, 1977; Brock and Engelder, 1977; Wallace and Morris, 1986; Chester and Logan, 1986; Sibson, 1986; Bruhn et al., 1990; Little, 1995). Concentrated shear often is recorded by the reorientation and destruction of primary structures, development of cataclastic foliations, and presence of extremely comminuted material such as gouge or

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

ultracataclasite. We refer to a zone of concentrated shear (cataclasite and ultracataclasite) as the fault core (Chester et al., 1993). At the macroscopic scale, a fault core represents the principal slip surface of the fault (Sibson, 1986). A fault zone may display a core near one or both boundaries of the damaged zone, a single core centralized in the damaged zone, or an anastomosing network of several cores within the damaged zone (Wallace and Morris, 1986; Rutter et al., 1986). Often, the fault core contains mesoscopic scale principal slip surfaces, i.e., features recording further concentration of shear at the mesoscopic scale (Chester et al., 1993; Arboleya and Engelder, 1995). Thus, the definition of a principal slip surface depends on scale of observation (Arboleya and Engelder, 1995). The Punchbowl fault, a deeply exhumed, large displacement fault of the San Andreas system, displays either a single principal slip surface or paired principal slip surfaces within broader zones of damaged host rock (Chester and Logan, 1986; Chester et al., 1993; Schulz and Evans, 1998). The intensity of deformation progressively increases toward the fault cores. The cores may contain foliated cataclasites, but always contain a single, continuous layer of ultracataclasite. The ultracataclasite consists almost entirely of matrix particles less than 10 µm in diameter with porphyroclasts of vein fragments (Chester and Logan, 1987). Fabric analyses indicate that nearly all of the shear displacement of the fault was accommodated within the fault cores, and that most of the displacement was localized to the ultracataclasite layer. Principal slip surfaces are always present within the fault cores, suggesting localization at the mesoscopic and microscopic scales as well. The purpose of this paper is to characterize the mesoscopic structure of the primary ultracataclasite layer of the Punchbowl fault in the Devil’s Punchbowl Los Angeles County Park (Fig. 1). At this location the Punchbowl fault juxtaposes fractured basement rock and arkosic sandstone of the Punchbowl Formation along a continuous and distinct, 0.3 m thick ultracataclasite layer (Figs. 2 and 3). Present-day exposures of the Punchbowl fault record shear at 2 to 4 km depth (Chester and Logan, 1986). The depth has been estimated using post-Pliocene uplift and erosion rates for the San Gabriel Mountains (Oakeshott, 1971; Ehlig, 1975; Morton and

201

Matti, 1987), thickness of the sedimentary sequence in the Devil’s Punchbowl basin cut by the fault, and mineral assemblages and microstructures of the fault rocks (Chester et al., 1993). We can not say definitively that the Punchbowl fault slipped seismically; however, active faults in the Central Transverse Ranges are seismogenic. These active faults cut rock types similar to or the same as those cut by the Punchbowl fault and are situated in the same tectonic setting. By comparison, we assume that the Punchbowl fault represents a mature, formerly seismogenic, fault zone of the San Andreas system. Thus, we use the mesoscopic structure of the ultracataclasite to infer the distribution of displacement during the final stages of faulting, and to constrain the characteristics of earthquake rupture surfaces and rupture processes.

2. Geology of the Punchbowl fault The Punchbowl fault is an inactive, exhumed fault of the San Andreas transform system in the central Transverse Ranges of southern California (Fig. 1). It is located approximately 5 km southwest of and is parallel to the active strand of the San Andreas fault. The Punchbowl fault is truncated to the northwest and to the southeast by the San Andreas, and can be considered an abandoned strand of the San Andreas System (Dibblee, 1967, 1968; Barrows et al., 1987; Matti and Morton, 1993). The fault has a sinuous trace, is often steeply dipping to the southwest, and in some locations displays anastomosing fault strands that bound slices of exotic rock types. The fault cuts crystalline rock of the San Gabriel basement complex, and along much of its length juxtaposes basement rocks and the Punchbowl Formation of the Devil’s Punchbowl basin (Noble, 1954). The Devil’s Punchbowl basin is a small, elongate basin located between the Punchbowl and San Andreas faults which is filled with more than a 1-kmthick sequence of the Punchbowl Formation (Noble, 1954; Woodburne, 1975). The basin probably originated as a pull apart basin during the early phases of Punchbowl faulting in the Middle Miocene. The Punchbowl Formation consists of a cobbly to pebbly arkosic sandstone with interbeds of siltstone and a basal breccia. The Punchbowl Formation overlaps

202

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

Fig. 1. Geologic map of the Punchbowl fault in the vicinity of the Devil’s Punchbowl Los Angeles County Park, northeast San Gabriel Mountains, California. The study area is in the Devil’s Punchbowl Park, at the southeastern end of the Punchbowl basin.

the angular unconformity with the Paleocene San Francisquito Formation along the northeastern edge of the basin showing that the sediments onlapped relief to the east (Weldon et al., 1993). The southwestern side of the basin is truncated by the Punchbowl fault and the marginal sedimentary deposits are not preserved. A less than 0.5 km wide outcrop of

the basal breccia extends approximately 11 km to the southeast of the basin along the Punchbowl fault (Noble, 1954). This basal breccia may record sedimentation associated with the fault-controlled margin of the Devil’s Punchbowl basin (Weldon et al., 1993). The basal breccia both overlies and is cut by subsidiary faults of the Punchbowl system suggest-

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

203

Fig. 2. Photographs of the Punchbowl fault zone and ultracataclasite layer. (a) View of the Punchbowl fault zone looking northwest from the Devil’s Chair overlook. White aplite of the San Gabriel basement complex (left) and Punchbowl Formation sandstones (right) are juxtaposed along the ultracataclasite layer. Note that the lithologic contacts in the basement have been rotated near the contact with the ultracataclasite layer as a result of distributed shear within the fault zone. Location of the slip-parallel exposure mapped at a scale of 1 : 10 is indicated. (b) A portion of the slip-parallel exposure of the ultracataclasite layer. Basement at the top and sandstone at the bottom. The contacts between the layer and cataclastic host rocks are sharp, and the ultracataclasite is texturally distinct. Note the wedge-like protrusion of ultracataclasite into the Punchbowl Formation sandstone.

204

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

ing contemporaneous deposition and fault movement (Chester, unpubl. mapping, 1995). In the Devil’s Punchbowl Los Angeles County Park, which covers the eastern portion of the Devil’s Punchbowl basin, the Punchbowl fault system is composed of two fault strands that bound a slice of fractured and faulted basement up to 0.5 km in thickness (Fig. 1; Noble, 1954). The slice of fractured basement between the northern and southern fault strands contains a heterogeneous assemblage of Precambrian biotite gneiss and quartzofeldspathic gneiss with alternating leucocratic and melanocratic bands, and massive to foliated Cretaceous plutonic rocks including quartz diorite, tonalite, granodiorite, and biotite monzogranite (Cox et al., 1983). Locally within the Devil’s Punchbowl Park, the monzogranite is a brilliant white aplite that complexly intrudes melanocratic rocks. The southern fault strand juxtaposes similar rock types (Cox et al., 1983), is not always well developed, and is segmented and discontinuous (Chester, unpubl. mapping, 1995). The northern fault strand is much better developed and is defined by a continuous layer of ultracataclasite along which the Punchbowl Formation and fractured basement are juxtaposed (Fig. 2). A total right-lateral separation on the Punchbowl fault system of 40 to 50 km is indicated by offset of the correlative San Francisquito and Fenner faults (Dibblee, 1967, 1968). This separation is consistent with the offset of the Devil’s Punchbowl basin from the inferred sediment source terrane (Weldon et al., 1993). The partitioning of total displacement between the northern and southern fault strands is unknown. The distribution of rock types along the Punchbowl fault in the Devil’s Punchbowl suggests that at least 10 km of separation occurred on the northern strand. Less than 10 km of separation would require an offset extension of the Fenner fault and associated Pelona Schist to occur in the slice between the northern and southern strands (Fig. 1). On the basis of internal structure and continuous character, we think that the northern strand accommodated a much greater percentage of the total slip than the southern strand. The timing of movement on the Punchbowl fault system is less clear. Although the fault appears to have been active during the formation of the Devil’s Punchbowl basin, at least half of the total displacement occurred during the Pliocene

and Pleistocene, after the deposition of the Punchbowl Formation was complete (Barrows et al., 1985; Matti et al., 1985; Weldon et al., 1993). Analysis of the subsidiary fault fabric and folds in the Punchbowl Formation suggests that the slip vector for the northern strand plunged approximately 30º to the southeast during later stages of faulting, which is consistent with the reverse fault geometry (Chester and Logan, 1987).

3. Method Two exposures of the ultracataclasite layer of the northern fault strand were mapped in detail. One exposure is located directly below the Devil’s Chair overlook in the Devil’s Punchbowl Los Angeles County Park, and the other is across the small canyon approximately 100 m to the northwest of the Chair. The orientation of the outcrop below the Chair is approximately perpendicular to the average slip vector (slip-perpendicular) inferred for the Punchbowl fault (Chester and Logan, 1987). The other outcrop surface is approximately perpendicular to the layer and parallel to the slip vector (slip-parallel). At each location, the ultracataclasite layer is located in a gully and was partly exposed by erosion. We removed additional dirt and talus with a shovel and broom to completely expose approximately 30 m of the layer. Two sections of the layer, totalling 16 m length, were mapped at a scale of 1 : 10 using a decimeter grid as a guide (Fig. 4). The grid was placed on the outcrop and repeatedly moved as mapping progressed. The grid was positioned according to a reference line that was attached to the outcrop. The projection for each strip map is normal to the fault plane. An additional map was made at a scale of 1:1 of a section of the slip-parallel outcrop (Fig. 5). The section mapped was excavated to a flat surface using a sharp chisel; debris was removed with an air blower, and the surface was rinsed with water. The area excavated was approximately twice the size of the map (Fig. 3). The 1 : 1 map was made by tracing the visible structure on plate glass that was firmly affixed to the outcrop surface. The maps show the locations of lithologic contacts, veins, small faults, and fractures. Fine- and medium-grained sandstone units are distinguished in

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

205

Fig. 3. Excavation of the ultracataclasite layer mapped at a scale of 1 : 1. Basement at the top and sandstone at the bottom. Note the porphyroclasts of basement rock in the ultracataclasite layer and the subsidiary fault in the sandstone. The prominent fracture surface is located in the center of the layer and identified by arrows. The less cohesive brown ultracataclasite occurs in a layer along the basement side of the pfs. The location of the 1 : 1 map (Fig. 5) is shown.

the Punchbowl Formation. Melanocratic and leucocratic units are distinguished in the basement. Two main ultracataclasite units are distinguished on the basis of color. Color was classified for freshly broken surfaces of the ultracataclasite by comparison with a rock color chart. The two ultracataclasite units are subdivided further on the basis of fracture, vein, and cohesion characteristics. The mineralogy of the host rocks and ultracataclasite were determined through petrographic and X-ray diffraction analyses of samples collected in the Devil’s Punchbowl area (Table 1).

4. Structure of the Punchbowl fault zone 4.1. Punchbowl Formation The damaged zone in the Punchbowl Formation is typically about 15 m thick. The zone is distin-

guished by the increase in mesoscopic fracture and subsidiary fault density above regional levels. Sedimentary layering and other sedimentary structures are evident within the zone, often to within 1 m of the ultracataclasite layer. All along the contact with the ultracataclasite layer, the sandstone is penetratively fractured and faulted, and displays cataclastic textures. The cataclastic sandstone always is texturally distinct from the ultracataclasite and forms a sharp contact with the ultracataclasite layer. In the region of the slip-parallel strip map, a discontinuous layer of medium-grained cataclastic sandstone exists between the fine-grained sandstone and the ultracataclasite layer (Fig. 4). Shear along the contact between the two sandstone units is indicated by a thin accumulation of reddish-brown ultracataclasite (Fig. 5). Subsidiary faults cut both sandstone units, however, the contact between the fine- and medium-grained sandstones is not offset. Many of the faults in the medium-grained sandstone

206

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

Table 1 Average mineralogy of the host rock and ultracataclasite of the Punchbowl fault zone, Devil’s Punchbowl County Park, California Minerals

Quartz Plagioclase feldspar Potassium feldspar Calcite, dolomite, siderite Hornblende, other amphiboles Illite, mica Kaolinite Chlorite Chlorite–smectite mixed layer Smectite Illite–smectite mixed layer Laumontite Clinoptillolite Analcime

Rock unit Punchbowl Formation a

Ultracataclasite b

Crystalline basement c

abundant abundant common trace trace minor trace trace trace minor trace common trace trace

abundant common trace trace trace trace none trace trace abundant trace trace minor minor

abundant abundant minor minor trace minor trace minor minor trace trace minor none trace

Relative mineral abundances determined by X-ray diffraction analysis: abundant >25%; common >10%; minor >3%; trace >0%; none D 0%. a 12 analyses of 9 samples. b 13 analyses of 11 samples (4 samples of olive black and 7 samples of dark yellowish brown ultracataclasite). c 18 analyses of 18 samples.

have a conjugate geometry consistent with extension parallel to the slip direction of the Punchbowl fault. The subsidiary fault and fracture geometries in the mapped exposures are similar to extensional and contractional duplexes described in detail by Swanson (1988, 1990). The contact between the ultracataclasite layer and Punchbowl Formation sandstone displays both relatively straight and somewhat undulatory segments, as well as small wedges of ultracataclasite protruding into the sandstone (Fig. 2b). The undulations between straight sections often appear as embayments into the sandstone. The two largest wedges observed are displayed in the slip parallel exposure. At this location, each wedge is approximately 0.3 m in length and is oriented at a low angle to the ultracataclasite layer. These wedges are not associated with intersections between larger subsidiary faults and the ultracataclasite layer (Fig. 2b). 4.2. Basement complex The slice of basement between the northern and southern fault strands is pervasively fractured and faulted. A progressive increase in deformation inten-

sity towards the ultracataclasite layer is not always evident. However, primary structures in the basement, such as gneissic layering and lithologic contacts, often are progressively reoriented with proximity to the ultracataclasite layer. At many locations, distributed shear of the basement along the ultracataclasite layer produced a zone of foliated cataclasites up to 5 m thick. The lithology of the porphyroclasts in the foliated cataclasites matches the lithology of the adjacent basement host rock. Where foliated cataclasites are derived from melanocratic protoliths, the contact between the foliated cataclasite and the ultracataclasite, and between the foliated cataclasite and the basement host rock, can be gradational. Where non-layered leucocratic rocks are in contact with the ultracataclasite, the zone of cataclasites is thinner and foliations are absent. At these localities the contact between the ultracataclasite and basement is distinct. In the region of the slip-parallel strip map, a 3-m-thick zone of foliated cataclasites in the basement borders the ultracataclasite layer. Foliations are subparallel to the ultracataclasite layer. An intrusive contact between the white aplite and older Precambrian gneiss, as well as other lithologic contacts, can

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221 Fig. 4. Structure of the ultracataclasite layer mapped at a scale of 1 : 10. (a) Slip-perpendicular exposure is shown in two panels: A–B and B–C. (b) Slip-parallel exposure is shown in four panels: D–E, E–F, F–G, and G–H. Location of the photographs shown in Fig. 2b and Fig. 3 are shown. 207

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

Fig. 4 (continued).

208

209

Fig. 4 (continued).

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

210

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

Fig. 5. Structure of the ultracataclasite layer mapped at a scale of 1 : 1.

be traced from the surrounding host rock into the zone of foliated cataclasites. The contact geometries are consistent with drag folding and right lateral shear. The contact between the ultracataclasite layer and basement rock is irregular. In addition to larger wavelength, low amplitude undulations, the contact displays small wavelength, high amplitude protrusions of ultracataclasite into the basement (Fig. 4). In

many cases the protrusions are associated with intersections between melanocratic layers or subsidiary faults and the ultracataclasite layer. 4.3. Ultracataclasite layer The roughness of a surface may be characterized by a dimensionless ratio of amplitude to wavelength within a geometric frequency interval (e.g., Power

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

and Tullis, 1991). In the vicinity of the Devil’s Chair, the ultracataclasite layer is exposed for over 100 m. For wavelengths between 5 and 50 m, the layer constitutes a continuous surface with a roughness that is on the order of 10 3 . Within the excavated and mapped region, the ultracataclasite layer varies between 0.15 and 0.55 m in thickness, and the average roughness of the layer boundaries is on the order of 10 1 to 10 2 . The ultracataclasite layer forms a continuous boundary between the two host rocks; cataclasite derived from the Punchbowl Formation was never found on the southern (basement) side of the ultracataclasite layer. Similarly, cataclasite derived from the basement is not present on the northern (Punchbowl Formation) side of the ultracataclasite layer. Except for the rare pebble- to cobble-size prophyroclasts, approximately 75% of the ultracataclasite is composed of matrix grains that are <10 µm in diameter, 20% is veins and vein fragments, and 5% is single crystal porphyroclasts (Chester and Logan, 1987). The veins are about 100 µm thick, and vein fragments are often about 100 µm in diameter. Rarely are the veins and porphyroclast greater than 1 mm in size. The ultracataclasite is relatively uniform in composition and mineralogically distinct from the surrounding Punchbowl Formation and basement rock. In general, the ultracataclasite is depleted in potassium feldspar, laumontite, chlorite, and illite=mica, and enriched in smectite, clinoptilolite, and analcime (Table 1). The contrasting mineralogy of the ultracataclasite reflects hydration reactions promoted by extreme comminution (e.g., Chester and Logan, 1986; Evans and Chester, 1995).

5. Internal structure of the ultracataclasite layer

211

after referred to as the brown ultracataclasite) always is found in contact with the Punchbowl Formation sandstone on the north side (Fig. 4). The contact between the black and brown ultracataclasites is sharp, continuous, and nearly planar; it is far less irregular than the contacts between the ultracataclasites and host rocks. Relationships noted at several exposures of the ultracataclasite layer between the two mapped segments, and at several additional exposures of the layer elsewhere in the Devil’s Punchbowl Park, suggest the contact between the brown and black ultracataclasites is at least 100 m long, and likely more than 1.5 km long. Inclusions of either ultracataclasite were never found in the other ultracataclasite. The relatively few pebble- to cobble-size porphyroclasts that are present in the ultracataclasite tend to occur in groups near the boundaries of the layer. Porphyroclasts of the Punchbowl Formation are confined to the brown ultracataclasite and porphyroclasts of basement rock are confined to the black ultracataclasite (Fig. 4). Within the region mapped, the brown ultracataclasite constitutes less than half of the layer, and varies between 0.005 and 0.2 m in thickness. Subsidiary faults in the damaged zones of the Punchbowl Formation and basement complex also contain ultracataclasite material. In contrast to the Punchbowl fault ultracataclasite layer, the subsidiary fault ultracataclasite layers are highly variable in color. The subsidiary fault ultracataclasite layers in the Punchbowl Formation display hues of red and yellow, whereas those in the basement display hues of green and red. The differences in color may reflect differences in composition and mineralogy. The mineral phases, major and trace element chemistry, and the stable isotope geochemistry of the fault rocks and adjacent host rocks are being studied and will be reported elsewhere.

5.1. Color 5.2. Fractures The ultracataclasite layer is composed of an olive black (5Y 2=1) unit and a dark yellowish brown (10YR 4=2) unit in both mapped segments. Each unit is remarkably uniform in color. The olive black ultracataclasite (hereafter referred to as the black ultracataclasite) always is found in contact with the crystalline basement on the south side of the layer, and the dark yellowish brown ultracataclasite (here-

Fracture surfaces within the cataclasites of the fault core and within the damage zones on either side of the core often display slip lineations. In contrast, fractures cutting the ultracataclasite layer are relatively smooth and undecorated, and do not display plumose structures, slip lineations or any other feature that would indicate a possible direction

212

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

of shear. Fracture surfaces generally do not offset internal contacts, such as between the brown and black ultracataclasite, or boundaries of the ultracataclasite layer. Except for one prominent fracture surface discussed below, the fractures are less than a meter in length. Of these, the longer fractures are slightly wavy, display a preferred orientation parallel to the layer, and are consistent with the ‘Y-, R1- and P-shear’ orientations of the Riedel array (e.g., Logan et al., 1979). There also is a distinct set of shorter fracture surfaces that display a quasi-conjugate geometry with bisector normal to the ultracataclasite layer, i.e., approximately in the ‘R2- and X-shear’ orientations. The prominent fracture surface (pfs) is relatively planar, perfectly continuous, and extends the entire length of both the slip-parallel and slip-perpendicular exposures (Fig. 4). In many locations, the pfs is coincident with the contact between the brown and black ultracataclasites. All other structures in the ultracataclasite either merge with or are truncated by the pfs. The pfs is far more planar than the boundaries of the ultracataclasite layer and displays a roughness on the order of 10 3 . The pfs is somewhat more undulatory in the slip-perpendicular than in the slip-parallel exposure. 5.3. Veins and cohesion The cohesion of the ultracataclasites varies significantly and systematically throughout the layer. All of the black ultracataclasite and approximately half of the brown ultracataclasite are cohesive. Microscale veins and vein fragments that are barely visible in outcrop are common in the cohesive ultracataclasites. Intact veins generally are oriented at low angles to the fault. However, most veins are distended or broken and strung out into layers subparallel to the fault. Most of the cohesive ultracataclasites display a lower density of fracture surfaces overall but a greater density of fractures oriented at large angles to the layer. In these regions the ultracataclasite weathers to blocky fragments. Other portions of the ultracataclasite, particularly the black ultracataclasite along the boundary of the layer, display a high density of anastomosing fracture surfaces and tends to weather into flakes. Locally, the flaky, cohesive black ultracataclasite displays small crenulations

with amplitudes and wavelengths on the order of 5 mm (Fig. 4). The layering in the cohesive ultracataclasites, defined by veins, fragments of veins, and fracture surfaces, is subparallel to the interior and boundary contacts of the ultracataclasite layer. The concordant relationship is most obvious along undulations of the contact between the brown and black ultracataclasites and along undulations of the boundaries of the ultracataclasite layer. However, truncation of the layering is evident locally, particularly along the pfs and at the boundaries of the wedge shaped protrusions of ultracataclasite into the Punchbowl sandstone cataclasites (Fig. 4). The remaining portion of the brown ultracataclasite is significantly less cohesive, does not contain veins or vein fragments, and tends to part along anastomosing surfaces to produce very small flakes. The less cohesive brown ultracataclasite only occurs as a thin, discontinuous layer along the pfs. In excavating the outcrop for the 1 : 1 mapping, we found that the less cohesive ultracataclasite could be cut easily with a sharp hand-held chisel, whereas the cohesive ultracataclasite had to be fractured with a chisel and hammer. In addition, when rinsing the surface with water, the less cohesive ultracataclasite became pasty when wet. The less cohesive brown ultracataclasite is a very distinct layer that occurs between the black ultracataclasite and the pfs (Figs. 3 and 5).

6. Distribution of slip 6.1. Localization of displacement Field observations and experimental studies support the general assumption that particle size and degree of sorting tend to decrease with increasing displacement along a fault if all other parameters are constant (e.g., Engelder, 1974; House and Gray, 1982; Sibson, 1986; An and Sammis, 1994). These data suggest that the ultracataclasite in the core of the Punchbowl fault records a very large magnitude of shear displacement, and that the abrupt change in texture at the contact between the ultracataclasite and bounding cataclasites records an abrupt change in the magnitude of shear strain. Extreme localization of displacement to the ultracataclasite

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

layer also is suggested by the geometry of lithologic contacts and other primary structures in the host rock surrounding the ultracataclasite layer. The primary lithologic contacts in the cataclasites and fractured rock bounding the ultracataclasite layer are offset and reoriented by distributed shear. However, in spite of the cataclastic deformation, primary lithologic contacts between basement lithologies in the Devil’s Punchbowl, such as between the white aplite and Precambrian gneiss, are easily traced into the fault zone and up to the ultracataclasite layer. The geometry of these contacts indicates that the sum of all fault parallel displacement in the cataclasites and fractured rock bounding the ultracataclasite layer totals less than approximately 100 m. Thus, nearly all of the ten-plus kilometers of displacement on the northern strand of the Punchbowl fault was accommodated by shear within the ultracataclasite layer. The fracture surfaces represent further localization of deformation within the ultracataclasite layer. Although we can not unequivocally demonstrate shear displacement on these fracture surfaces because of the lack of offset markers and surface decoration, several features indicate that at least some of the surfaces are sites of shear. The mesoscale fractures are preferentially oriented subparallel to the ultracataclasite layer, and the fabric is similar to the Riedel shear fracture fabric documented in many experimental and natural brittle deformation zones (e.g., Logan et al., 1979; Swanson, 1988; Arboleya and Engelder, 1995). The mesoscale fabric also is similar to the microscale shear band fabric of the Punchbowl ultracataclasite documented by Chester and Logan (1987). Unlike the mesofractures, however, shear offsets at the microscale are demonstrated by the cutting and offset of veins, vein fragments, and color banding. The orientations and sense of shear on the microscopic shear bands are consistent with both the orientations and sense of shear expected for the Riedel array. It is possible that the presence of mesofractures is accentuated by near surface phenomena associated with uncovering and weathering. Nonetheless, the similarity of the mesoscale and microscale fabrics suggest that in the very least, the mesofractures follow planes of weakness in the ultracataclasite resulting from the microscale layering and shear bands. As such,

213

the mesoscale fracture fabric reflects the syntectonic structure of the ultracataclasite if not the geometry of actual shear surfaces. We infer that the pfs represents a late-stage, throughgoing, mesoscale principal slip surface within the ultracataclasite because (1) the pfs is present and continuous in all exposures of the ultracataclasite, (2) all contacts, layering, and fracture surfaces in the ultracataclasite either merge with or are truncated by the pfs, (3) it forms the contact between different rock types such as between the brown and black ultracataclasites or between the brown ultracataclasite and Punchbowl Formation, (4) the pfs displays a lower roughness in the slip-parallel than in the slip-perpendicular exposures, consistent with mesoscale corrugations of the surface aligned with the inferred direction of slip on the Punchbowl fault, and (5) the pfs is spatially associated with the thin layer of less cohesive ultracataclasite (Figs. 3 and 5). 6.2. Timing and processes of formation of the brown and black ultracataclasites The general structure of the fault zone, including the continuous ultracataclasite layer, must have been established fairly early in the fault history because the ultracataclasite layer has accommodated almost all of the shear displacement on the fault. Probably most of the early-formed features in the ultracataclasite layer were destroyed during subsequent deformations when the large shear strain was imposed. The formation and juxtaposition of the brown and black ultracataclasites may represent two of the more ancient events recorded in the fault core. We do not know if the black and brown ultracataclasites were derived solely from the basement and Punchbowl Formation, respectively. However, significant mixing of comminuted basement and Punchbowl Formation is not consistent with the restricted distribution of porphyroclasts. We infer that the brown ultracataclasite was largely derived from the Punchbowl Formation sandstone because (1) the brown ultracataclasite is always in contact with the sandstone, (2) the brown ultracataclasite contains only porphyroclasts of sandstone, (3) porphyroclasts of sandstone are restricted to the brown ultracataclasite, and (4) the brown ultracata-

214

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

clasite is very similar in appearance to the ultracataclasites in the subsidiary faults of the Punchbowl Formation. Similar reasoning implies that the black ultracataclasite must have originated from comminution of the basement rock. The similar texture and mineralogy of the two ultracataclasites suggests that they formed at fairly similar conditions. A simple model for the formation and juxtaposition of the brown and black ultracataclasites involves (1) generation of the black ultracataclasite layer along a section of the fault within the basement and generation of the brown ultracataclasite layer along a section of the fault within the Punchbowl Formation, followed by (2) large displacement on the pfs to place the black and brown ultracataclasites in contact (Fig. 6). The contact between the two ultracataclasites must have formed prior to the very

last increment of displacement along the pfs because the contact either merges with or is truncated by the pfs. We interpret that the less cohesive brown ultracataclasite formed during the last stage of deformation in the fault core. Cohesion of the ultracataclasite is partly a result of mineralization and the vein cementation. Almost all of the rocks in the fault core contain microscale veins that record several episodes of cementation. Veins and vein porphyroclasts were reoriented to form layers parallel to the clay foliation by offset on distributed networks of microscale shear bands and particle-scale flow in the clay matrix (Chester and Logan, 1987). The cementation events must have been syntectonic because the veins and microscopic shear bands are mutually cross cutting. The fact that the less cohesive brown ultracataclasite

Fig. 6. Schematic illustrating a simple model for the formation and juxtaposition of the dark yellowish-brown and olive-black ultracataclasites. During the early stages of faulting some segments of the fault were wholly contained in the Punchbowl Formation or in the basement. In these segments the ultracataclasite is derived from a single host rock type. At late stages of faulting, after large displacement on the fault, the Punchbowl Formation and the basement are juxtaposed. Translation of the brown and black ultracataclasites with the host rock places the ultracataclasite in contact.

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

generally lacks veins and vein fragments, suggests that it represents the most recently reworked and uncemented material of the fault core. 6.3. Mode of failure during the last phase of faulting The relations mapped indicate that all of the ultracataclasite underwent an earlier phase of deformation involving cementation and microscale shearing during which cohesion developed, and a later phase involving displacement along the pfs in the interior of the ultracataclasite layer. This later phase produced a narrow layer of reworked, less cohesive ultracataclasite, which we interpret to be the product of abrasive wear from slip on the pfs. Although the contact between the black and brown ultracataclasites could have formed through distributed shearing flow, we suggest a simple model in which both the formation of the contact and all subsequent displacement on the fault occurred by localized slip on the pfs. Such a model is consistent with the observations that the pfs (1) does not cut across and offset the contact between the brown and black ultracataclasites, and (2) either coincides with the contact or, as seen in the excavation mapped at a scale of 1 : 1, is separated from the contact by the less cohesive brown ultracataclasite. Slip on the pfs would have reworked the cohesive ultracataclasite by abrasive wear. At sites of abrasion and removal of ultracataclasite, the pfs would continue to form the contact between the two ultracataclasites, and layer-

215

ing in the ultracataclasites would be truncated by the pfs. At sites of wear product accumulation, the pfs would progressively migrate away from the contact between the two ultracataclasites as the wear product accumulated (Fig. 7). Regardless of the process that formed the contact between the two ultracataclasites, the present geometry would only be preserved if no significant mixing of the two ultracataclasites occurred since juxtaposition (Fig. 8). After the contact was formed, shear displacement on any surface that cut across the contact would have generated discontinuities and duplications of the contact. If displacement had occurred on an anastomosing set of localized slip surfaces, or on slip surfaces that branched or cut across the ultracataclasite layer, then one would expect to see interfingering and stacking of fault bounded slices of brown and black ultracataclasites rather than a single continuous contact between the two units (Fig. 8b). The geometry of the contacts between the ultracataclasite and surrounding rock and the fact that the basement and sandstone are confined to their respective sides of the ultracataclasite layer are additional evidence that shears did not cut in and out of the layer. Although the wedge structures at the boundary of the ultracataclasite layer could record some flow of the ultracataclasite into the surrounding rock, there is no evidence of mixing flow between the brown and black ultracataclasites. Features such as ultracataclasite filled cracks or injection structures that record flow or mobilization of ultracataclasite

Fig. 7. Figure illustrating wear and wear product accumulation associated with slip along the undulatory pfs. Along some segments of the surface, such as at restraining bends where high normal stress leads to high wear rates, the cohesive ultracataclasite is removed at the surface through abrasion. Along other segments, such as at releasing bends where low normal stress leads to low wear rates or separation, thin layers of abrasive wear product tends to accumulate along the surface. Wear product should accumulate as thin lenticular layers, and accumulation results in migration of the slip surface away from the contact with the cohesive ultracataclasite.

216

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

Fig. 8. Schematic showing the internal structure of the ultracataclasite layer expected for four endmembers as a function of slip localization and degree of mixing at the mesoscopic scale. (a) Slip on a single, stationary, throughgoing surface will lead to little mixing. This mode of failure results in juxtaposition of contrasting host rock and ultracataclasite across the localized slip surface. This structure is most similar to that of the Punchbowl ultracataclasite layer (see text). (b) Contemporaneous or alternating slip on multiple, anastomosing, throughgoing surfaces leads to mixing across the layer. Fault slices of host rock and ultracataclasite may be distributed and juxtaposed throughout the layer. For the case of closely spaced surfaces and large slip, the layer may appear homogeneous at the mesoscopic scale. (c) Ultracataclasite components may be juxtaposed but remain segregated by distributed flow within the layer if streamline or laminar flow is maintained. (d) Components may mix if flow is turbulent or non-laminar. Fluidization and injection of ultracataclasite is an extreme example of non-laminar flow that could lead to mixing of components. As in the case (b), flow processes could homogenize the ultracataclasite.

(e.g., Brock and Engelder, 1977; Lin, 1996) also are not observed (Fig. 8d). Some distributed fracturing and flow of the ultracataclasite and host rock would have been necessary to accommodate movement along the nonplanar pfs. However, it appears that such deformation was minor during the final phase of faulting. Thus, only a small fraction of the ultracataclasite layer was actively undergoing shear, and most of the ultracataclasite retained the cohesion and microfabric formed during earlier deformation events. The amount of displacement on the fault since formation of the contact between the brown and black

ultracataclasites must be at least equal to the length of the contact measured parallel to the slip direction. Field relations suggest that this length is greater than 1.5 km. A similar estimate is derived from the outcrop pattern of the Punchbowl Formation. The Punchbowl Formation extends along the north side of the Punchbowl fault for approximately 10 km to the southeast of the study location and probably to 2 km in depth (Fig. 1). The outcrop pattern and the right-lateral, reverse-oblique slip on the Punchbowl fault (Chester and Logan, 1987) suggest that the sandstone and basement were in contact for at least the final 2 km of slip along the section of the fault mapped.

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

7. Friction and earthquake mechanisms The mesoscopic structure of the Punchbowl fault ultracataclasite layer records the mode of failure during the final several kilometers of slip on a large displacement fault of the San Andreas system. Assuming that the fault slipped seismically, the final several kilometers of slip would have generated up to 103 large magnitude earthquake slip events, or an even greater number of smaller magnitude events. Thus, the structure of the ultracataclasite layer reflects deformation associated with the passage of numerous earthquake ruptures and implies that: (1) during the later stages of faulting, repeated earthquake ruptures occurred along the same well defined pfs; (2) slip along this pfs generated a frictional wear product from a pre-existing, layered, cohesive ultracataclasite; and (3) newly formed surfaces and splays did not accommodate a significant amount of slip. To date, few features have been identified that are unique to seismic slip or aseismic creep. Pseudotachylytes are the best evidence for paleoseismic slip events (e.g., Sibson et al., 1975; Swanson, 1992). Other features that may indicate seismic slip are implosion breccias, fluidized cataclasite, and injection structures. These latter structures are thought to result from sudden pressurization or depressurization processes (Scholz, 1990; Lin, 1996). At the other extreme, crystal fiber growth on a slip surface, cataclastic foliations, and layering often are attributed to slow, aseismic creep processes (Sibson, 1986; Groshong, 1988; Miller, 1996). However, because creep compaction may be the typical mode of failure during interseismic periods of the seismic cycle (Sibson, 1989; Chester et al., 1993), the presence of features formed by creep cannot be used as evidence precluding seismic slip. The existence of structures recording both localized slip and distributed flow may indicate repetitive deformation sequences characteristic of seismic cycling (e.g., Hobbs et al., 1986; Power and Tullis, 1989). Except for foliation and layering, none of the features diagnostic of seismic or aseismic slip are present in the Punchbowl fault zone. Some characteristics of internal fault structure associated with unstable slip have been suggested by rock deformation experiments investigating fault friction. Faults often are modeled in experiments by

217

sliding two blocks of intact rock along a contacting surface or a thin layer of simulated gouge (e.g., Logan, 1975; Beeler et al., 1996). In simulated gouge experiments, the imposed displacement initially is accommodated in the layer by distributed slip on a network of Riedel shears (Logan et al., 1979). With greater displacement, localized shears parallel to the layer develop within and at the boundaries of the layer and subsequently accommodate the imposed displacements (Engelder et al., 1975; Logan et al., 1979, 1992). Abrasive wear along a localized shear surface produces extremely fine-grained material that tends to accumulate along irregularities of the surface. At relatively large displacements, as achieved in rotary shear experiments on simulated gouge (e.g., Beeler et al., 1996), the mode of failure is quite similar to that in the Punchbowl ultracataclasite layer. That is, the displacement in an experiment is accommodated along distinct layer-parallel shear surfaces, and apparently only a narrow zone within the layer is undergoing active shear at any one time (Marone and Kilgore, 1993; Beeler et al., 1996). Fine-grained wear product accumulates as the shear surface slowly migrates through the layer. A process that weakens a fault is necessary for a dynamic earthquake rupture to nucleate (e.g., Rice, 1983). Experiments suggest that the extreme localization of slip to discrete surfaces, such as occurs at large displacements in simulated gouge experiments, favors the small critical slip distance and rate weakening friction behavior that promotes dynamic instability (Marone and Kilgore, 1993; Beeler et al., 1996). Thus, the localization of displacement in the Punchbowl fault to the pfs associated with narrow zones of less cohesive ultracataclasite would appear consistent with rate weakening behavior and seismic slip. Theoretical models suggest that both the nucleation patch size of an earthquake rupture and the dimensions of the breakdown zone at the earthquake rupture tip scale with the critical slip distance (Rudnicki, 1980; Dieterich, 1992). The breakdown process at the rupture tip involves distributed fracture and fracture linkage to form the rupture surface. Assuming that the zone of distributed fracture, occurrence of discontinuities, and roughness of the rupture surface increase with the dimension of the breakdown zone, we infer small nucleation and breakdown dimensions for ruptures in the Punchbowl fault

218

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

zone. The preservation of preexisting structures near the pfs, such as the contact between the black and brown ultracataclasite, suggests that the breakdown processes did not significantly disrupt a large volume of rock. Repeated seismic rupture on a single slip surface in the Punchbowl ultracataclasite layer is opposite to the kinematics of pseudotachylyte bearing faults where slip surfaces are rarely reruptured (e.g., Grocott, 1981; Swanson, 1988, 1989, 1990). Pseudotachylyte bearing faults are effectively welded upon solidification of the friction melts and have strengths that probably approach or exceed the host rock. Thus, subsequent ruptures tend to form in the host rock. Sidewall ripouts have been described for pseudotachylyte bearing faults, and represent a mechanism for bypassing a segment of a fault that seizes during slip (Swanson, 1989). Although strength recovery also must have occurred along the slip surface of the Punchbowl fault, the repeated rupture of the pfs and the absence of sidewall ripouts implies that the strength contrast between the slip surface and wall rock remained pronounced. Processes that lead to weakening during dynamic slip may be important for the occurrence of great earthquakes (e.g., Brune, 1991). In addition, such dynamic weakening mechanisms may be necessary to explain the low shear strength of the San Andreas fault. A number of processes have been proposed that could lead to dynamic weakening including thermal pressurization of pore fluids (e.g., Sibson, 1973), liquefaction, fluidization of granular material (e.g., Melosh, 1996), and reduction of normal stress by interface separation during dynamic slip (Brune et al., 1993). Frictional heating may lead to significant temperature changes in cases where rapid, large slip occurs in narrow zones (Brune et al., 1969; Mase and Smith, 1987). Under extreme conditions of seismic slip, temperature changes can be great and cause local melting of rock. For seismic faulting in fluid saturated rock, heating should cause a local increase in pore fluid volume (e.g., Sibson, 1973). In conditions where permeability and changes in porosity are small, heating can lead to pressurization of fluid, a local reduction in effective stress, and dynamic weakening (Mase and Smith, 1987). Changes in porosity and permeability in the Punchbowl ultracataclasite layer during the last phase of faulting

were probably small, as most of the ultracataclasite retained its cemented and cohesive character. Large shear on surfaces or within thin layers of less cohesive ultracataclasite sandwiched between cohesive ultracataclasite, as observed along the Punchbowl fault, could be consistent with thermal pressurization during seismic slip. Such a structure also may be consistent with interface separation during seismic slip. Acoustic fluidization can occur in cohesionless, granular material with sufficient concentration of high frequency acoustic energy (Melosh, 1979, 1996). Fluidization also could occur in a fluid saturated granular material through liquefaction. For the case of seismic faulting, acoustic fluidization or liquefaction in a fault core should produce structures recording flow within the layer. Flow structures generated since the juxtaposition of the black and brown ultracataclasites generally are absent in the Punchbowl ultracataclasite layer. In fact, deformation involving distributed flow of granular material as occurs during liquefaction, fluidization, Coulomb plasticity of thick gouge zones (Byerlee and Savage, 1992), and rolling on space filling bearings with compatible kinematic rotations (Herrmann et al., 1990), probably was not important during the final several kilometers of slip on the Punchbowl fault.

8. Conclusions In the Devil’s Punchbowl area, the Punchbowl fault contains a single, continuous ultracataclasite layer centrally located in a broader zone of extremely fractured rock. The ultracataclasite layer constitutes a macroscopic scale principal slip surface that accommodated 10 km or more of slip and juxtaposed the Punchbowl Formation lithic sandstone and igneous and metamorphic rocks of the San Gabriel basement complex. We characterized the mesoscopic structure of the ultracataclasite by mapping two exposures at a scale of 1 : 10 and an excavation at 1 : 1. Different types of ultracataclasite were distinguished on the basis of color, cohesion, fracture and vein fabric, and porphyroclast lithology. The ultracataclasite is internally layered, and the geometry and crosscutting relations of layer contacts record the mode

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

of failure during the final stages of fault activity. By analogy with modern faults of the San Andreas in the Transverse Ranges, we assume that the Punchbowl fault slipped seismically, and that the structure of the ultracataclasite layer probably records deformation resulting from the passage of numerous earthquake ruptures. Structural relations indicate that the ultracataclasite underwent an earlier phase of deformation involving cementation and microscale flow during which cohesion developed, and a later phase involving displacement along a prominent fracture surface (pfs) in the interior of the ultracataclasite layer and production of a thin layer of reworked, less cohesive ultracataclasite. Prior to the latest stage of slip along the pfs, a brown ultracataclasite and black ultracataclasite associated with the Punchbowl Formation and the basement, respectively, were juxtaposed. The amount of displacement on the fault since formation of the contact between the brown and black ultracataclasites is greater than 1.5 km. We infer that the contact between the brown and black ultracataclasites was formed by localized slip on the pfs. With subsequent slip, attrition wear and wear product accumulation occurred along the pfs. At sites of attrition, the layering in the ultracataclasite was truncated at the pfs and the contact between the brown and black ultracataclasite coincides with the pfs. At sites of wear product accumulation, the pfs is separated from the contact between the brown and black ultracataclasites by a layer of less cohesive ultracataclasite that represents the accumulation of reworked ultracataclasite. There is no evidence for significant mixing of the brown and black ultracataclasites as would occur by offset on anastomosing shear surfaces in the layer or by mobilization and injection of one ultracataclasite into the other. The internal structure of the Punchbowl fault implies that earthquake ruptures were not only confined to the ultracataclasite layer, but largely localized to the pfs. During the latest phase of deformation, the bulk of the ultracataclasite retained cohesion and layer contacts were relatively undisturbed. As such, the earthquake ruptures must have followed the pfs without significant branching or jumping to other locations in the ultracataclasite. By comparison with laboratory studies of rock friction, the localization of displacement in the Punchbowl ultracataclasite im-

219

plies rate weakening behavior with small critical slip distance, and thus small nucleation and breakdown dimensions for ruptures. Of the various mechanisms proposed to explain of the low strength of the San Andreas and to produce dynamic weakening of faults, those that require or predict wide zones of less cohesive granular material that flows and mixes appear incompatible with our observations. Mechanisms that assume or are consistent with extreme localization of slip, such as thermal pressurization of pore fluids and possibly interface separation waves, are more consistent with structures produced during the final phase of movement on the Punchbowl fault.

Acknowledgements It was under John Logan’s guidance that we both began our research on crustal faulting. His influence on our research is obvious, particularly regarding the importance we place on combining field and experimental approaches to geologic problems. We are pleased to dedicate this paper to John, and thank him for his support and friendship over the years. Discussions with J.P. Evans were helpful. Reviews by Sue Agar and Mark Swanson motivated additional study and helped us to clarify our thinking. This research was supported by the U.S. Geological Survey (USGS), Department of the Interior, under awards 1434-94-G-2457 and 1434-HQ-96-GR02709, and by a National Science Foundation (NSF) award EAR92-05973.

References An, L.J., Sammis, C.G., 1994. Particle size distribution of cataclastic fault materials from southern California: A 3-D study. Pure Appl. Geophys. 143, 203–227. Arboleya, M.L., Engelder, T., 1995. Concentrated slip zones with subsidiary shears: Their development on three scales in the Cerro Brass fault zone, Appalacian valley and ridge. J. Struct. Geol. 17, 519–532. Barrows, A.G., Kahle, J.E., Beeby, D.J., 1985. Earthquake hazards and tectonic history of the San Andreas fault zone, Los Angeles County, California. Calif. Div. Mines Geol. Open File Rep. 85-10LA, 139. Barrows, A.G., Kahle, J.E., Beeby, D.J., 1987. Earthquake hazards and tectonic history of the San Andreas fault zone, Los

220

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221

Angeles County, California. In: Hester, R.L., Hallinger, D.E. (Eds.), San Andreas fault-Cajon Pass to Palmdale. Pac. Sect. Am. Assoc. Pet. Geol. Vol. Guideb. 59, 1–92. Beeler, N.M., Tullis, T.E., Blanpied, M.L., Weeks, J.D., 1996. Frictional behavior of large displacement experimental faults. J. Geophys. Res. 101, 8697–8715. Brock, W.G., Engelder, T., 1977. Deformation associated with the movement of the Muddy Mountain Overthrust in the Buffington Window, southeastern Nevada. Geol. Soc. Am. Bull. 88, 1667–1677. Bruhn, R.L., Yonkee, W.A., Parry, W.T., 1990. Structural and fluid-chemical properties of seismogenic normal faults. Tectonophysics 218, 139–175. Brune, J.N., 1991. Seismic source dynamics, radiation and stress. Rev. Geophys. Suppl. 29, 688–699. Brune, J.N., Henyey, T.L., Roy, R.F., 1969. Heat flow, stress and rate of slip along the San Andreas fault, California. J. Geophys. Res. 74, 3821–3827. Brune, J.N., Brown, S., Johnson, P.A., 1993. Rupture mechanisms and interface separation in foam rubber models of earthquakes — a possible solution to the heat-flow paradox and the paradox of large overthrusts. Tectonophysics 218, 59– 67. Byerlee, J., 1990. Friction, overpressure and fault normal compression. Geophys. Res. Lett. 17, 2109–2112. Byerlee, J., 1993. Model for episodic flow of high pressure water in fault zones before earthquakes. Geology 21, 303–306. Byerlee, J.D., Savage, J.C., 1992. Coulomb plasticity within the fault zone. Geophys. Res. Lett. 19, 2341–2344. Chester, F.M., Logan, J.M., 1986. Implications for mechanical properties of brittle faults from observations of the Punchbowl fault zone, California. Pure Appl. Geophys. 124, 79–106. Chester, F.M., Logan, J.M., 1987. Composite planar fabric of gouge from the Punchbowl fault, California. J. Struct. Geol. 9, 621–634. Chester, F.M., Evans, J.P., Biegel, R.L., 1993. Internal structure and weakening mechanisms of the San Andreas fault. J. Geophys. Res. 98, 771–786. Cowie, P.A., Scholz, C.H., 1992. Physical explanation for the displacement–length relationships of faults using a post-yield fracture mechanics model. J. Struct. Geol. 14, 1133–1148. Cox, B.F., Powell, R.E., Hinkle, M.E., Lipton, D.A., 1983. Mineral resource potential map of the Pleasant View Roadless Area, Los Angeles County, California. U.S. Geol. Surv. Misc. Field Stud. Map MF-1649-A, scale 1 : 62,500. Cox, S.J.D., Scholz, C.H., 1988. On the formation and growth of faults: an experimental study. J. Struct. Geol. 10, 413–430. Dibblee, Jr., T.W., 1967. Areal geology of the western Mojave desert. U.S. Geol. Surv. Prof. Pap. 522. Dibblee, T.W., Jr., 1968. Displacements on the San Andreas fault system in the San Gabriel, San Bernardino, and San Jacinto Mountains, southern California. Proc. Conf. on Geological Problems of the San Andreas fault system. Stanford Univ. Publ. Geol. Sci. 11, 260–276. Dieterich, J.H., 1992. Earthquake nucleation on faults with rateand state-dependent strength. Tectonophysics 211, 115–134. Ehlig, P.L., 1975. Basement rocks of the San Gabriel Mountains,

south of the San Andreas fault, southern California. Calif. Div. Mines Geol. Spec. Rep. 118, 177–186. Engelder, J.T., Logan, J.M., Handin, J., 1975. The sliding characteristics of sandstone on quartz fault-gouge. Pure Appl. Geophys. 113, 69–86. Engelder, T., 1974. Microscopic wear grooves on slickensides: indicators of paleoseismicity. J. Geophys. Res. 79, 4387. Evans, J.P., Chester, F.M., 1995. Fluid–rock interaction in faults of the San Andreas system: Inferences from San Gabriel fault rock geochemistry and microstructures. J. Geophys. Res. 100, 13007–13020. Flinn, D., 1977. Transcurrent faults and associated cataclasis in Shetland. J. Geol. Soc. London 133, 231–248. Gay, N.C., Ortlepp, W.D., 1979. Anatomy of a mining-induced fault zone. Geol. Soc. Am. Bull. 90, 47–58. Grocott, J., 1981. Fracture geometry of pseudotachylyte generation zones: A study of shear fractures formed during seismic events. J. Struct. Geol. 3, 169–178. Groshong Jr., R.H., 1988. Low-temperature deformation mechanisms and their interpretation. Geol. Soc. Am. Bull. 100, 1329–1360. Herrmann, H.J., Mantica, G., Bessis, D., 1990. Space-filling bearings. Phys. Rev. Lett. 65, 3223–3226. Hickman, S.H., 1991. Stress in the lithosphere and the strength of active faults. Rev. Geophys. Suppl. 29, 759–775. Hobbs, B.E., Ord, A., Teyssier, C., 1986. Earthquakes in the ductile regime? Pure Appl. Geophys. 124, 309–336. House, W.M., Gray, D.R., 1982. Cataclasites along the Saltville thrust, U.S.A., their implications for thrust-sheet emplacement. J. Struct. Geol. 4, 257–269. Hull, J., 1988. Thickness-displacement relationships for deformation zones. J. Struct. Geol. 10, 431–435. Lin, A., 1996. Injection veins of crushing-originated pseudotachylyte and fault gouge formed during seismic faulting. Eng. Geol. 43, 213–224. Little, T.A., 1995. Brittle deformation adjacent to the Awatere strike-slip fault in New Zealand: Faulting patterns, scaling relationships, and displacement partitioning. Geol. Soc. Am. Bull. 107, 1255–1271. Logan, J.M., 1975. Friction in rocks. Rev. Geophys. Space Phys. 13, 358–361. Logan, J.M., Friedman, M., Higgs, N.G., Dengo, C., Shimamoto, T., 1979. Experimental studies of simulated gouge and their application to studies of natural fault gouge, Proc. Conf. VIII, Analysis of Actual Fault Zones in Bedrock. U.S. Geol. Surv. Open-file Rep., pp. 305–340. Logan, J.M., Dengo, C.A., Higgs, N.G., Wang, Z.Z., 1992. Fabrics of experimental fault zones: Their development and relationship to mechanical behavior. In: Evans, B., Wong, T. (Eds.), Fault Mechanics and Transport Properties of Rocks. Academic Press, London, pp. 33–68. Marone, C., Kilgore, B., 1993. Scaling of the critical slip distance for seismic faulting with shear strain in fault zones. Nature 362, 618–621. Mase, C.W., Smith, L., 1987. Effects of frictional heating on the thermal, hydrologic, and mechanical response of a fault. J. Geophys. Res. 92, 6249–6272.

F.M. Chester, J.S. Chester / Tectonophysics 295 (1998) 199–221 Matti, J.C., Morton, D.M., 1993. Paleogeographic evolution of the San Andreas fault in southern California: A reconstruction based on a new cross-fault correlation. In: Powell, R.E., Weldon II, R.J., Matti, J.C. (Eds.), The San Andreas Fault System: Displacement, Palinspastic Reconstruction, and Geologic Evolution. Geol. Soc. Am. Mem. 178, 107–160. Matti, J.C., Morton, D.M., Cox, B.F., 1985. Distribution and geologic relations of fault systems in the vicinity of the central Transverse Ranges, southern California, U.S. Geol. Surv. Open-file Rep. 85-365. Means, W.D., 1995. Shear zones and rock history. Tectonophysics 247, 157–160. Melosh, H.J., 1979. Acoustic fluidization: a new geologic process? J. Geophys. Res. 84, 7513–7520. Melosh, H.J., 1996. Dynamical weakening of faults by acoustic fluidization. Nature 379, 601–606. Miller, M.G., 1996. Ductility in fault gouge from a normal fault system, Death Valley, California: A mechanism for fault-zone strengthening and relevance to paleoseismicity. Geology 24, 603–606. Morton, D.M., Matti, J.C., 1987. The Cucamonga fault zone, geologic setting and Quaternary history. U.S. Geol. Surv. Prof. Pap. 1339. Noble, L.F., 1954. Geology of the Valyermo Quadrangle and vicinity. U.S. Geol. Surv. Map GQ-50. Oakeshott, G.B., 1971. Geology of the epicentral area. Calif. Div. Mines Geol. 196, 19–30. Power, W.L., Tullis, T.E., 1989. The relationship between slickenside surfaces in fine-grained quartz and the seismic cycle. J. Struct. Geol. 11, 879–894. Power, W.L., Tullis, T.E., 1991. Euclidean and fractal models for the description of rock surface roughness. J. Geophys. Res. 96, 415–424. Reches, Z., Lockner, D.A., 1994. Nucleation and growth of faults in brittle rocks. J. Geophys. Res. 99, 18159–18173. Rice, J.R., 1983. Constitutive relations for fault slip and earthquake instabilities. Pure Appl. Geophys. 121, 443–475. Rice, J.R., 1992. Fault stress states, pore pressure distributions, and the weakness of the San Andreas fault. In: Evans, B., Wong, T.F. (Eds.), Fault Mechanics and Transport Properties in Rocks. Academic Press, London, pp. 475–503. Rudnicki, J.W., 1980. Fracture mechanics applied to the earth’s crust. Annu. Rev. Earth Planet. Sci. 8, 489–525. Rutter, E.H., Maddock, R.H., Hall, S.H., White, S.H., 1986. Comparative microstructures of natural and experimentally produced clay-bearing fault gouges. Pure Appl. Geophys. 124, 3–30. Scholz, C.H., 1990. The Mechanics of Earthquakes and Faulting. Cambridge Univ. Press, New York, NY, 439 pp. Scholz, C.H., Dawers, N.H., Yu, J.Z., Anders, M.H., Cowie, P.A.,

221

1993. Fault growth and fault scaling laws: Preliminary results. J. Geophys. Res. 98, 21951–21961. Schulz, S.E., Evans, J.P., 1998. Spatial variability in microscopic deformation and composition of the Punchbowl fault, southern California: Implications for mechanisms, fluid–rock interaction, and fault morphology. Tectonophysics 295, 225–246. Segall, P., 1991. Fault mechanics. Suppl. Rev. Geophys. 29, 864– 876. Sibson, R.H., 1973. Interactions between temperature and fluid pressure during earthquake faulting, a mechanism for partial or total stress relief. Nature 243, 66–68. Sibson, R.H., 1977. Fault rocks and fault mechanisms. J. Geol. Soc. London 133, 191–213. Sibson, R.H., 1986. Brecciation processes in fault zone: Inferences from earthquake rupturing. Pure Appl. Geophys. 124, 159–175. Sibson, R.H., 1989. Earthquake faulting as a structural process. J. Struct. Geol. 11, 1–14. Sibson, R.H., Moore, J.M., Rankin, A.H., 1975. Seismic pumping — a hydrothermal fluid transport mechanism. J. Geol. Soc. London 131, 653–659. Sleep, N., Blanpied, M.L., 1992. Creep, compaction, and the weak rheology of major faults. Nature 359, 687–692. Swanson, M.T., 1988. Pseudotachylyte-bearing strike-slip duplex structures in the Fort Foster Brittle Zone, S. Maine. J. Struct. Geol. 10, 813–828. Swanson, M.T., 1989. Sidewall ripouts in strike-slip faults. J. Struct. Geol. 11, 933–948. Swanson, M.T., 1990. Extensional duplexing in the York Cliffs strike-slip fault system, southern coastal Maine. J. Struct. Geol. 12, 499–512. Swanson, M.T., 1992. Fault structure, wear mechanisms and rupture processes in pseudotachylyte generation. Tectonophysics 204, 223–242. Wallace, R.E., Morris, H.T., 1986. Characteristics of faults and shear zones in deep mines. Pure Appl. Geophys. 124, 107– 125. Weldon II, R.J., Meisling, K.E., Alexander, J., 1993. A speculative history of the San Andreas fault in the central Transverse Ranges, California. In: Powell, R.E., Weldon II, R.J., Matti, J.C. (Eds.), The San Andreas Fault System: Displacement, Palinspastic Reconstruction, and Geologic Evolution. Geol. Soc. Am. Mem. 178, 161–198. Wojtal, S., Mitra, G., 1988. Nature of deformation in some fault rocks from Appalachian thrusts. Geol. Soc. Am. Spec. Pap. 222, 17–33. Woodburne, M.O., 1975. Cenozoic stratigraphy of the Transverse Ranges and adjacent areas, southern California. Geol. Soc. Am. Spec. Pap. 162, 91.